The Precambrian Earth: tempos and events

Developments in Precambrian Geology, 12 THE PRECAMBRIAN EARTH: TEMPOS AND EVENTS Developments in Precambrian Geology,...

Author: P. G. Eriksson

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Developments in Precambrian Geology, 12

THE PRECAMBRIAN EARTH: TEMPOS AND EVENTS

Developments in Precambrian Geology, 12

THE PRECAMBRIAN EARTH: TEMPOS AND EVENTS

DEVELOPMENTS IN PRECAMBRIAN GEOLOGY Advisory Editor Kent Condie

Further titles in this series 1. B.F.WlNDLEY and S.M. NAQVI (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.MEI'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. NAQVl (Editor) Precambrian Continental Crust and Its Economic Resources 9. D.V. RUNDQVIST and F.P. MITROFANOV (Editors) Precambrian Geology of the USSR 10. K.C. CONDIE (Editor) Proterozoic Crustal Evolution 11. K.C. CONDIE (Editor) Archean Crustal Evolution

Developments in Precambrian Geology, 12

THE PRECAMBRIAN EARTH: TEMPOS AND EVENTS Edited by P.G. E R I K S S O N Department of Geology, University of Pretoria Pretoria 0002, Republic of South Africa

W. A L T E R M A N N Centre Biophysique Mo16culaire (CBM) Centre National de la Recherche Scientifique (CNRS) 45071 Orl6ans, Cedex 2, France

D.R. NELSON Department of Applied Physics, Curtin University of Tectmology Perth, W.A. 6845, Australia and Geological Survey of Western Australia, Mineral House, 100 Plain Street East Perth, 6004, Australia

W.U. MUELLER Departement Sciences de la Terre Universit6 du Qu6bec ~tChicoutimi Chicoutimi, Qu6bec G7H 2B 1, Canada

O. CATUNEANU Department of Earth and Atmospheric Sciences University of Alberta, Edmonton Alberta T6G 2E3, Canada

2004

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CONTRIBUTING AUTHORS D.H. ABBOTT

Lamont-Doherty Earth Observatory, University of Columbia, Palisades, New York, NY 10964, USA (dallass@ ldeo.columbia.edu; phone: 845-3658664;fax: 845-3658156) F.E ALKMIM

Departamento de Geologia, Universidade Federal de Ouro Preto, Ouro Preto, MG-Brazil, CP 35400 ([email protected]; fax: 31-5512334) W. ALTERMANN

Centre Biophysique Moleculaire, Exobiologie, Centre National de la Recherche Scientifique, Rue CharlesSadron, 45071 Orleans Cedex 2, France ([email protected]; phone: 33-238-255569; fax: 33-238631517) EW.U. APPEL

Geological Survey of Denmark and Greenland, r geus.dk; phone: 45-38-142214; fax: 45-38-142050)

Voldgade 10, DK-1350 Copenhagen, Denmark (pa@

N. ARNDT

LGCA, Universitd de Grenoble, 38041 Grenoble Cedex, France ([email protected]; phone: 33-4-76635931; fax: 33-4-76514058) L.B. ASPLER

23 Newton Street, Ottawa, Ontario K1S 2S6, Canada ([email protected]; phone: 908-6470180x7152; fax: 908-5803523) L.D. AYRES

Department Geological Sciences, University of Manitoba, Winnipeg, Manitoba R3T 2N2, Canada (ayres@ms. umanitoba, ca; phone: 204-4749371; fax: 204-4747623) S. BANERJEE

Dept. of Earth Sciences, liT Bombay, Pawai, Mumbai 400 076, India ([email protected]; phone: 22-5767282; fax: 5767253) J.G. BLOCKLEY

76 Beach Street, Bicton, WA 6157, Australia ([email protected]; phone: 8-93171775) P.K. BOSE

Dept. of Geological Sciences, Jadavpur University, Kolkata 700 032, India ([email protected]; fax: 33-4731484) M.D. BRASIER

Earth Sciences Department, Oxford University, Parks Road, Oxford OX1 3PR, United Kingdom (Martin. [email protected]; phone: 1865-272074;fax: 1865-272072) K.L. BUCHAN

Geological Survey of Canada, 601 Booth St., Ottawa, Ontario KIA OE8, Canada ([email protected]; phone: 613-9477341;fax: 613-9477396)

Contributing authors

vi

A.J. BUMB Y

Dept. of Geology, University of Pretoria, Pretoria 0002, South Africa ([email protected]; phone: 27-12-4202238; fax: 27-12-3625219) G.R. BYERLY

Dept. of Geology and Geophysics, Louisiana State University, Baton Rouge, LA 70803-4101, USA (gbyerly@ geol.lsu.edu; phone: 225-5785318; fax: 225-5782300) O. CATUNEANU

Dept. of Earth and Atmospheric Sciences, 1-26 Earth Sciences Building, University of Alberta, Edmonton, Alberta T6G 2E3, Canada ([email protected]; phone: 780-4926569; fax: 780-4922030) J.R. CHIARENZELLI

Dept. of Geology, State University of New York, Potsdam, NY 13676, USA ([email protected]; phone: 315-2673401; fax: 315-2672695) E.H. CHOWN

90 Dickens Drive, Kingston, Ontario K7M 2M8, Canada ([email protected]; phone: 1-613-5475632) K.C. CONDIE

Dept. of Earth and Environmental Sciences, New Mexico Tech., Socorro, NM 87801, USA ([email protected]; phone: 505-8355531; fax: 505-8356436) P.L. CORCORAN

Dept. of Earth Sciences, University of Western Ontario, London, Ontario N6A 5B7, Canada (pcorcor@ uwo.ca; phone: 519-6612111x86836; fax: 519-6613198) B.L. COUSENS

Dept. of Earth Sciences, Carleton University, 1125 Colonel By Drive, Ottawa, Ontario KIS 5B6, Canada (brian_cousens@ carleton.ca; phone: 613-5202600x4436; fax: 613-5202569) R. DAIGNEAULT

Sciences de la terre, Universit~ du Quebec ~ Chicoutimi, 555 Blvd. De l'Universite, Chicoutimi, Quebec G7H 2B1, Canada ([email protected]; phone: 418-5455011x5636; fax: 418-5455012) J.R. DEVANEY

Petro-Canada Oil and Gas, P.O. Box 2844, Calgary, Alberta T2P 3E3, Canada (ion.devaney @backpacker,corn; phone: 403-2964243; fax: 403-2963030) S.T. DE VRIES

Department of Sedimentology, Institute of Earth Sciences, Utrecht University, Postbus 80021, 3508 TA Utrecht, The Netherlands ([email protected]) J.A. DONALDSON

Ottawa-Carleton Geoscience Centre, Department of Earth Sciences, Carleton University, Colonel By Drive, Ottawa, Ontario K1S 5B6, Canada ([email protected]; phone: 613-5203515; fax: 613-5202569) J. DOSTAL

Department of Geology, Saint Mary's University, Halifax B3H 3C3, Canada ([email protected]; phone: 902-4205747; fax: 902-4968104) A.F. EMBRY

Geological Survey of Canada, 3303 33rd Street N.W., Calgary, Alberta T2L 2A7, Canada (AEmbry@ NRCan.gc.ca; phone: 403-2927125; fax: 403-2925377)

Contributing authors

vii

K.A. ERIKSSON

Dept. of Geological Sciences, Virginia Tech., Blacksburg, VA 24061, USA ([email protected]; phone: 540-2316521; fax: 540-2313386) P.G. ERIKSSON

Dept. of Geology, University of Pretoria, Pretoria 0002, South Africa ([email protected]; phone: 12-4202238; fax: 12-3625219) R.E. ERNST

Geological Survey of Canada, 601 Booth St., Ottawa, Ontario KIA OE8, Canada ([email protected]; phone: 613-9477341; fax: 613-9477396) A.D. FOWLER

Department of Earth Sciences, University of Ottawa, Ontario KIN 6N5, Canada ([email protected]; phone: 613-5625800x6273; fax: 613-5625192) H.E. FRIMMEL

Dept. of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa ([email protected]. za; phone: 21-6502901; fax: 21-6503783) A.M. GELLATLY

Dept. of Geological Sciences, University of Missouri, Columbia, MO 65211, USA ([email protected]; phone: 573-8826328; fax: 573-8825458) J.T. HAGSTRUM

U.S. Geological Survey, 345 Middlefield Road, MS 937, Menlo Park, CA 94025, USA ([email protected]; phone: 650-3294672; fax: 650-3294664) A.H. HICKMAN

Geological Survey of Western Australia, Department of Industry and Resources, Mineral House, 100 Plain Street, East Perth, WA 6004, Australia ([email protected]; phone: 9-2223333; fax: 9-2223633) A.W. HOFMANN

Max-Planck-Institut fur Chemie, Abteilung Geochemie, 55020 Mainz, Germany ([email protected]) G. HUDAK

Department of Geology, University of Wisconsin Oshkosh, W154901-8649, USA ([email protected]; phone: 920-4244463; fax: 920-4240240) L.C. KAH

Dept. of Geological Sciences, University of Tennessee, Knoxville, TN 37996, USA ([email protected]; phone: 865-9742366; fax: 865-9742368) J. KAZMIERCZAK

Institute of Paleobiology, Biogeology Division, Polish Academy of Sciences, Twarda 51/55, PL-00818 Warsaw, Poland ([email protected]; phone: +48-22-697887;fax: +48-22-6206225) S. KEMPE

Institute of Applied Geosciences, University of Technology Darmstadt, Schnittspahnstr. 9, D-64287 Darmstadt, Germany (kempe @geo.tu-darmstadt.de; phone: +49-6151-162471; fax: +49-6151-166539) A.N. KONILOV

Laboratory of the Early Precambrian Tectonics, Geological Institute of the Russian Academy of Sciences, Pyzhevsky Street 7, 109017 Moscow, Russia ([email protected]; phone: 095-2308346; fax: 095-9510443)

Contributing authors

viii

J.E LINDSAY

JSC Astrobiology Institute, NASA,JSC (SA-13), Houston, Texas, TX 77058, USA (john.f [email protected]; phone: 281-2445119; fax: 281-4831573) D.G.E LONG

Department of Earth Sciences, Laurentian University, Sudbury, Ontario P3E 2C6, Canada (dlong@nickel. laurentian.ca; phone: 705-6751151x2268;fax: 705-6754898) D.R. LOWE

Geological and Environmental Sciences Dept., Stanford University, Stanford, CA 94305-2115, USA (dlowe@ pangea.stanford.edu; phone: 650-7253040;fax: 650-7250979) T.W. LYONS

Dept. of Geological Sciences, University of Missouri, 101 Geological Sciences Building, Columbia, MO 65211, USA ([email protected]; phone: 573-8826328; fax: 573-8825458) M.A. MARTINS-NETO

Departamento de Geologia, Universidade Federal de Ouro Preto, Caixa Postal 173, 35400-000 Ouro Preto/MG, Brazil ([email protected]; fax: 31-5512334) NUPETRO N~cleo de Geologia do Petr61eo/ Fundafao Gorceix, Ouro Preto, Caixa Postal 173, 35400-000, Ouro Preto/TVIG, Brazil ([email protected]) -

J.G. MEERT

Dept. of Geological Sciences, 274 Williamson Hall University of Florida, Gainesville, FL 32611, USA ([email protected]; phone: 352-8462414; fax: 352-3929294) L.T. MIDDLETON

Northern Arizona University, Department of Geology, P.O. Box 4099, Building 12, Rm 100, Flagstaff, AZ 86011-4099, USA ([email protected]; phone: 928-5234561; fax: 928-5239220) M.V. MINTS

Laboratory of the Early Precambrian Tectonics, Geological Institute of the Russian Academy of Sciences, Pyzhevsky Street 7, 109017 Moscow, Russia ([email protected]; phone: 095-2308346; fax: 095-9510443) E.M. MOORES

Dept. of Geology, University of California, Davis, CA 95616, USA ([email protected]; phone: 530,7520352; fax: 530-7520951) P. MOSTERT

Dept. of Geology, University of Pretoria, Pretoria 0002, South Africa ([email protected]; phone: 12'4202238;fax: 12,3625219) W.U. MUELLER

Sciences de la terre, Universit~ du Qudbec ~ ChicoutimL 555 Blvd. De l'Universite, Chicoutimi, Quebec G7H 2B1, Canada (wmueller@uqac,ca; phone: 418,5455013; fax: 418-5455012) J.S. MYERS

Department of Earth Sciences, Memorial University, St John's, Newfoundland AIB 3E5, Canada (jmyers@ esd.mun.ca; phone: 709-7378000;fax: 709-7374569) D.R. NELSON

Geological Survey of Western Australia, Mineral House, 100 Plain Street, East Perth, WA 6004, Australia. Department of Applied Physics, Curtin University of Technology, GPO Box U1987, Perth, WA 6001, Australia ([email protected],au; phone: 89-2663736;fax: 89-2662377)

Contributing authors

ix

H.W. NESBITT

Dept. of Earth Sciences, University of Western Ontario, London, Ontario N6A 5B7, Canada ([email protected]; phone: 519-6613194; fax: 519-6613198) W. NIJMAN

Department of Sedimentology, Institute of Earth Sciences, Utrecht University, Postbus 80021, 3508 TA Utrecht, The Netherlands ([email protected]) H. OHMOTO

Department of Geosciences, 435A Deike Building, Pennsylvania State University, University Park, PA 16802, USA ([email protected]; phone: 814-8654074) A. POLAT

Department of Earth Sciences, University of Windsor, Windsor, Ontario N9B 3P4, Canada (polat@uwindsor ca; phone: 519-253 3000, ext. 2498) M. POPA

Dept. of Geology, University of Pretoria, Pretoria 0002, South Africa ([email protected]; phone: 0124202242; fax: 012-3625219) A. PROKOPH

SPEEDSTAT, 36 Corley, Ottawa, Ontario K1V 877, Canada (aprokocon@aoLcom; phone: 613-9477341; fax: 613-9477396) R.H. RAINBIRD

Geological Survey of Canada, Continental Geoscience Division, 615 Booth Street, Room 609, Ottawa, Ontario K1A OE9, Canada ([email protected]; phone: 613-9477341; fax: 613-9477396) E RAMAEKERS 832 Parkwood Drive, S.E., Calgary, Alberta T2J 3W7, Canada ([email protected]; paulramaekers@

home.corn; phone: 403-2789423; fax: nil) S. SARKAR

Dept. of Geological Sciences, Jadavpur University, KoIkata 700 032, India (jugeoss@vsnLnet fax: 33-4731484) J. SCHIEBER

Dept. of Geological Sciences, Indiana University, Bloomington, IN 47405, USA ([email protected]; phone: 812-8564740;fax: 812-8557899) J.W. SCHOPF

CSEOL- Center for the Studies of the Evolution of Life, Geology Building, University of California Los Angeles (UCLA), CA 90024-1567, USA ([email protected]; phone: +1-310-8251170) B.M. SIMONSON

Dept. of Geology, Oberlin College, Oberlin, OH 44074-1044, USA ([email protected]; phone: 440-7758347; fax: 440-7758038) E.L. SIMPSON

Dept. of Physical Sciences, Kutztown University of Pennsylvania, Kutztown, PA 19530, USA (simpson@ kutztown.edu; phone: 610-6834445;fax: 610-6831352) J. STIX

Department of Earth and Planetary Sciences, McGill University, Montreal, Quebec H3A 2A7, Canada ([email protected]; phone: 514-3985391;fax: 514-3984680)

x

Contributing authors

E. TAMRAT

Dept. of Geological Sciences, University of Florida, Gainesville, FL 32611, USA ([email protected]; phone: 352-8462414; fax: 352-3929294) P.C. T H U R S T O N

Department of Earth Sciences, Laurentian University, Sudbury, Ontario P3E 2C6, Canada (pthurston @nickel. laurentian.ca; phone: 705-6751151x2372; fax: 705-6736508) A.E T R E N D A L L

Dept. of Applied Physics, Curtin University of Technology, Perth, WA 6004, Australia ([email protected]. edu.au; trendall @iinet.net.au; phone: 898-451006) R. VAN D E R M E R W E

Dept. of Geology, University of Pretoria, Pretoria 0002, South Africa ([email protected]; phone: 27-(0)82-5775004; fax: 27-12-3625219) M.J. VAN K R A N E N D O N K

Geological Survey of Western Australia, Department of Industry and Resources, Mineral House, 100 Plain Street, East Perth, WA 6004, Australia ([email protected]; phone: 9-2223333; fax: 9-2223633) E WESTALL

Centre Biophysique Moleculaire, Exobiologie, Centre National de la Recherche Scientifique, Rue CharlesSadron, 45071 Orleans Cedex 2, France ([email protected]; phone: 33-238- 2557912; fax: 33-238-631517) J.D.L. W H I T E

Geology Department, University of Otago, Dunedin, New Zealand ([email protected]; phone: 63-3-4799009; fax: 63-3-4797527) G.E. W I L L I A M S

Department of Geology and Geophysics, University of Adelaide, Adelaide, SA 5005, Australia (george. williams@ adelaide.edu.au; phone: 8-83035843; fax: 8-83034347) G.M. YOUNG

Dept. of Earth Sciences, Faculty of Science, University of Western Ontario, Biology and Geology Building, London, Ontario N6A 5B7, Canada ([email protected]; phone: 519-6613193; fax: 519-6613198) T.E. ZEGERS

European Space Agency, ESTEC, Keplerlaan 1, 2201 AZ Noordwijk, The Netherlands ([email protected]; phone: 31-71-5656585)

CONTENTS

Contributing authors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . PREFACE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W.U. Mueller, W. Altermann, O. Catuneanu, P.G. Eriksson and D.R. Nelson

Chapter 1. 1.1. 1.2. 1.3. 1.4. 1.5.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . D.R. Nelson Earth's Formation and First Billion Years . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . D.R. Nelson The Early Precambrian Stratigraphic Record of Large Extraterrestrial Impacts . . . . . . . . . . . . . B.M. Simonson, G.R. Byerly and D.R. Lowe Strategies for Finding the Record of Early Precambrian Impact Events . . . . . . . . . . . . . . . . . D.H. Abbott and J.T. Hagstrum Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ........... D.R. Nelson

Chapter 2. 2.1. 2.2.

2.3.

2.4.

2.5. 2.6.

2.7. 2.8.

THE EARLY EARTH . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Editor: D.R. Nelson

GENERATION OF CONTINENTAL CRUST Editors: D.R. Nelson and W.U. Mueller

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

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

W.U. Mueller and D.R. Nelson Isua Enigmas: Illusive Tectonic, Sedimentary, Volcanic and Organic Features of the > 3.7 Ga Isua Greenstone Belt, Southwest Greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.S. Myers Geochemical Diversity in Volcanic Rocks of the > 3.7 Ga Isua Greenstone Belt, Southern West Greenland: Implications for Mantle Composition and Geodynamic Processes . . . . . . . . . . . . . A. Polat, A.W. Hofmann and P. W. U. Appel Abitibi Greenstone Belt Plate Tectonics: The Diachrononous History of Arc Development, Accretion and Collision . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . R. Daigneault, W.U. Mueller and E.H. Chown Granite Formation and Emplacement as Indicators of Archaean Tectonic Processes . . . . . . . . . . T.E. Zegers Diapiric Processes in the Formation of Archaean Continental Crust, East Pilbara Granite-Greenstone Terrane, Australia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A.H. Hickman and M.J. Van Kranendonk Early Archaean Crustal Collapse Structures and Sedimentary Basin Dynamics . . . . . . . . . . . . W. Nijman and S.T. de Vries Crustal Growth Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . N. T. Arndt

v xvii

1

1 3 27 45 62

65 65

66

74

88 103

118 139 155

xii

Contents

2.9. C o m m e n t a r y . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . D.R. Nelson and W.U. Mueller

158

Chapter 3.

161

TECTONISM AND MANTLE PLUMES THROUGH TIME . . . . . . . . . . . . . . . . . . Editors: P.G. Eriksson and O. Catuneanu

3.1.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson and O. Catuneanu 3.2. Precambrian Superplume Events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . K. C. Condie .3. Large Igneous Province Record through Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . R.E. Ernst, K.L. Buchan and A. Prokoph 3.4. Episodic Crustal Growth During Catastrophic Global-Scale Mantle Overturn Events . . . . . . . . . D.R. Nelson 3.5. An Unusual Palaeoproterozoic Magmatic Event, the Ultrapotassic Christopher Island Formation, Baker Lake Group, Nunavut, Canada: Archaean Mantle Metasomatism and Palaeoproterozoic Mantle Reactivation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . B.L. Cousens, J.R. Chiarenzelli and L.B. Aspler 3.6. A Commentary on Precambrian Plate Tectonics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P, G. Eriksson and O. Catuneanu 3.7. Precambrian Ophiolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.R. Chiarenzelli and E.M. Moores 3.8. The Limpopo Belt of Southern Africa: A Neoarchaean to Palaeoproterozoic Orogen . . . . . . . . . A.J. Bumby and R. van der Merwe 3.9. Geodynamic Crustal Evolution and Long-Lived Supercontinents During the Palaeoproterozoic: Evidence from Granulite-Gneiss Belts, Collisional and Accretionary Orogens . . . . . . . . . . . . . M. V. Mints and A.N. Konilov 3.10. Formation of a Late Mesoproterozoic Supercontinent: The South Africa-East Antarctica Connection H.E. Frimmel 3.11. A Mechanism for Explaining Rapid Continental Motion in the Late Neoproterozoic . . . . . . . . . J.G. Meert and E. Tamrat 3.12. Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson and O. Catuneanu

Chapter 4. 4.1. 4.2. 4.3.

4.4.

PRECAMBRIAN VOLCANISM: AN INDEPENDENT VARIABLE THROUGH T I M E . . . Editor: W.U. Mueller

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W. U. Mueller and P. C. Thurston Terminology of Volcaniclastic and Volcanic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . W.U. Mueller and J.D.L White Komatiites: Volcanology, Geochemistry and Textures . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.1. Physical Volcano!ogy of Komatiites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W. U. MuelIer 4.3.2. Komatiite Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J. Dostal and W. U. Mueller 4.3.3. Textures in Komatiites and Variolitic Basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . N.T. Arndt and A.D. Fowler Archaean and Proterozoic Greenstone Belts: Setting and Evolution . . . . . . . . . . . . . . . . . . . P.C. Thurston and L.D. Ayres

161 163 173 180

183 201 213 217

223 240 255 267

271 271 273 277 277 290 298 311

Contents

4.5. 4.6. 4.7.

xiii

Explosive Subaqueous Volcanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.D.L. White Archaean Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W.U. Mueller, J. Stix, J.D.L. White and G.J. Hudak Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W. U. Mueller

Chapter 5.

THE EVOLUTION OF THE PRECAMBRIAN ATMOSPHERE: CARBON ISOTOPIC EVIDENCE FROM THE AUSTRALIAN CONTINENT . . . . . . . . . . . . . . . . . . . . Editors: P.G. Eriksson and W. Altermann

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson and W. Altermann 5.2. Archaean Atmosphere, Hydrosphere and Biosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . H. Ohmoto 5.3. Evolution of the Precambrian Atmosphere: Carbon Isotopic Evidence from the Australian Continent J.E Lindsay and M.D. Brasier 5.4. Precambrian Iron-Formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A.F. Trendall and J.G. Blockley 5.5. The Precambrian Sulphur Isotope Record of Evolving Atmospheric Oxygen . . . . . . . . . . . . . . T.W. Lyons, L.C. Kah and A.M. Gellatly 5.6. Earth's Two Great Precambrian Glaciations: Aftermath of the "Snowball Earth" Hypothesis . . . . . G.M. Young 5.7. The Paradox of Proterozoic Glaciomarine Deposition, Open Seas and Strong Seasonality Near the Palaeo-Equator: Global Implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . G.E. Williams 5.8. Neoproterozoic Sedimentation Rates and Timing of Glaciations--A Southern African Perspective . H.E. Frimmel 5.9. Earth's Precambrian Rotation and the Evolving Lunar Orbit: Implications of Tidal Rhythmite Data for Palaeogeophysics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . G.E. Williams 5.10. Ancient Climatic and Tectonic Settings Inferred from Palaeosols Developed on Igneous Rocks . . . H.W. Nesbitt and G.M. Young 5.11. Aggressive Archaean Weathering . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.L. Corcoran and W. U. Mueller 5.12. Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson and W Altermann 5.1.

Chapter 6. 6.1. 6.2. 6.3. 6.4. 6.5.

EVOLUTION OF LIFE AND PRECAMBRIAN BIO-GEOLOGY . . . . . . . . . . . . . . . Editor: W. Altermann

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W. Altermann Earth's Earliest Biosphere: Status of the Hunt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.W. Schopf Evolving Life and Its Effect on Precambrian Sedimentation . . . . . . . . . . . . . . . . . . . . . . . W. Altermann Microbial Origin of Precambrian Carbonates: Lessons from Modern Analogues . . . . . . . . . . . . J. Kazmierczak, S. Kempe and W. Altermann Precambrian Stromatolites: Problems in Definition, Classification, Morphology and Stratigraphy . . W. Altermann

334 345 356

359 359 361 388 403 421 440

448 459

473 482 494 505

513 513 516 539 545 564

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6.6. 6.7.

Contents

Precambrian Geology and Exobiology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . E Westall Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . W. Altermann

Chapter 7.

SEDIMENTATION THROUGH TIME Editor: P.G. Eriksson

575 587

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

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson, A.J. Bumby and M. Popa 7.2. Sedimentary Structures: An Essential Key for Interpreting the Precambrian Rock Record . . . . . . J.A. Donaldson, L.B. Aspler and J.R. Chiarenzelli 7.3. Archaean Sedimentary Sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.L. Corcoran and W. U. Mueller 7.4. Discussion of Selected Techniques and Problems in the Field Mapping and Interpretation of Archaean Clastic Metasedimentary Rocks of the Superior Province, Canada . . . . . . . . . . . . . . J.R. Devaney 7.5. Precambrian Tidalites: Recognition and Significance . . . . . . . . . . . . . . . . . . . . . . . . . . . K.A. Eriksson and E.L. Simpson 7.6. Sedimentary Dynamics of Precambrian Aeolianites . . . . . . . . . . . . . . . . . . . . . . . . . . . E.L. Simpson, F.E Alkmim, P.K. Bose, A.J. Bumby, K.A. Eriksson, P.G. Eriksson, M.A. Martins-Nero, L.T. Middleton and R.H. Rainbird 7.7. Early Precambrian Epeiric Seas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson, A.J. Bumby and P. Mostert 7.8. Precambrian Rivers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . D. G.E Long 7.9. Microbial Mats in the Siliciclastic Rock Record: A Summary of Diagnostic Features . . . . . . . . . J. Schieber 7.10. Microbial Mat Features in Sandstones Illustrated . . . . . . . . . . . . . . . . . . . . . . . . . . . . . S. Sarkar, S. Banerjee and P.G. Eriksson 7.11. Sedimentation Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson, P.K. Bose, S. Sarkar and S. Banerjee 7.12. Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson and M.A. Martins-Nero 7.1.

Chapter 8.

SEQUENCE STRATIGRAPHY AND THE PRECAMBRIAN Editor: O. Catuneanu

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

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A.E Embry, O. Catuneanu and P.G. Eriksson 8.2. Concepts of Sequence Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . O. Catuneanu, A.E Embry and P.G. Eriksson 8.3. Development and Sequences of the Athabasca Basin, Early Proterozoic, Saskatchewan and Alberta, Canada . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P. Ramaekers and O. Catuneanu 8.4. Third-Order Sequence Stratigraphy in the Palaeoproterozoic Daspoort Formation (Pretoria Group, Transvaal Supergroup), Kaapvaal Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson and O. Catuneanu 8.5. Commentary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . O. Catuneanu and P.G. Eriksson

8.1.

593 593 602 613

625 631 642

657 660 663 673 675 677

681 681 685

705

724 735

Contents

Chapter 9. 9.1. 9.2. 9.3. 9.4. 9.5. 9.6. 9.7. 9.8. 9.9.

xv

TOWARDS A SYNTHESIS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.G. Eriksson, O. Catuneanu, D.R. Nelson, W.U. Mueller and W. Altermann

Evolution of the Solar System and the Early Earth . . . . . . . . . . . . . . . . . . . . . . . . . . . . Generation of Continental Crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonism and Mantle Plumes through Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Precambrian Volcanism, an Independent Variable . . . . . . . . . . . . . . . . . . . . . . . . . . . . Evolution of the Hydrosphere and Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Evolution of Precambrian Life and Bio-Geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sedimentation Regimes through Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sequence Stratigraphy through Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tempos and Events in Precambrian Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

References

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

Subject Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

739 739 743 747 749 751 755 758 761 762 771 923

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xvii

PREFACE

Many facets of Earth's evolution are intriguing, yet most of our attention is drawn to the more recent phases, as from the Cambrian epoch (c. < 540 Ma). The Precambrian Era, from 4.5 to 0.54 Ga, covers almost 90% of this planet's history but our knowledge of this segment of time remains elusive, save for selective aspects. Even though extensive research has been conducted on some topics, a significant diversity of opinions amongst researchers is evident and generally concerns the origin, the mechanism or the process inherent to a specific criterion. For example, did the atmosphere change with time? Are palaeoweathering surfaces or regoliths indicative of anoxic or oxygenated climatic conditions? Many other unresolved questions remain. What were the prevalent large-scale driving forces during continent or greenstone belt formation? Have sedimentary patterns evolved with time, or is there just a shift in the prevalent transport process? Can sequence stratigraphy be applied to Archaean sequences, or is better temporal resolution required? Can tidal sequences be identified and characterised in 3.2 Ga sedimentary rocks? Can we reconstruct celestial mechanics within the solar system using modern tidal models and calculations? Have volcanic processes changed over geological time? Why are komatiites largely constrained to the Archaean era? Have supercontinents driven global glaciation periods? The dictum "the present is key to the past" was borne out of the necessity that specific geological problems or features in the past can not be adequately explained or addressed due to limited information or outcrop exposure; there is also a large measure of truth inthe reverse axiom. Ancient surface-forming sedimentary structures (i.e., bedforms and macroforms) can be compared to present-day structures because we assume processes operative on land and sea that produced this structure were comparable and independent of the era. Actualism is a non-gradualistic uniformitarian approach advocating that similar (the same?) processes and invariant natural physical, chemical and biological laws were operative, and hence applicable to both the Precambrian and present. Relative rates of processes or events, such as mid-ocean ridge spreading and subduction, weathering, genesis of continental crust, the rotation of Earth, and the atmospheric evolution probably contrasted with those of Phanerozoic successions, but the large and small scale, surface-forming processes or mechanisms are certainly comparable. For specific aspects, the Precambrian rock record is more useful in explaining processes or mechanisms (of early Earth) than observation and reading of recent processes which are difficult to access. For example, ancient subaqueous volcanic sequences can often be better loci for identifying and quantifying volcanic eruption mechanisms or depositional processes of the early Earth than modern ocean floor settings. This is simply because we cannot access the ocean floor without expensive high technology machinery (i.e., submersibles and ROVs), or map the ocean floor systematically, but extensive subaqueous

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Preface

volcanic sequences are exposed on land. Despite being highly deformed and metamorphosed, ancient Precambrian rocks do offer advantages. For example, mountain-building phases regularly expose thick ancient sequences of oceanic crust, and the nuclei of most continents have well-exposed remnants of Archaean crust (e.g., Superior Province, Canada and Kaapvaal craton, South Africa). These extensive Archaean cratons, present on all major continents, permit us to evaluate how the mechanisms and processes of crustal evolution changed over time. The details of the mechanisms driving the Earth and its development inevitably changed with time. Interaction between erosional and depositional surface processes, mantle dynamics and large-scale horizontal plate motions, mantle- and crust-derived igneous rocks, as well as the complex interplay between atmosphere, climate and biodiversity, reveal the hallmarks of evolution through time, hence the title of this book "The Preeambrian Earth: Tempos and Events". The title conveys the notion that nothing remains completely the same through time and each of the chapters focuses on how this change came about and how it can be explained. It is the latter that arouses geologists' interest. Geologists are renowned for their divergent opinions, and this book adheres to this tradition. Controversy often drives research, and in all chapters there is never one unique solution, opinion, or view but rather a representative cross-section of possible explanations. The editors have sought to synthesise each chapter so that the reader can assimilate more easily all this new and highly condensed information. Throughout the book a diversity of opinion was encouraged and supported. So what are "tempos and events"? Continents have been formed, accreted and dismembered regularly throughout Earth's 4.5 Ga history. The conditions and parameters driving such continental evolutionary events will have changed. This book attempts to place change in an overall context. A continent is an amalgamation of accreted material including volcanic ocean floor and arcs, sedimentary trench deposits, continental platform deposits, but also abundant plutonic material. In the Precambrian, supercontinents formed at various times, with the best known being: (1) a Neoarchaean "northern" (present-day frame of reference) "Kenorland"; (2) a "southern" continent at c. 2.2-1.8 Ga, with (3) the approximately coeval, "northern" Laurentia at c. 2.0-1.7 Ga, to be followed (4) by Columbia in the Mesoproterozoic, (5) the Neoproterozoic had a succession of supercontinents (e.g., Rodinia), ending with the Phanerozoic Gondwana landmasses. The term tempos refers to the rate over a time span at which a certain event (i.e., formation of a supercontinent) or process (i.e., alluvial fan or volcanic edifice construction) occurred. Significant events include superplume volcanism, palaeo-atmospheric changes, advances or retreats of living organisms, orogenic mountain-building phases, continental breakup, as well as global sedimentation patterns (glaciogenic deposits, iron-formations or giant carbonate platforms) and associated sea level changes. Intuitively with the changing dynamics of large-scale endogenous and exogenous processes, the time scale of events and processes must have changed, but some processes or mechanisms responsible for a temporal change of events are difficult to quantify. Each chapter is a reflection of specific events, and the term evolution includes the notion of a temporal change. Not all authors agree on how and over what time scale specific events have occurred, but we are all of the opinion that Earth has evolved. Many

Preface

xix

roads lead to Rome, with some being more direct and evident than others. This book is the outgrowth of numerous research avenues documenting change with time, or as the rifle indicates, tempos and events. The book is a comprehensive entity, but each chapter and individual paper is designed so that it can be read and understood in isolation. The first chapter examines the celestial origins of our planet and solar system and the early differentiation into core, mantle, crust and primitive atmospheres. Chapter 2 is devoted to understanding the formation of granitegreenstone crust. Plate tectonics (and any possible forebear mechanisms) and mantle superplumes surely are related to continental (and oceanic) crustal growth, but are treated separately in chapter 3, as these concepts are both complex and contentious. Chapter 4 considers Precambrian volcanism as an independent variable through time; again, there are strong relationships with the preceding two chapters. Overlap and common themes naturally abound for the successive chapters 5 (evolution of the atmosphere-hydrosphere), 6 (life and bio-geology), 7 (sedimentation through time) and 8 (Precambrian sequence stratigraphy). We are the first to admit that many different structures and chapter outlays could have been adopted, but they grew naturally as the book developed and helped modify some of the prejudices the editors had, as the data set enlarged and new viewpoints were presented. Every chapter has an introduction by one or two editors to the relevant theme(s) being covered, followed by a succession of differently-authored papers, arranged in a specific order chosen by the editorial team. An editorial commentary terminates the chapter incorporating and discussing the views presented. The final chapter synthesises an enormous range of geological "events" occurring at different time periods ("tempos") over c. 4 Ga of Precambrian Earth evolution. The introduction, the various papers, and the closing commentary are termed "sections". The sections have a different authorship team, and are numbered for ease of cross-referencing between sections and chapters. Figures and tables are similarly numbered, based on the adopted section system. Throughout the book we have adopted a few conventions, quite apart from an "English" spelling system that may be strange to some North AmeriCan readers. Terms "Ma" and "Ga" represent millions and billions (10 9) of years before Present, with "My" being used to denote a time period unrelated to any datum. Emphasised text has been underlined and italics are used for non-English terms of common usage. We have found the editing and writing of this book to be an extremely rewarding and educational exercise, and hope the reader will also find it beneficial. We owe our large community of authors from all over the world our gratitude for their tolerance of our numerous requests and changes.

ACKNOWLEDGEMENTS Firstly, Friso Veenstra of Elsevier who approved this book on the recommendation of Kent Condie, the series editor for Elsevier's book-series "Developments in Precambrian Geology". Kent Condie has trodden deep steps within Precambrian geological thought and

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literature, and we have all learnt much from him over the years; we also acknowledge his great encouragement of the book as a whole and for advice on the theme chosen. Patricia Massar of Elsevier was of great assistance in explaining the technical details of manuscript preparation. Without the skills, patience and infinite capacity of Magda Geringer (Department of Geography, University of Pretoria), this book would not have seen the light of day. She did most of the drafting and upgraded all figures, quite apart from dealing with a large array of different software packagesmwe are in her debt. Financially, the University of Pretoria, the University of Alberta and the University of Quebec at Chicoutimi supported drafting costs generously. Wulf Mueller would like to thank LITHOPROBE for its significant contributions in the last 10 years because much of our (my) knowledge concerning Canadian Precambrian craton evolution and characteristics stems from this pan-Canadian research project. Apart from myself, numerous authors in this book were supported by LITHOPROBE. Wlady Altermann is indebted to LE STUDIUM and the CNRS, Orl6ans, and to the GeoBioCenter LMU for support during the major phase of preparing this book. Pat Eriksson acknowledges generous financial support from the University of Pretoria towards this book. No endeavour of this magnitude is possible without a very large team of qualified and willing reviewers. They are gratefully acknowledged and listed below: A. Altenbach (Ludwig-Maximilians University, Munich) W. Altermann (Centre National de la Recherche Scientifique, Orl6ans) A.D. Anbar (Univ. of Rochester, New York) L.B. Aspler (Consultant, Ottawa) S.-J. Barnes (Univ. of Quebec, Chicoutimi) R.S. Blewett (Geoscience Australia, Canberra) A. Brack (Centre National de la Recherche Scientifique, Orl6ans) G. Brandl (Council for Geoscience, South Africa) A.J. Bumby (Univ. of Pretoria, South Africa) D. Champion (Geoscience Australia, Canberra) J.R. Chiarenzelli (State Univ. of New York, Potsdam) RL. Corcoran (Univ. of Western Ontario) E. Cotter (Bucknell Univ., USA) R Cousineau (Univ. of Quebec, Chicoutimi) J. Dann (Massachusetts Institute of Technology) J.R. de Laeter (Curtin Univ., Australia) M. de Wit (Univ. of Cape Town, South Africa) J.A. Donaldson (Carleton Univ., Ottawa) B.G. Els (Univ. of Pretoria, South Africa) K.A. Eriksson (Virgina Polytechnic, USA) R G. Eriksson (Univ. of Pretoria, South Africa) R.E. Ernst (Geological Survey of Canada, Ottawa) D.A.D. Evans (Yale Univ.) C.M. Fedo (George Washington Univ.) R Fralick (Lakehead Univ.)

Preface

T. Frisch (Geological Survey of Canada, Ottawa) W.E. Galloway (Univ. of Texas at Austin) A. Gaucher (Univ. of Montevideo) G. Gerdes (Universit~t Oldenburg, Wilhelmshaven) G.J.B. Germs (Rand Afrikaans Univ., Johannesburg, South Africa) B.E Glass (Univ. of Delaware) C. Glatz (Univ. of Maine) R. Gorbatschev (Lund Univ., Sweden) J.W. Hagedorn (Amherst College, USA) R.E. Hanson (Univ. of Fort Worth, Texas) S. Hassler (California State Univ., Hayward) B.M. Jakosky (Univ. of Colorado, Boulder) C.W. Jefferson (Geological Survey of Canada, Ottawa) B.S. Kamber (Univ. of Queensland, Australia) E King (Univ. of Western Ontario) C. Klein (Univ. of New Mexico, Albuquerque) B. Lafrance (Laurentian Univ., Ontario) D. Lescinsky (Univ. of Western Ontario) J.H. Lipps (Univ. of California, Berkeley) D.G.E Long (Laurentian Univ., Ontario) M.A. Martins-Neto (Univ. of Ouro Preto, Brazil) J. McPhie (Univ. of Tasmania, Australia) J. Menzies (Brock Univ., Ontario) A.D. Miall (Univ. of Toronto) R. Morris (CSIRO, Australia) W.U. Mueller (Univ. of Quebec; Chicoutimi) D.R. Nelson (Geol. Surv. Western Australia and Curtin Univ., Australia) A.A. Nemchin (Curtin Univ., Australia) A. Nunes (Barnard College, New York) C.D. Ollier (Univ. of Western Australia) R.T. Pidgeon (Curtin Univ., Australia) M. Popa (Univ. of Bucharest) H.W. Posamentier (Anadarko Canada Corporation, Calgary) D. Powars (US Geological Survey, Stephens City, Virginia) R.H. Rainbird (Geol. Survey of Canada, Ottawa) K.J.R. Rosman (Curtin Univ., Australia) D. Schmidt (Ludwig-Maximilians Univ., Munich) C.A. Smit (Rand Afrikaans Univ., Johannesburg, South Africa) R.H. Smithies (Geol. Surv. Western Australia) G.M. Stott (Ontario Geol. Survey, Sudbury) H. Strauss (Westf~lische Wilhelms-Universit~t, MUnster) L. Tack (Royal Museum for Central Africa, Belgium) H. Tirsgaard (Maersk Oil, Copenhagen)

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T. Torsvik (Norwegian Geological Survey, Trondheim) A.E Trendall (Curtin Univ., Australia) D.W. Valentino (State Univ. of New York, Oswego) R. van der Voo (Univ. of Michigan) A.J. van Loon (Benitachell Univ., Spain) M.R. Walter (Macquarie Univ., Australia) R. Weinberg (Monash Univ., Australia) G. Whitmore (Univ. of Natal, Durban, South Africa) M.C. Wizevich (Southern Connecticut State Univ.) Additional acknowledgements are listed below, related to specific sections of the book. D.R. Nelson (section 1.2) dedicates this section to Emeritus Professor John R. de Laeter, A.O., Cit (WA), B.Ed. (Hons), B.Sc. (Hons), Ph.D.D.Sc. (W. Aust.), Hon Tech (Curtin), FTSE, FInstP, FAIP, whose inspiration, leadership and example over the last 15 years constitutes an incalculable personal, professional and scientific contribution. The Perth Sensitive High-Resolution Ion MicroProbe is operated by a consortium consisting of Curtin University, the Geological Survey of Western Australia and the University of Western Australia with the support of the Australian Research Council. Published with permission of the Director, GSWA. D.H. Abbott and J.T. Hagstrum (section 1.4) thank S. Hoffman for editing and D. Breger for her help and expertise on the SEM. A. Polat, A.W. Hoffmann and P.W.U. Appel (section 2.3) thank S. Moorbath, J.S. Myers, S. Hanmer, M. Rosing, R. Frei and H. Rollinson for scientific discussion on the Isua greenstone belt. R. Kerrich and A. Trenhaile are acknowledged for their comprehensive critique of the initial draft of the manuscript. A. Polat is: grateful to the Max-Planck-Institut (Mainz) and the Alexander yon Humboldt Foundation for financial and logistic support. NSERC Grant #250926-02 to A. Polat is acknowledged. This is a contribution of the Isua Multi-disciplinary Research Project. R. Daigneault, W.U. Mueller and E.H. Chown (section 2.4) note that this Abitibi synthesis study was the outgrowth of 20 years mapping by the three of us and many students that paved the way with their data and theses. The Quebec Ministbre des Richesses naturelles (MRNQ) is thanked for constant support, as are the exploration companies. NSERC, FUQUAC, and LITHOPROBE funding (LITHOPROBE contribution no. 1325) are gratefully acknowledged. A.H. Hickman and M.J. van Kranendonk (section 2.6) note that thiscontribution uses information from numerous sources, but particularly from colleagues at the Geological Survey of Western Australia (GSWA), Geo5 science Australia, the University of Newcastle (W.J. Collins and M. Pawley), and Sipa Resources Ltd., Perth (E Morant and C. Brauhart). Their interpretations have benefited from discussions with all these people, but especially with R.H. Smithies, D.R. Nelson, L. Bagas, I.R. Williams, T. Farrell, and E Pirajno. Both authors publish with permission of the Director, Geological Survey of Western Australia. W. Nijman and S.T. de Vries (section 2.7) note that some of the observations and conclusions form part of a Ph.D. study of the second author, to be published in 2003. Gratefully we acknowledge the financial support of the project on Earth's Earliest Sedimentary Basins by the Foundation Dr. Schtirmannfonds in the Netherlands and the cooperation with Maarten de Wit, Frances Westall, Manfred van Bergen, Hanan Kisch, Dave Nelson, and Richard Armstrong. The manu-

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xxiii

script benefited much from comments from Poppe de Boer. Last but not least, we thank the many former Utrecht M.Sc. students for their contributions to the project. R.E. Ernst, K.L. Buehan and A. Prokoph (section 3.3) note that this is Geological Survey of Canada publication 2002188. B.L. Cousens, J.R. Chiarenzdli and L.B. Aspler (section 3.5) acknowledge funding provided through contracts with the Yellowknife Geology Division, Indian and Northern Affairs Canada, facilitated by Bill Padgham and Carolyn Relf, as part of the Western Churchill NATMAP project. Rex Brommecker of WMC originally suggested we examine the Rack Lake drill core. Thanks to A1 Donaldson, Hamish Sandeman, Tony Peterson, Tony LeCheminant, Eva Zaleski, Rob Rainbird, Thomas Hadlari, Derek Smith, Yannick Beaudoin, Ken Ashton, Russell Hartlaub, and Simon Hanmer for providing samples. A1 Armitage and Ryan Morelli kindly provided unpublished analyses from MacQuoid-Gibson dykes and Martin Formation volcanic rocks, respectively. Thanks to the XRF facility at the University of Ottawa and the Ontario Geological Survey laboratories for major and trace element analyses. Donna Switzer, Julie Thompson, Muy Ngo, Brenda Obina and Samantha Seigel performed much of the isotopic lab work. We benefited from discussions with Hamish Sandeman, Tony Peterson, Tony LeCheminant, Rob Rainbird, A1 Donaldson and A1 Miller. M.V. Mints and A.N. Konilov (section 3.9) point out that their paper developed from discussions at workshops sponsored by the COPENA IGCP Project 371 and SVEKALAPKO, a EUROPROBE project during recent years. The work was supported by the Russian Foundation for Basic Research, Projects No. 00-0564241 and No. 01-05-64373. The authors are very grateful to colleagues who have been participating in ongoing studies of the Early Precambrian in the Kola Peninsula and Karelia. H.E. Frimmel (section 3.10) thanks S. Perrit and S. Helferich for providing zircon separates from samples collected in the Borgmassivet area and southern Kirwanveggen, respectively. G. Doyle and A. Bisnath are thanked for the companionship during the 2001/02 Antarctic field season. Both analytical and field work in Antarctica were funded by the Department of Environmental Affairs and Tourism within the framework of the South African National Antarctic Programme. J.G. Meert and E. Tamrat (section 3.11) wish to thank Monika Lipinski for an early review of the manuscript. H. Ohmoto (section 5.2) thanks the following persons for providing stimulating discussions, over years, on the various topics related to this paper: Dick Holland, Jim Kasting, Lee Kump, Mike Arthur, Chris House, Kate Freeman, Tony Lasaga, Nic Beukes, Jenz Gutzmer, Lawry Minter, Mike Kimberley, Ken Towe, Yumiko Watanabe, Kosei Yamaguchi, Takeshi Kakegawa, Ken Hayashi, Hiroshi Naraoka, Munetomo and Yoko Nedachi, and Yasu Kato. Financial support from NASA Astrobiology Program (NCC2-1057), NASA Exobiology Program (NAG5-9089) and NSF (EAR-9706279) is gratefully acknowledged. T.W. Lyons, L.C. Kah and A.M. Gellafly (section 5.5) acknowledge financial support by NSF awards EAR-9596079 and EAR-9725538. The Indiana University Stable Isotope Facility, and in particular Jon Fong and Steve Studley, assisted with many of the isotopic analyses. The authors also benefited from many fruitful conversations with Don Canfield, Mike Formolo, Tracy Frank, Matt Hurtgen, Jim Luepke, and Don Winston. G.M. Young (section 5.6) is grateful to the National Scientific and Engineering Research Council of Canada for generous support of investigation of Precambrian glaciogenic rocks over the years and to indi-

xxiv

Preface

viduals too numerous to mention for discussions and field trips related to glaciogenic rocks. H.E. Frilnmel (section 5.8) acknowledges research funding by the South African National Research Foundation. W.J. Sehopf (section 6.2) notes that this research was supported by NASA, though Grant NAG 5-12357 and the Astrobiology Institute. J. Kazmierezak, S. Kempe and W. Altermann (section 6.4) acknowledge financial support from the Foundation for Polish Science (Scholar Grant 2000 to J.K.) and from the Deutsche Forschungsgemeinschaft (to S.K.). W.A. is grateful to the University of Pretoria, the GeoBio-Center LMU Germany and to the Centre Biophysique Moleculaire, CNRS, Orleans, France for technical support. J.R. Devaney (section 7.4): aside from a thesis completed at Lakehead University and supervised by Phil Fralick, my sedimentologically-oriented studies of Archaean greenstone belts were done while employed by the Ontario Geological Survey. J. Sehieber (section 7.9) is grateful to Wolfgang Krumbein for his enthusiastic encouragement of microbial mat research in the early stages of his career, as well as for his sustained motivation and advice since then. Support for research on sediments that yielded microbial mat features was provided by the National Science Foundation (Grants #EAR-9117701 and EAR,9706178) and the Donors of the Petroleum Research Fund, administered by the American Chemical Society (Grants #25134-AC2 and 33941-AC8). A. Embry, O. Catuneanu and P.G. Eriksson (sections 8.1 and 8.2) acknowledge financial support from the University of Alberta and NSERC Canada (OC); AFE thanks the Geological Survey of Canada for encouraging the research on sequence stratigraphy and for allowing the publication of this paper. PGE acknowledges research support from the University of Pretoria. E Ramaekers and O. Catuneanu (section 8.3): this study benefited from work carried out by P.R. at the Saskatchewan and Alberta Geological Surveys, the Saskatchewan Mining Development Corporation (now Cameco), ongoing work in the Athabasca Basin as a consultant, discussion with industry staff, especially those of Cameco and Cogema, and coworkers of the Extech IV Research project (Jefferson et al., 2002), especially Charlie Jefferson, Gary Yeo, and Darrel Long. O.C. acknowledges research support from the University of Alberta and NSERC Canada. Last, and by no means least, we thank our long-suffering partners, whose support and understanding during preparation of this book, and at all times, remain very important to us: Patricia, Renate, Gabriela, Mari~inne and Janet. June 2003

Wulf U. Mueller (Chicoutimi) Wladyslaw Altermann (Orldans) Octavian Catuneanu (Edmonton) Patrick G. Eriksson (Pretoria) David R. Nelson (Perth)

The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) Published by Elsevier B.V.

Chapter 1

THE EARLY E A R T H

1.1. INTRODUCTION D.R. NELSON Inferences about the pre-4.0 Ga geological history of the Earth have been based traditionally either on the study of the oldest identified remnants on the Earth's surface (e.g., Maas et al., 1992; Nutman et al., 1996; Amelin et al., 1999; Nelson et al., 2000; Ryder et al., 2000; Wilde et al., 2001; Mojzsis et al., 2001), or on modelling of the differentiation of global chemical reservoirs (e.g., Arndt and Chauvel, 1990; Bennett et al., 1993; Bowring and Housh, 1995; Kramers and Tolstikhin, 1997; Snow and Schmidt, 1998; Albarbde et al., 2000; Canfield et al., 2000; Nutman et al., 2001). A major limitation of these approaches arises from the limited tangible evidence available for study of early Earth--the preserved rock record commences at 4030 Ma (Stern and Bleeker, 1998; Bowring and Williams, 1999), more than 500 My after the Earth's formation. As a consequence, these approaches have so far provided only broad constraints on the mechanisms and time scales of accretion and early differentiation of the Earth, and of physicochemical conditions on the Earth's surface during this time. In section 1.2 of this chapter, a new approach to the study of the early Earth, based on detailed chemical and isotopic studies of meteorites in combination with advances in our understanding of nucleosynthesis, has been investigated. In 1960, the remarkable discovery by J.H. Reynolds of the isotope xenon-129 (129Xe) within the earliest-forming phase of a primitive meteorite (Reynolds, 1960; see also Jeffery and Reynolds, 1961) was eventually to lead to a breakthrough in our understanding of the timing of accretion and differentiation of the Earth. The 129Xe detected by Reynolds had accumulated in situ from the radiogenic decay of the long-extinct nuclide iodine- 129 (129I), which has a half-life of only c. 16 My. The daughter products of a number of other extinct nuclides have since been identified within primitive meteorites, and it is now generally accepted that their short-lived radioactive parent nuclides were synthesised during supernova explosions in the vicinity and shortly before the formation of our solar system. These catastrophic nucleosynthesis events mark the time at which the radioactive isotopes that are widely used for geochronology were formed. As they are now long extinct, short-lived nuclides cannot be used directly to obtain absolute dates relative to the present-day, but their short half-lives have been used to constrain precisely the relative chronologies of planetary formation milestones for the early solar system (see Fig. 1.1-1). The Earth and other terrestrial planets formed by the collision and amalgamation of smaller rocky planetesimals within the early solar system's protoplanetary disk. During the later stages of this accretion process, progressively larger planetary embryos were formed

2

Chapter 1: The Early Earth

Type II supernova event (4571 Ma) 4600 triggered collapse of precursor interstellar molecular cloud

4500 4540

4400

formation of CAI's (4570 Ma) first chondrules (4565 Ma) first planetesimals (c. 4565-4550 Ma)

4300

proto-Earth accretion, core formation

4200

?Mars-sized impactor, formation of Moon o

4100 4000

I

detrital zircons (ZrSiO4) from the Yilgam Craton Acasta orthogneisses

3900 ~ 3800

"

Mt Sones, Enderby Land

3700

Isua greenstone belt, Greenland Manfred Complex, NarryerTerrane

3600

Ancient Gneiss Complex, Kaapvaal Craton orthogneiss in Warrawagine Complex, Pilbara Craton

3500

Coonterunah Formation, Pilbara Craton

3400

granite-greenstonecrust, Pilbara and Kaapvaalcratons

Fig. 1.1-1. Chronology of major events during formation of the solar system and the early Earth (see section 1.2 for further details).

and collided. These violent collisions resulted in the episodic reforming of the growing proto-Earth, along with the destruction of much of the evidence of the extent of earlier differentiation. The Earth's Moon also probably formed as a result of such a catastrophic collision during the later stages of Earth accretion. As the planetary embryos grew, the impact rate decreased and the chances of survival of these early-formed fragments of the Earth's surface increased. In section 1.3 of this chapter, Simonson et al. argue that terrestrial impact structures predating the Proterozoic era (> 2.5 Ga) are unlikely to have survived, due to the fragmentary state of preservation of the Earth's rock record from this time. Fortunately, evidence of such early impact events may be preserved in the Earth's stratigraphic archive, as thin layers rich in distinctive sand-sized spherules. In section 1.4, Abbott and Hagstrum estimate that in the time interval between 3.8 and 2.5 Ga, there were more than 350 impact events large enough to produce an impact layer of global extent. It has also been proposed (section 1.4) that major magmatic and (by implication) crustformation events during the Archaean could have been related to major impact episodes. Although it is widely acknowledged that major impacts must have played an important role in the formation of the Earth's early continental crust, this "extraterrestrial" influence has largely been overlooked in most previous studies of the Earth's Archaean terranes. Recognition and detailed investigation of impact-related sedimentary rocks preserved in the Earth's stratigraphic record currently is still in its infancy, but the way ahead is clearer from studies such as those documented in sections 1.3 and 1.4 of this chapter.

1.2. E a r t h ' s k b r m a t i o n a n d First Billion Years

3

1.2. EARTH'S FORMATION AND FIRST BILLION YEARS D.R. NELSON Introduction

In this section, a new approach to the study of the early Earth, commencing before the time of formation of our solar system at 4571 Ma and working forward in time towards 3500 Ma, has been investigated. This approach explores recent insights into the processes active during formation of the early Earth arising from detailed chemical and isotopic studies of meteorites, combined with advances in our understanding of nucleosynthesis. Many meteorites are fragments of asteroids formed early in the evolutionary history of the solar system, that were too small to have undergone much internal heating (see Hutchison et al., 2001). Some contain refractory calcium- and aluminium-rich inclusions that condensed from the nebula when temperatures were so high that other elements were volatile, shortly after formation of the Sun and during dissipation of the nebula. Others represent disrupted fragments of planetesimals and differentiated planetary bodies, including the Moon and Mars, formed later in the accretion history of the solar system. Some meteorite classes are samples of the interiors of disrupted planetary bodies, and have formed prior to, during and after active differentiation of these bodies. They may therefore provide unique information about the processes operating during the early differentiation of the Earth into silicate crust and mantle, and metallic core. The identification of short-lived (with half-lives less than 100 My) radioactive, now extinct, nuclides within some classes of meteorites has imposed important new constraints on the early evolution of the solar system and on accretion and differentiation rates for planetary bodies such as the Earth. Short-lived nuclides potentially offer the means to precisely constrain early solar system chronology, and of planetary accretion and differentiation processes, in relation to the time of nucleosynthesis. To fully appreciate the insights offered by extinct nuclides into the chronology of the early solar system and formation and differentiation its planets including the Earth, an understanding of the processes involved in the synthesis of the elements prior to the formation of our solar system is required. Details of nucleosynthesis within stars were formulated by the pioneering work of E.M. Burbidge, G.R. Burbidge, Fowler, Hoyle and co-workers (Burbidge et al., 1957) and independently, by Cameron (1957). With the exception of the element hydrogen (H) and possibly some of the helium (He), lithium (Li), beryllium (Be) and boron (B) which may have been synthesised during the Big Bang or by spallation reactions, elements lighter than iron (Fe) now present in our solar system were created primarily by fusion reactions within the interiors of stars. Elements heavier than Fe were mostly synthesised by two major neutron-capture processes; the "slow" or s-process, which refers to the slow capture, relative to the rate of fl-decay, of neutrons within stars, and by the "rapid" or r-process, mostly in catastrophic supernovae events during which unstable intermediate isotopes form by the capture of neutrons in a neutron-dense environment and so rapidly that they do not have time to decay. (Some less abundant neutron-deficient, proton-rich The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann. D.R. Nelson, W.U. Mueller and O. Catuneanu

4

Chapter 1: 7"he Early Earth

nuclei were synthesised by a third px:ocess, the "proton" or p-process.) In this section, the processes by which short-lived nuclides were synthesised and implications of their identification within early-formed solar system materials for the mechanisms and time scales of formation of the solar system and its planets will be reviewed briefly. Additional important areas of investigation of the early Earth have recently arisen largely as a result of advances in experimental techniques. These include the recent development and application of methods for the isotopic analysis of elements with high ionisation potentials using multicollector inductively coupled plasma mass spectrometry (MC-ICP-MS). In addition, geochemical and isotopic studies of the noble gases have provided constraints on the evolution of the atmosphere (see chapter 5) and history of degassing of the Earth's mantle. The ion microprobe has also recently been applied to the investigation of the only remnants of the pre-4000 Ma Earth available for study, the few small (c. 200 ~tm long) 4400 to 4000 Ma grains of the common accessory mineral zircon (ZrSiO4) found in sedimentary and granitic rocks in Western Australia. It is the purpose of this contribution to provide an overview of these new areas of investigation and to summarise their implications for the pre-3500 Ma history of the Earth.

Synthesis of the Heavy Elements and Short-Lived Nuclides Many heavier nuclides, including unstable nuclides with short lifetimes and that are now extinct, are synthesised by the r-process during core-collapse supernova events. In this section, the stellar environments in which these radionuclides are synthesised and the evidence for and some implications of the presence of extinct nuclides within early-formed materials of our solar system are outlined. Star formation commences with the fragmentation and collapse of the denser parts of interstellar dust and gas within so-called "dark nebula" regions into "proto-stars" of typically 10 to 50 solar masses and 0.1 parsecs ( p a r s e c - 3.09 • 1013 km or 3.26 light-years) in diameter. The proto-star mass continues to increase by the accretion of in-falling material until the increasing density and temperature within its core triggers thermonuclear fusion. Strong stellar winds may then reduce the in-fall and inhibit further growth of the star. An isolated star more massive than about 10 solar masses will evolve very rapidly and have a very short pre- "main sequence" (pre- hydrogen-burning) history of ~< 105 years. Once formed, it may consume all of its available hydrogen fuel within 10 My. Its core will contract progressively until the temperature becomes high enough for He burning to commence. As the core temperature continues to increase, He will be able to combine by the triple-or reaction to build carbon (lZc), and (by or-capture) oxygen (160), with all the He consumed after about 1 My. The core progressively contracts until "carbon-burning" can commence. Carbon nuclei then react with one another to form neon (Ne), sodium (23Na) and magnesium (about 10,000 years; 25Mg is the essential "seed" required for production of the short-lived nuclide 26A1, mainly via the reaction 25Mg(p, y)26A1; see Table 1.2-1), then Ne to O and Mg (about 12 years), O to silicon (Si) and sulphur (S; about 4 years) and finally, burning of Si to 54Fe and a variety of other neutron-rich nuclides with masses near 50 or 60 (about 1 week). Because Fe will not fuse to produce more energy, the star's

1.2. Earth's bbrmation and kirst Billion Years

5

available nuclear fuel nears exhaustion and the radiation pressure generated via fusion is not sufficient to counterbalance the compressive forces of gravitation. The core contracts and its temperature increases until, by highly endoergic reactions, Fe-group nuclei photodisintegrate into c~-particles and neutrons. The star will then collapse catastrophically (see Meyer, 1997, for a recent review), generating a shock wave in the star's outer layers (a socalled Type II supernova event; see Fig. 1.2-1-Type II supernovae are distinguished from those of Type I by the presence of H-emission lines in their spectra), during which r-process nucleosynthesis reactions generate the heavier neutron-rich nuclides (see Table 1.2-1). The discovery in some meteorites of the radioactive decay products of short-lived nuclides synthesised during supernovae events has placed important constraints on the mechanisms and timing of formation of our solar system. The first report of the presence of 129Xe from the radiogenic decay of the short-lived, now extinct, nuclide 129I was made by Reynolds (1960; see also Jeffery and Reynolds, 1961). The presence of a num-

Table 1.2-1. Short-lived radioactive parent-daughter decay systems that may be detected in early solar system materials, and their stellar production sites (adapted from the Chart of the Nuclides, Goswami (2000) and references cited therein) Half-life Radioactive Synthesis Stellar Decay Daughter nuclide

processes

production site

processes

nuclide(s)

26A1 36C1 41Ca 44Ti 53Mn 60Fe 92Nb 99Tc 107pd 129I 135Cs 146Sm 182Hf 205pb 244pu 247Cm

p, EC s EC EC, p r, EC r p s s, r r r p, EC, o~ r, EC, ~ s r,c~ r, ot

SN, N, AGB, W-R SN, AGB, W-R SN, AGB, W-R SN SN SN, AGB SN SN, W-R, RG SN, AGB, W-R SN AGB, W-R SN SN AGB, W-R, RG SN SN

EC /4EC EC,/3 + EC 2/3EC /3/-3/4,6~ 2/-3EC SF 3o~, 2/3-

26Mg 36Ar* 41K 44Ca 53Cr 60Ni 92Zr 99Ru* 107Ag 129Xe 135Ba 142Nd 182W 205T1" 132,4,6Xe (238U) 235U*

0.717 My 0.301 My 0.103 My 59 yr 3.74 My 1.50 My 34.7 My 0.21 1 My 6.50 My 15.7 My 2.30 My 103 My 9.0 My 1.5.3 My 81 My 15.6 My

Nuclide synthesis and decay processes: EC = electron capture, s = slow neutron capture, r = rapid neutron capture, p = "p-process'; photodisintegration of the products of the s-process and proton capture, fl- = electron emission,/3 + = positron emission, SF = spontaneous fission, ~ = alpha capture/decay. Stellar production sites: SN = Types Ib, Ic and II (core-collapse) supernova, N = nova, AGB = asymptotic giant branch star, W-R = "Wolf-Rayet" star, RG = Red Giant star. *Evidence of radioactive decay from short-lived parent radionuclide in presolar or early solar system materials as yet unconfirmed.

6

Chapter 1: The Early Earth

Fig. 1.2-1. SN1987A in the Large Magellanic Cloud 170,000 light-years away, imaged using the Wide Field and Planetary Camera 2 (WFPC2) aboard the Earth-orbiting Hubble Space Telescope. The remains of the supernova star and the material thrown off when the star exploded are visible as the bright object in the centre of the inner ring. The three rings in this false-colour image are illuminated by emissions from atomic H (red) and doubly-ionised O (yellow). The rings have been interpreted to lie on the surface of an hourglass-shaped bubble of gas, created by the interaction of stellar winds from the star as it evolved from a red to a blue supergiant, long before the supernova event. The bright central ring, with a diameter of 1.5 light-years, represents the waist of the hourglass, made to glow by radiation from the supernova explosion. The origin of the rings is uncertain; one theory suggests that the outer rings were formed by a pair of wobbling, narrowly directed jets emerging from the vicinity of the central star and slamming into the hourglass walls. Nebular structures found around other dying stars indicate that such jets and the death of stars are intimately associated. Courtesy of Chris Burrows (Space Telescope Science Institute), the WFPC2 Science Team, and NASA.

1.2. Earth's Formation and Fir~t Billion Years

7

ber of other short-lived nuclides within the earliest-forming phases of primitive meteorites has since been confirmed (Table 1.2-1; see also Podosek and Nichols, 1997). The inventory of the short-lived nuclides calcium-41 (41Ca), 26A1,6~ and palladium-107 (107pd) could have been generated within a nearby Asymptotic Giant Branch (AGB) s t a r u a star of up to c. 8 solar masses with an inert C- and O-rich core, He-burning inner shell and an outer H-burning shell from which most of the star's energy will be derived (Busso et al., 1999). However, others such as manganese-53 (53Mn) and 1291, could not have been produced within AGB stars, but can only be created within massive, short-lived stars by the r-process during supernovae events. Their presence within the earliest-formed materials of our solar system requires that a Type II supernova explosion occurred in the vicinity and within 2 My of solar system formation (Cameron and Truman, 1977). Alternative explanations, that the short-lived nuclides were produced by spallation interactions between suitable target nuclei and high-energy cosmic rays, are not consistent with the abundances of these nuclides (particularly the observed relative abundances of and correlations between the nuclides 41Ca and 26A1) found in early-formed materials (e.g., Wasserburg and Arnould, 1987; Cameron, 1995; Wasserburg et al., 1996; Sajipal et al., 1998; Lee and Halliday, 2000a). Nor does the evidence favour the interpretation (Clayton, 1982, 1986) that the decay products of these nuclides were carried by microscopic (or "fossil") phases formed in the interstellar medium long before formation of the meteorite inclusions (see MacPherson et al., 1995). Current evidence favours the synthesis of the short-lived nuclides with atomic masses > 140, along with a proportion of the heavy elements within our solar system, during a core-collapse supernova event c. 4571 Ma ago (Lugmair and Shukolyukov, 2001; Gilmour and Saxton, 2001). Some short-lived nuclides with atomic masses < 140 (such as l~ and 1291) were present in early-formed solar system materials in much lower abundance than anticipated by this scenario, and it has been proposed that these radionuclides were mostly synthesised during earlier supernovae events (Wasserburg et at., 1996; Cameron, 1998) or by a second r-process mechanism (Qian and Wasserburg, 2000). Nevertheless, these nucleosynthesis events mark the time at which a proportion of the unstable isotopes that are widely used for geochronology were newly synthesised and commenced their radioactive decay. The shock waves that originated from the last supernova explosion to occur in the vicinity could have triggered the rapid collapse of a more slowly evolving molecular cloud of interstellar dust and gas nearby (Cameron and Truman, 1977; Foster and Boss, 1996), thus initiating the formation of our own Sun and solar system. Based on the relative abundance of the short-lived nuclides found within the earliest preserved remnants ot" our solar system, a 10-solar-mass supernova located between 2 to 10 parsecs distant, may have generated the newly-synthesised heavy elements within, and probably also triggered formation of, our solar system (Cameron et at., 1995, 1997).

Formation of Our Solar System Due at least in part to the lower mass of its precursor molecular cloud, the development of our solar system was fortunately very different from that outlined above for supernovae.

8

Chapter 1: The Early Earth

Smaller stars of up to a few solar masses have longer pre-main sequence histories, and substantially longer lifetimes, than more massive stars. The interstellar cloud from which our solar system formed was derived from the ejecta of a range of stellar sources, including red giants and supergiants, AGB, nova, supernova and possibly also Wolf-Rayet stars (massive, high-temperature stars with extremely high mass-loss rates). A nearby supernova event injected newly synthesised, short-lived nuclides into the interstellar cloud and triggered its collapse to form a proto-Sun with radius about 5 times that of the present Sun over a period of < 105 years (see Cameron, 1995). Collapse will have occurred progressively from the inner to the outer part of the cloud, with conservation of angular momentum causing the collapsing cloud to spin faster. Collisions of dust and gas particles orbiting the proto-Sun in the same direction caused the loss of their energy, resulting in the flattening of the cloud, particularly near the centre where the densities are highest. Rotation of both the disk and the proto-Sun around a common centre of mass generated spiral density waves in the surrounding nebula. Within the evolving nebula, gravitational energy can be converted to heat during collapse and can initially be radiated away, so temperatures initially decrease with increasing distance from the cloud core. However, as the density of the cloud increased, heat could not be lost efficiently. At some time within 105 years from the onset of collapse of the cloud core, the proto-Sun commenced the violent early H-burning (or T-Tauri) phase of its evolution (Cameron, 1995). The infall rate of material from the rotating accretion disk increased episodically during this phase, although a significant proportion of the mass of the Sun was lost or recycled back into the disk via energetic bipolar outflows emitted along the axis of rotation. Magnetic field instabilities emanating from the Sun and vigorous flares, violent eruptions and strong stellar winds will have caused turbulence and mixing of ionised gas and dust within the accretion disk. Temperatures within the accretion disk will have changed dynamically as the disk evolved, with parts shielded from the increasing temperatures of the inner nebula by the increasing density of the accumulating dust and gas closer to the cloud core and near the midplane of the disk (see Meibom et al., 2000). Where temperatures dropped below 1500 K, high-temperature refractory elements, such as Ca, Ti (titanium) and A1, condensed to form fragile "fluffy" sub-micron sized grains. These dust grains collided and accreted to form precursors of the Ca- and Al-rich inclusions (CAIs) of primitive meteorites. These inclusions are composed of a variety of high condensation temperature minerals, such as corundum (A1203), perovskite ([Ca, Na, Fe, Ce]TiO3), melilite ([Ca, Na]z[Mg, Fe, AI, Si]307), hibonite ([Ca, Ce][AI, Ti, Mg]12019) and spinel ([Mg, Fe, Ni, Cr]A1204). The more abundant Fe, Ni (nickel), and silicate-rich components condensed within parts of the nebula at lower temperatures, whereas volatile components such as water, ammonia, and methane ices, condensed only in the cold outer regions of the accretion disk. Volatile components in the inner solar system may have been carried to the outer regions of the solar system as ionised gas and dust by the solar wind (Shu et al., 1994). The Sun would have settled into the hydrogen-burning "main sequence" phase within 3-30 My of the onset of collapse of the cloud core (Strom et al., 1993; Cameron, 1995).

1.2. Earth's Formation and First Billion Years

9

Mechanisms and Ttime Scales o f Early Condensation

Spectroscopic observations suggest that, for those young solar-type T-Tauri stars that have accretion disks (so-called "classical" T-Tauri stars), such disks persist for only a few million years from the time of star formation (Podosek and Cassen, 1994; Bertout and M6nard, 2001; Briceno et al., 2001; Throop et al., 2001). Such brief time scales for planetary formation are also suggested by mathematical simulations (Wetherill and Stewart, 1993), which indicate that progressive aggregation of non-volatile components into planetary-sized bodies will occur within <~ 105 years. Simulations also indicate that gas drag would cause planetesimals smaller than about 1 km in diameter formed before the nebula had dissipated to lose angular momentum and spiral into the Sun in time frames of 105 years (Weidenschilling, 1977). Only larger planetesimals would avoid this fate. Because coagulation arising from collision between icy particles is more efficient than that between metal or silicate particles, coagulation of icy interstellar grains within an ice sublimation belt in the cold outer parts of the nebula led to the early formation of large gas-rich giant proto-planets, the precursors to Jupiter and Saturn, before the nebula gas had dissipated (Cameron, 1995). Evidence of hypervelocity impacts preserved within chondrites has also been interpreted to indicate formation of massive planets within ~< 2 My of solar system formation (Hutchison et al., 2001). The rapid formation and large mass of Jupiter was to have a profound influence on the subsequent evolution of the terrestrial planets within the inner accretion disk (cf. Chambers and Wetherill, 1998). Planetesimals in the inner part of the disk formed at a slower rate than those within the ice sublimation belt, by the collision and amalgamation of chemically refractory dust particles. As the planetesimals became progressively larger, the random component of their heliocentric velocities progressively decreased. Their lower relative velocities further enhanced their rapid growth by collision and accretion (see Wetherill, 1994). Such collisions resulted in shock melting, brecciation and accretion of part of the masses of smaller planetesimals to the larger bodies. These collisions also triggered large-scale melting and magmatic differentiation within the silicate components of the larger planetesimals. Although planetary-sized bodies had formed in the inner part of the accretion disk within <~ 2 My of the onset of collapse of the precursor molecular cloud (see below), these probably continued to be extensively reworked by late-stage collisions for at least a further 100 My. The presence of decay products from the short-lived nuclides, particularly 26A1, 41Ca, 53Mn, 6~ 107pd, 1291and hafnium-182 (182Hf; see Table 1.2-1), within primitive meteorites affirms that these meteorites must have formed within a few million years of the synthesis of these unstable radionuclides in the supernova event. The abundances of the decay products of many short-lived nuclides have been shown to be correlated commonly with the abundances of their parent elements (see, e.g., MacPherson et al., 1995, for the case of 26A1), confirming that the daughter products were not inherited as "fossil" components from interstellar dust but rather, that the radioactive decay of their parent nuclides occurred within the mineral phases in which the daughter products reside (see Fig. 1.2-2). Because they are now extinct, short-lived nuclides cannot be used directly to obtain absolute dates relative to the present-day, but their short half-lives allow their use to investi-

10

Chapter 1: The Early Earth

0.160| /

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S,ine. ree.s.anO,

F

+ ano hi,e

~ 0.150~

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me,i,te

~0"~~0.145It

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27AI/24Mg Fig. 1.2-2. 26AI-26Mg isochron diagram (from Hsua et al., 2000) for a spinel-free island (SFI#3) within CAI 5421 from the Allende chondrite meteorite. Spinel-free islands are the earliest recognisable phases in CAIs and may represent clasts that were entrapped in a liquid of CA/composition that crystallised around them. Melilite-rich mantles were commonly formed around these phases, which are the principal phases that condensed early from the solar nebula. This diagram is like a conventional isochron diagram except 27A1 is used on the X-axis instead of the parent isotope 26A1, because all 26A1 is now extinct, having been converted to 26Mg. The slope of array has no age significance, but can provide the initial 26A1/27A1 ratio at the time the sub-system was isolated from a common reservoir. Six anorthite measurements were taken from a single SFI exhibiting a moderate range of A1/Mg ratios (200-300) and excesses that are well correlated with its AI/Mg. A range of initial 26A1/27A1compositions have been determined for Allende CAIs, indicating that the meteorite consists of an aggregation of material formed over a range of times early in the history of the solar system.

gate the relative chronology of the earliest history of the solar system. The abundance ratio of a short-lived parent nuclide relative to a stable nuclide of the same element (for instance, 26A1/27A1) in a chemically undisturbed mineral phase is a function of both the abundance of the parent nuclide in the starting material (which will have been the solar nebula for the most refractory and therefore earliest condensing phases of meteorites) and the time at which the mineral formed. This ratio can be calculated from the slope of the correlation between the measured daughter nuclide/daughter element abundance ratio (i.e., 26Mg/24Mg) and the measured parent/daughter element ratio (27Al]24Mg). An absolute time scale for the radioactive decay of the parent nuclide can be obtained by the determination of the absolute age of formation of a host mineral phase that has a well-constrained parent nuclide

1.2. Earth's Formation and bTrst Billion Years

11

abundance. The uranium-lead (U-Pb) radiogenic decay schemes have been most successfully applied to the precise determination of the absolute chronologies of mineral phases bearing daughter products of extinct nuclides (see below). Provided an absolute time scale for the radioactive decay of the parent nuclide can be established, absolute chronologies for other host phases can be obtained by comparison of their initial 26AI/Z7AI ratios with that calculated for the dated host phase. Chronologies determined using such initial ratio comparisons are, however, based on the assumption that both parent and daughter nuclides were distributed homogeneously, both throughout the regions in which the phases were formed and for the duration over which the comparison is made. If some or all of the inventory of short-lived nuclides was introduced from a nearby supernova, which also triggered the rapid collapse of the precursor molecular cloud, complete mixing of the solar nebula prior to the onset of condensation seems unlikely. The preservation of "presolar" isotopic anomalies in sub-micron sized corundum (A1203), diamond, graphite, silicon carbide (SC) and silicon nitride (Si3N4) grains (see Zinner, 1997; Ott, 2001, for recent reviews) and of 160 heterogeneities in CAIs (Young and Russell, 1998; Yurimoto et al., 1998; however, see Thiemens, 1999, for an explanation invoking mass-independent chemical fractionation for such O isotopic variation) also indicates that the nebula did not reach complete chemical homogeneity before planetesimals started to condense. Moreover, although the manganese-53-chromium-53 (53Mn-53Cr) chronologies determined for many meteorite classes are consistent with other evidence (see Carlson and Lugmair, 2000), 53Mn-53Cr isotopic trends between meteorite classes suggest that the 53Mn/54Mn ratio within the inner solar nebula may have varied radially with distance from the Sun (Lugmair and Shukolyukov, 1998). Violation of the assumption of initial chemical homogeneity throughout the entire solar nebula prior to and during early condensation may introduce significant but unquantifiable uncertainties into chronologies based on initial ratio comparisons and such chronologies should be treated as indicative only, unless supported by additional evidence. The radioactive nuclide 26A1 decays to 26Mg with a half-life (tl/2) of 0.730 My (see Table 1.2-1). The presence of live 26A1 in the early solar system was first demonstrated by Gray and Compston (1974) and Lee and Papanastassiou (1974) for refractory CAIs of the Allende carbonaceous chondrite. Lee et al. (1976) reported a range of 26Mg/Z4Mg values for minerals from Allende CAIs that were correlated with AI/Mg. Recent studies (for a summary, see Podosek and Cassen, 1994) have confirmed that many CAIs had similar initial 26AI/27A1 ratios (c. 5 x 105) and can be inferred to have formed within a short interval of < 105 yrs duration (see Fig. 1.2-2). The identification of excess potassium-41 (41K) from the in situ decay of the parent nuclide 41Ca (tl/2 c. 0.1 My) within high Ca/K phases in CAIs from the Efremovka carbonaceous chondrite (Srinivasan et al., 1996) also requires that some CAIs formed within <~ 1 My of the supernova event. An absolute age of 4571 Ma has been determined for CAI formation (see Lugmair and Shukolyukov, 2001 ). This date therefore corresponds both to the time of the supernova event and also formation of the earliest preserved millimetre- to centimetre-scale materials formed within our solar system.

12

Chapter 1: The Early Earth

Chondrules are millimetre-sized spherical objects found in chondritic meteorites, that were formed by flash heating to 1700-2100 K within the solar nebula (see, for example, Shu et al., 1996; Alexander et al., 2001, for recent reviews). The majority of chondrules examined lack any evidence of live 26A1, implying that the oldest chondrules formed at least several million years later than the time of CAI formation (Carlson and Lugmair, 2000). Evidence of live 26A1, but with generally lower initial 26A1/Z7A1 ratios than in CAIs, has been found in only a few chondrules (Hutcheon and Hutchison, 1989; Zinner and G6pel, 1992; Russell et al., 1996; Galy et al., 2000). Comparison of the initial 53Mn/55Mn ratios of chondrules from the Chainpur and Bishunpur chondrites (Nyquist et al., 1997, 1999) with those of CAIs (Birck and All6gre, 1988) is also consistent with a time of formation of the oldest chondrules 2-6 My later than CAI formation (see Carlson and Lugmair, 2000). Similar comparisons using initial 53Mn/55Mn ratios indicate that enstatite chondrites formed between 9 and 24 My later than CAIs (Birck and All6gre, 1988; Shukolyukov and Lugmair, 1999; see Carlson and Lugmair, 2000).

Mechanisms and Time Scales of Planetary Differentiation For planetesimals larger than a few tens of kilometres in diameter, internal gravitational pressure and heating from radioactive decay may be sufficient to cause partial melting within their interiors, resulting in melt segregation and igneous differentiation. The eucrite and augite meteorite classes are remnants of such differentiated planetary bodies. A Pb-Pb isochron date of 4557.8 -+- 0.5 Ma was determined by Lugmair and Galer (1992) for the time of igneous differentiation processes occurring within the angrite parent body. The difference in the initial 53Mn/55Mn ratio determined for the eucrite parent body compared to that of the angrite parent body indicates that the former differentiated 5.8 4-0.8 My prior to that of angrites (Lugmair and Shukolyukov, 2001 ). A date of 4564.8 • 0.9 Ma is therefore indicated for the time of igneous differentiation within the eucrite parent body (see Lugmair and Shukolyukov, 2001). These dates indicate that planetesimals at least several tens to hundreds of kilometres in diameter (the probable howardite-eucrite-diogenite parent body, the asteroid 4 Vesta, is thought to have had a radius of c. 250 km) had formed and had undergone internal magmatic differentiation less than c. 10 My after the supernova event and formation of CAIs. A time difference of ~< 10 My between formation of CAIs and igneous differentiation within the eucrite and angrite parent bodies is also consistent with the initial 26A1, 53Mn and 182Hf abundances calculated for eucrites (Lugmair et al., 1995; Srinivasan et al., 1999; Quitte et al., 2000; Kline et al., 2002) and the higher initial 87 Sr/86Sr ratios of eucrites and angrites compared to CAIs (see Podosek and Cassen, 1994; Halliday and Porcelli, 2001). Early differentiation of metallic cores and silicate mantles within planetary bodies is also indicated by rhenium- 187-osmium- 187 (187Re-187Os), hafnium- 182-tungsten- 182 (ISZHf-182W) and palladium- 107-silver- 107 (l~176 isotopic data from the iron and enstatite meteorites. Formation of metallic cores within ~< 50 My of formation of the solar system is suggested by 187Re-187Os isotopic data from pallasites and iron meteorites (Morgan et al., 1995; Shen et al., 1998; Horan et al., 1998), although a large uncertainty

1.2. Earth's Formation and First Billion Years

13

of -+-3% in the half-life of 187Re renders absolute chronologies determined by this method imprecise. The short-lived nuclide 182Hf decays to 182W with a half-life (tl/2) of 9 My (see Table 1.2-1 ). Hf is a refractory lithophile element and will preferentially partition into silicates during metal-silicate melt segregation, whereas W is moderately siderophile and will be concentrated into metal phases. Depletions in l S2W relative to that within terrestrial silicates have been detected in iron meteorites and the metal fractions of ordinary chondrites (Harper and Jacobsen, 1996a; Horan et al., 1998; Lee and Halliday, 2000a, b), indicating segregation of low Hf/W metal phases early in solar system history and prior to complete decay of 182Hf. No differences in W isotopic compositions amongst the different iron meteorite groups have yet been resolved; calculated initial 182Hf/18~ ratios for metal phases within enstatite chondrites are also comparable with those determined for carbonaceous chondrites (Kline et al., 2002; Schoenberg et al., 2002a; Yin et al., 2002), indicating that the metal phases of differentiated meteorites segregated within <~ 2 My of the formation of the solar system. This is in general terms consistent with the l~176 evidence, which indicates a formation interval of < 9 My for most of the iron meteorites (Carlson and Lugmair, 2000). This interval corresponds to the time at which the temperatures of the parent bodies of the iron meteorites fell below the diffusional blocking temperature for Pd, Ag, Hf and W, and so corresponds to the time at which these planetesimals had either ceased to differentiate or were disrupted by impacts (Carlson and Lugmair, 2000). Intriguing evidence of early differentiation involving hydrous fluids within planetesimals is provided by the identification of monoisotopic 129Xe, from the radiogenic decay of the extinct parent nuclide 129I, within halide crystals in the chondritic meteorite Zag (Whitby et al., 2000). These findings indicate that aqueous fluids circulated, presumably as part of hydrothermal systems, within planetesimals <~ 2 My after solar system formation. Further evidence in support of the operation of early aqueous alteration processes in planetesimals has been provided by 53Mn-53Cr isotopic investigations of carbonate (CaMg[CO3]2) fragments from the carbonaceous (CI) chondrites Orgueil and Ivuna (Endress et al., 1996) and of fayalite from the altered (CV3) chondrite Mokoia (Hutcheon et al., 1998), which indicate pervasive aqueous alteration within 20 My of the time of solar system formation. Other studies (Young et al., 1999) have recently argued that the mineralogical and O isotopic variability observed within the carbonaceous chondrites is attributable to fluid-rock interaction within their parent body (or bodies) shortly after CAI and chondrule formation. Earth Accretion, Formation of the Moon and the Magma Ocean

It is apparent from the above discussion that planetary embryos (Mercury or lunar size, c. 1026 g) that were to eventually collide and combine in the final stages of formation of the Earth had themselves formed within <~ 5 My of the supernova event that triggered formation of the solar system. This short time scale for formation of planetary embryos indicates that some of the short-lived nuclides, particularly 26A1, were still present in sufficient abundance to provide an internal heat source, enabling planetary embryos to form and differentiate extensively into silicate mantle and metallic core components within this

14

Chapter 1: The Early Earth

time interval. The terrestrial planets were mainly derived by the collision and accretion of planetary embryos that formed within a c. 0.5-2.5 AU range of the Sun (Wetherill, 1994), but major contributions of the more volatile components could have been sourced from the more volatile-rich asteroid belt and outer parts of the solar system during the later stages of accretion of the terrestrial planets. It is also apparent that liquid water was present during accretion of the terrestrial planets and would almost certainly have played a major part in the early differentiation of the planetary embryos that ultimately were to become part of the Earth. Provided accretion was rapid, the earliest atmospheres of these planetary embryos would have been acquired from the solar nebula before its dissipation and may have been close to solar in composition. Atmospheres of those planetary embryos formed closer to the Sun were probably more depleted in volatile components such as water, ammonia, and methane ices compared to those formed more distant from the Sun. However, the final chemical compositions and internal structures of the terrestrial planets would have been determined largely by the last major collisions that occurred between planetary embryos. Statistical simulations of the final stages of planetary accretion in the terrestrial zone (Chambers and Wetherill, 1998; Agnor et al., 1999) predict the formation and eventual collision of a small number of massive planetary bodies. Whatever the chemical nature or degree of differentiation of the last major planetary bodies that were to collide and ultimately form the Earth, there is little doubt that they were to be transformed irreversibly by the process that was to bring about their amalgamation. Although not universally accepted (see, e.g., Jones and Palme, 2000; Ruzicka et al., 2001; Dreibus and W/~nke, 2002, for recent critical analyses), the most widely accepted current hypothesis for the origin of the Earth's Moon argues that it originated from the impact of a planetary embryo ("Theia") of at least the mass of Mars (Mars mass is c. 11% of present-day mass of the Earth) with the proto-Earth, after the latter had accreted to about 50-90% of its present mass (Giant Impact Hypothesis; see Stewart, 2000, and references therein) (see also, section 5.9). This model is generally consistent with the angular momentum of the Earth-Moon system, with stochastic planetary accretion models (Canup and Angor, 2000, and references therein), with the extreme depletion in volatile elements found in the Moon and with the alignment of the O isotopic compositions of lunar samples along the terrestrial O isotope fractionation line (Weichert et al., 2001). The best available estimate of the time of this possible collision, between 25 and 35 My after solar system formation (i.e., between about 4550 and 4540 Ma ago), has been provided by 182Hf-182W model dates obtained on lunar rocks interpreted to be derived from different sources within an isotopically stratified lunar mantle (Kline et al., 2002; Yin et al., 2002). This estimate is unaffected by the recent recognition (Lee et al., 2002) that ISZw in some lunar samples is of cosmogenic origin, and is consistent with the oldest dates obtained from lunar samples (Swindle et al., 1986; Snyder et al., 2000). Determination of the exact time of impact is dependent on poorly-constrained variables such as accretion rates, the possible secular variation in the oxidation state (as pointed out by Ringwood, 1966, the presentday silicate mantle is too oxidised to be in equilibrium with its metallic iron core) and W metal-silicate partition coefficient within the bulk silicate Earth, and the ratio of the masses of the proto-Earth and impactor. But the W and Pb isotopic evidence is consistent

1.2. Earth's Formation and First Billion Years

15

with collision between the proto-Earth and an impactor 30 My after formation of the solar system, with a proto-Earth to impactor mass ratio of 7:3 and with between 65 to 90% of the current mass of the Earth-Moon system having been accreted at the time of impact (Halliday, 2000; Kline et al., 2002; Yin et al., 2002). This formation time for the Moon is generally consistent with (comparatively imprecise) estimates based on 1291-244pu-Xe (Pu, plutonium) isotopic constraints (reviewed by Swindle et al., 1986). A wide range of other plausible scenarios is also consistent with the 1291-244pu-Xe, 182Hf/182W, O and U-Pb isotopic constraints, however. An initial 182W/184W ratio of the bulk Moon within uncertainty of the bulk silicate Earth, inferred from the least radiogenic 182W/184W value measured so far for lunar samples (D.-C. Lee et al., 1997), is consistent with its derivation from the silicate portion of the proto-Earth and the impactor. Collision of a Mars-sized body with the proto-Earth will have caused shock melting of a significant part of the mass of at least one hemisphere of the proto-Earth. Rapid isostatic readjustment of the remaining mass of the proto-Earth mantle will have occurred immediately after impact, followed over the next c. 103 years by the propagation of a solidification front towards the surface (Solomatov, 2000). A low-viscosity, turbulently-convecting shallow magma ocean formed rapidly over a probably subsolidus (depending on presently poorly-constrained initial conditions) lower mantle, but the magma ocean will have cooled rapidly to a more viscous, low melt fraction magma ocean consisting of 20 to 30% partial melt within c. 1 My of the impact (Abe, 1993). Segregation of metal from silicate phases would have been facilitated by formation of a magma ocean, promoting the release of gravitational energy associated with further growth of the metallic core. The proto-Earth atmosphere was probably severely depleted as a consequence of the impact (Chen and Ahrens, 1997). Silicate vapour and steam released into the atmosphere by the impact may have slowed the rate of heat loss by surface radiation, but solidification of a thin surface layer overlying the magma ocean will have occurred rapidly, possibly over a period of tens of years (Mukhin and Pinenov, 2002). A partially molten magma ocean in the upper part of the Earth's mantle may have persisted for up to 200 My after the impact (Abe, 1997).

Latter Stages of Earth Accretion and Early Earth Differentiation Two main areas of direct investigation have provided significant insights into the early differentiation history of the Earth; 182Hf-182W isotopic systematics, which have facilitated our understanding of the formation of the Earth's metallic core, and the noble gases, which have provided constraints on the evolution of the atmosphere and history of degassing of the Earth's mantle. Theoretical and experimental studies of the Earth's Moon and detailed investigations of the oldest preserved remnants of the Earth's surface crust are additional important areas of investigation. During large-scale differentiation of the Earth into silicate mantle and metallic core, W is expected to be partitioned preferentially into the metallic core, whereas Hf should be concentrated into the silicate mantle. As anticipated, the silicate component of the Earth has a Hf/W ratio of c. 15, substantially higher than the chondritic value of c. 1 (cf. Harper and Jacobsen, 1996a; Halliday and Lee, 1999). As outlined earlier, remnants of the early-

16

Chapter 1: The Early Earth

formed metallic cores of other planetary bodies, such as the iron meteorites and the metal fractions of ordinary chondrites, have low Hf/W ratios and are depleted in l sew relative to carbonaceous chondrites (Lee and Halliday, 1995; Harper and Jacobsen, 1996a; Lee and Halliday, 2000a, b; Halliday et al., 2000; Kline et al., 2002; Schoenberg et al., 2002a; Yin et al., 2002), indicating that these metal phases segregated within ~< 2 My of the formation of the solar system. If the Earth's core also formed rapidly when live 182Hf was present, it might be anticipated that the core would also be depleted in 1sew, in common with the iron meteorites. Provided that differentiation of the Earth into metallic core and silicate mantle components occurred rapidly and within a closed system, the high Hf/W of the silicate mantle resulting from preferential incorporation of W relative to Hf into the core should have resulted in the Earth's silicate mantle having a radiogenic W isotopic composition relative to the chondritic value. Measurement of the W isotopic compositions of terrestrial silicates (Lee and Halliday, 1996; see Halliday et al., 2000) initially suggested that the silicate mantle component of the Earth had a 182W/184Wvalue indistinguishable from the chondritic value. However, recent investigations (Kline et al., 2002; Schoenberg et al., 2002a; Yin et al., 2002) indicate that there were analytical problems with earlier measurements of 182w[lS4w for carbonaceous chondrites and have confirmed that terrestrial silicates have significantly more radiogenic W isotopic ratios than carbonaceous chondrites. Based on the difference in W isotopic compositions of terrestrial silicates and carbonaceous chondrites, these studies have argued that the Earth's metallic core must have formed within 35 My of formation of the solar system. However, metal-silicate segregation within planetesimals may have occurred significantly more rapidly than the rate of planetary accretion. The rate of terrestrial core growth was therefore probably largely limited by the rate of planetary accretion, rather than the rate of planetary differentiation and metal-silicate segregation. If this was the case, 182w/lS4w values in terrestrial metal and silicate components will largely be a function of the degree of mixing of metal and silicate components derived from the impactors and proto-Earth, following partial rehomogenisation of these components as a result of late impact events and prior to metal-silicate resegregation. As a consequence, 182Hf_182W model ages are unlikely to represent meaningful estimates of the time of terrestrial core formation. If the terrestrial planets were accreted rapidly enough to have captured atmospheres derived from the nebula, it is likely that these early atmospheres were subsequently removed by intense ultraviolet luminosity and solar winds generated during the Sun's T-Tauri phase. However, such early solar atmospheric gases may have been trapped by dissolution in extensive magma oceans if these existed on the surfaces of the terrestrial planets prior to nebula dissipation. The relative abundances and isotopic compositions of the noble gases have been applied successfully to investigate this possibility. The element helium (He) consists of two isotopes, 3He and 4He, of which 4He is the by-product of c~-decay. Neon (Ne) consists of three isotopes, ZONe, 21Ne and 22Ne, of which 21Ne and 22Ne are produced principally by nucleogenic (n, or) reactions on 180 and 24Mg, and 25Mg respectively, from local decay of U and thorium (Th; although 21Ne may also be generated by spallation reactions on Mg, Na, Si and A1). Because both 4He and 21Ne are generated by the radioactive decay of U and Th, variations in 21Ne arising from nucleogenesis within a closed system

1.2. Earth's Formation and First Billion Years"

17

will be correlated with 4He. Based on the relative elemental abundances of the noble gases and isotopic compositions of He and Ne, three principal componentsmsolar (inferred from the solar wind as measured in lunar surface materials), primitive meteoritic (or "planetary") and atmospheric--can be distinguished. Isotopic compositions of the noble gases He and Ne within fresh basaltic glasses from Hawaii have been interpreted as resulting from mixing of nucleogenic and atmospheric components with a solar component having high 2~ (c. 13.6; Honda et al., 1991). Solar-like Ne isotopic ratios have also recently been detected in basaltic glasses from Iceland (Dixon et al., 2000). The existence of solar-like Ne within the Earth requires preservation since the time of Earth's formation of source regions that have retained high [Ne]/[U+Th], so that the isotopic composition of trapped initial solar Ne is not modified by the addition of nucleogenic 21Ne generated by in situ decay of U and Th. The high 2~ values of some mid-ocean ridge basalt samples are also accompanied by low 38Ar/36Ar (argon) ratios (Pepin, 1998), consistent with the presence of a low 38Ar/36Ar, "solar wind"-like component. The identification of high 3He and "solar" Ne (high 2~ and Ar (low 38Ar/36Ar) components within the Earth supports the proto-Earth's early capture and dissolution within a magma ocean of a solar atmosphere prior to nebula dissipation. Basalts derived from the Earth's upper mantle, well gases and the present-day atmosphere contain excesses of 129Xe from the radiogenic decay of the extinct parent nuclide 129I, and of the heavier isotopes of Xe (particularly 136Xe) from fission of 238U and of the extinct nuclide 244pu (Kunz et al., 1998; see Table 1.2-1). The presence of these excesses provides further confirmation of the rapid accretion of much of the mass of the Earth before the complete decay of the short-lived parents 129I and 244pu. Excesses of 129Xe in midoceanic ridge basaltic glasses are correlated with 4~ (Staudacher, 1987). As 4~ is generated by radiogenic decay of 4~ the observed correlation of 129Xe with 4~ in ridge basalt glasses is consistent with the rapid early depletion of incompatible elements (including K) and the degassing of the noble gases in the Earth's upper mantle during the decay of 129I. Rapid degassing and early depletion of the noble gases within the upper mantle led to high K/Ar and I/Xe ratios and to correlated enhancement in 4~ and 129Xe. Estimates of the terrestrial 129Iand atmospheric 129Xe inventories indicate that the initial terrestrial 1291/127Iratio was more than 2 orders of magnitude lower than that of the prevailing solar system ratio and indicate a 129Xe retention formation interval for the Earth of between 60 to 110 My after formation of the solar system (Swindle et al., 1986; Kunz et al., 1998; Podosek and Ozima, 2000; Porcelli et al., 2001). This Earth formation 1291-129Xe date has generally been interpreted as the average time of closure of the Earth to Xe loss (Swindle et al., 1986), although its exact chronological significance is unclear. Jephcoat (1998) suggested that 129I,and its "missing" daughter 129Xe, may have been occluded into the Earth's deep mantle and core during the early stages of accretion, whereas Podosek and Ozima (2000) speculated that 129Xe is missing because of the loss of volatiles during violent impact degassing late in the accretion history of the Earth. On the basis of the Earth's noble gas characteristics, Harper and Jacobsen (1996) argued that the Earth accreted in two stages. During the first stage, solar He and Ne (and also solar Ar) were incorporated into a magma ocean blanketed under a massive proto-atmosphere

18

Chapter 1: The Early Earth

of molecular hydrogen and helium prior to dissipation of the solar nebula. The "planetary" (i.e. meteoritic) compositions of the heavier noble gases were subsequently acquired by the late accretion of planetesimals after dissipation of the solar nebula during the second phase. This two-stage model also accounts for the oxidised state of the Earth's silicate mantle relative to the core. This interpretation was, however, disputed by Trieloff et al. (2000), who argued that the Hawaiian and Icelandic plume (see also, sections 3.2 and 3.3) components have 2~ ratios more consistent with a meteoritic solar (Ne-B) component with 12.52 • 0.18 rather than a solar component with present-day solar wind ratio of 13.80 -+-0.10. These authors argued that the Earth's solar-type rare gas inventory was inherited from the small planetesimals that accreted to form the Earth, following their irradiation by solar wind after the nebula had dissipated. From the above discussion, it is clear that the processes responsible for the early differentiation of the Earth into metallic core, silicate mantle, chemically differentiated crust and atmosphere overlapped in time and operated in unison with, and were even closely linked to, the progressive accretion of the planet. Earth's Earliest Crust

Given the rapid time scales of formation of the Earth, its later-stage accretion by the collision of a small number of massive planetesimals, and its internal heating by radioactive decay and release of gravitational energy associated with core formation, there seems little doubt (regardless of the applicability of the Giant Impact Hypothesis) that the Earth passed through a "magma ocean" stage in its accretion history. Insight into the nature of the first surface crusts formed under such conditions may be inferred from observations of the crusts of the Earth's Moon and other terrestrial planets. Because of its lower mass, the Moon lost its internal heat rapidly, so its earliest crust is relatively well preserved. The oldest recognised lunar crust is preserved as 4530 Ma ferronian anorthosites that are generally considered to represent flotation cumulates derived from a differentiated residual global lunar magma ocean (see Snyder et al., 2000; Shearer and Floss, 2000; see also O'Hara, 2000, for an alternative view). Between about 4200 and 2500 Ma, lunar mare basalts were episodically erupted with very high effusion rates, filling multi-ring and irregular-shaped depressions on the Earth-facing hemisphere of the Moon up to a thickness of 4.5 km (see Shearer and Floss, 2000). These melts are considered to have been derived by decompression melting of mafic cumulate sources, with the depth of source melting becoming progressively deeper in the lunar mantle with time (Nyquist et al., 1995; see Snyder et al., 2000). The masses, cooling rates and the time scales of crust formation of Venus, the Earth and the Moon were very different, yet there are similarities in their major physiographic features. Based on counts and sizes of impact craters, Venus was completely resurfaced comparatively late in its historymthe average surface age of Venus is estimated to be about 750 Ma (McKinnon et al., 1997). Three major physiographic features---crustal plateaus, volcanic plains and volcanic risesmmay be distinguished over > 80% of the surface area of Venus (Phillips and Hansen, 1998). Crustal plateaus are postulated to have formed by large degrees of plume melting beneath thin lithosphere, followed by lithospheric thick-

1.2. Earth's Formation and First Billion Years

19

ening by magmatic intrusion and underplating. The widespread volcanic plains may have formed over broad areas of mantle upwelling associated with mantle convection. Although volcanic plains probably formed throughout the time of crustal plateau formation, Phillips and Hansen (1998) proposed that Venusian plains volcanism culminated at about 700 Ma, when the stress state of the lithosphere switched from a "mobile-lid" state, during which some form of plate-tectonic crustal recycling may have operated, to a rigid lithosphere "thick-lid" state. Volcanic rises are considered to be younger features formed over a thicker lithosphere by lower degrees of plume melting (Phillips and Hansen, 1998). A partially molten magma ocean may have persisted in the upper part of the early Earth's mantle for several hundred million years. Surface temperatures were probably so high at this time that water resided mostly in the atmosphere as steam rather than in the oceans, with about 1 wt.% H20 dissolved in the magma ocean (Abe, 1993). Much of the Earth's accreted volatile inventory residing in the early atmosphere may have been lost by impact erosion and hydrodynamic escape, whereas a substantial proportion of the presentday atmosphere and oceans may have been reacquired from comets following accretion (cf. Delsemme, 1995). There are few reliable constraints on the magnitudes and rates of these competing processes. Depending on the effectiveness of the atmosphere's blanketing on the rate of heat loss from the Earth's surface, a thin solid surface layer could have formed over the magma ocean within a period of tens of years (Mukhin and Pinenov, 2002). As was the case for the Moon and the other terrestrial planets, volcanic plains are inferred to have been a dominant physiographic feature of the earliest surface crusts of the Earth. Plagioclase is stable only at depths less than about 40 km and will sink in wet terrestrial magmas (Taylor and McLennan, 1985), so early terrestrial crusts are likely to have been largely komatiitic and/or high-Mg basaltic, rather than anorthositic as on the Moon. Thin, transitory high-Mg basaltic and komatiitic crusts derived by very large degrees of partial melting of both shallow and deeper upwelling mantle sources may have been melted, fragmented and reformed many times over the surface of a chaotically convecting, partially molten mantle (see also, section 3.6). On cooling and solidification, such early crusts were probably rapidly resurfaced by continuous, voluminous sheet-like volcanic flows triggered by new upwelling events. Such early Earth crusts would have had short lifetimes and were brecciated by meteorite impacts and rapidly recycled back into the mantle. Early terrestrial crusts were probably periodically recycled on a global scale, as a consequence of both major impacts and of convective mantle overturn (Nelson, 1998a; see section 3.4). The nature of the recycling mechanisms for the earliest mafic "thin-lid" plates back into the Earth's mantle is unknown (a speculative hypothesis is discussed in section 3.6). Mismatches between the rate at which new crust was created and its surface distribution over convective mantle upwellings, and the rate at which crust could be destroyed via an orderly process of recycling into the mantle, probably meant that crust recycling initially did not operate by one orderly mechanism as it apparently does on the Earth today. Because they were zones in which water and other volatiles were recycled into the mantle and the adjacent crust could be magmatically reworked by remelting, intrusion and differentiation, siliceous high metamorphic grade continental crust buoyant enough to resist recycling into the mantle probably formed within these recycling zones (cf. "plughole"

20

Chapter 1: The Early Earth

model in section 3.6). As cooling of the planet proceeded, the prevailing chaotic mantle convection regime probably gave way to transitory stable convective regimes during which stable spreading-ridge systems could form over regions of mantle upwelling for increasingly longer durations. Because of the higher mantle temperatures, the solidus of upwelling mantle was intersected at greater depths than is the case for the present-day Earth, resulting in the formation of substantially thicker and more magnesian oceanic crust during these early stable convective periods compared to that which was to be formed on a younger, cooler Earth (see Arndt and Chauvel, 1990). Crustal thickening by flood-basalt resurfacing also operated throughout the Archaean; the more than 15 km thick, predominantly high-Mg basaltic sequences of the c. 3480-3410 Ma Warrawoona Group of the Pilbara craton, Western Australia, represents one of the best preserved early Archaean examples. Tonalite, trondhjemite and granodorite melts were generated by partial melting within this early, hydrothermally altered mafic crust. The oldest identified terrestrial zircons could have crystallised within such felsic differentiates. Earth's Oldest Preserved Crustal Remnants

The oldest identified remnants of the Earth's early crust are 4400 to 4000 Ma grains of the common accessory mineral zircon (ZrSiO4), found both as detrital grains within c. 3050 Ma metasedimentary rocks (Froude et al., 1983; Compston and Pidgeon, 1986; Wilde et al., 2001; Wyche et al., in press) and as xenocrysts in c. 2700 Ma gneissic monzogranites (Nelson et al., 2000) of the Yilgarn craton, Western Australia. Major- and traceelement distribution patterns and the presence of fine euhedral growth zoning and siliceous inclusions within these ancient zircons suggest that they crystallised within moderately to highly silicic (i.e., granitic sensu lato) melts. They were probably derived from a 4400 to 3874 Ma composite terrane comprised of igneous and high-grade metamorphic rocks (see Nelson et al., 2000). Their ancient source rocks may survive in the northwest Yilgarn but have not been identified. Zircon will accommodate U but not Pb within its mineral structure and is thus suitable for U-Pb dating. Furthermore, zircon is highly resilient and can survive high-grade metamorphism, intensive detrital reworking and incorporation into silicic melts. As a consequence, the chemical, microstructural and U-Pb isotopic characteristics of individual zircon crystals may preserve a record of complex geological histories. Such histories have been investigated for several >/4000 Ma zircon grains by combined in situ U-Th-Pb isotopic microanalysis using the SHRIMP (Sensitive High-Resolution Ion MicroProbe), back-scattered electron and cathodoluminescence imaging and trace-element microanalytical methods (Nelson et al., 2000; Wilde et al., 2001; Nelson, 2002). Due at least in part to the higher abundance of the parent isotope 235U (tl/2 c. 700 My) prior to 4000 Ma and consequent higher accumulated abundance of the daughter isotope 2~ ion-microprobe analyses of typically 25 ~tm diameter sites on sectioned and polished surfaces of these ancient zircons can yield very precise 2~176 dates, with typical 95% confidence uncertainties of less than a few million years. Ion-microprobe analyses indicating recent loss of radiogenic Pb from the analysis sites are commonly obtained from sites with U greater

1.2. Earth's Formation and First Billion Years

21

than c. 400 ppm, consistent with analyses obtained on younger zircons from samples taken elsewhere throughout the Yilgarn craton (cf. Nelson, 1997). An unusual feature of the ~> 4000 Ma zircons so far examined in detail is the wide range of e~176 dates found within each grain. Some, but not all, of the range in e~176 dates obtained from individual grains is attributable to the overlap of the ion-microprobe analysis sites onto younger rim zircon (cf. Nelson et al., 2000). The oldest e~176 date obtained from each crystal has usually been interpreted as a minimum estimate of the time of igneous crystallisation of the zircon, with any younger dates not accounted for by the presence of younger rim zircon instead attributed to loss of radiogenic-Pb from micronscale domains within each grain (Nelson et al., 2000). Diffuse euhedral primary igneous zoning is evident in the cathodoluminescence and back-scattered electron images obtained for most/> 4000 Ma zircons so far examined in detail (see Figs. 1.2-3 and 1.2-4), indicating that the e~176 variation cannot be attributed to patchy recrystallisation of the zircon microstructure. Nor can the 2~176 variation be related to the network of fractures, which are later features (see Figs. 1.2-3 and 1.2-4). Such radiogenic-Pb loss could have occurred episodically, during a series of thermal events occurring in the interval 4405 to 3750 Ma, or during one or a few younger events occurring before c. 3750 Ma. Maas et al. (1992) subdivided the older U-Pb zircon analyses from the Mount Narryer quartzites into c. 4190 Ma and 4150 Ma sub-populations; both ages are represented within grain 11 from the Churla Well granitic gneiss examined by Nelson et al. (2000; see Fig. 1.2-3). In addition, e~176 dates of 4360, 4340 and 4318 Ma obtained from grain 10 from a sample (169075) of quartzite from the Southern Cross granite-greenstone terrane, central Yilgarn, closely match the dates of 4363, 4340 and 4320 Ma indicated by analyses obtained by Wilde et al. (2001) from their 4404 Ma grain W74/2-36 from the Jack Hills, northwest Yilgarn (see Figs. 1.2-4 and 1.2-5). This concurrence between e~176 dates obtained from within and between zircon grains favours the episodic loss of radiogenic Pb during a series of thermal events between 4360 and 3750 Ma. Significant Pb diffusion from the structure of zircon will occur during metamorphism when temperatures exceed 850~ (J.K.W. Lee et al., 1997; see Fig. 1.2-6), so the events recorded within these ancient zircons: may correspond to high-grade thermal episodes during which accumulated radiogenic Pb was lost from susceptible domains within these grains. It is conceivable that at least some of the concordant e~176 dates at c. 4360, 4340, 4320, 4185, 4150, 4005, 3978, 3945 and 3874 Ma recorded within these ancient zircons (see Fig. 1.2-7) correspond to events during which temperatures of their host rocks exceeded 850~ perhaps during episodic mantle upwelling or overturn episodes. Although it is not certain how representative their thermal histories are of the early Earth's near-surface environment, evidence of ancient radiogenicPb loss by diffusion from micron-scale domains preserved within these t> 4000 Ma zircons is consistent with near-surface temperatures exceeding 850~ as late as 4000 Ma. Wilde et al. (2001) and Mojzsis et al. (2001) reported 6180 values significantly higher than the value of c. 5.3%o expected for mantle-derived zircons for > 3900 Ma detrital zircons from the Jack Hills metasedimentary belt. It was argued that their high 6180 values indicated that these ancient zircons crystallised within granitic melts derived from the partial melting of high-180/160 sedimentary or low-temperature hydrothermally-altered

22

Chapter 1: The Early Earth

Fig. 1.2-3. Back-scattered electron (BSE) and cathodoluminescence (CL) images (adapted from Nelson et al., 2000) of grain 11 extracted from a leucocratic granitic gneiss (sample 105007: Churla Well), Yilgarn craton. The surface microstructure and SHRIMP analysis sites are visible on the BSE image. Note that because the zircon structure and composition has been modified by high-intensity electron bombardment during chemical analysis along linear traverses, these can be seen on the CL images as bright lines. Diffuse primary igneous zonation is evident in the CL image. Inset show zones within the grain which are distinbuishable on the basis of the BSE and CL images and petrographic investigation. An interpretive schematic structural diagram sununarising SHRIMP U-Pb dates, with + 1o" uncertainties, obtained for this grain is also shown.

1.2. Earth's Formation and First Billion Years

23

Fig. 1.2-4. Back-scattered electron (BSE) and cathodoluminescence (CL) images of grain 10 extracted from a quartzite from the Southern Cross Province of the Yilgarn craton (sample 169075: Kohler Bore). The surface microstructure and SHRIMP analysis sites are visible on the BSE image. Faint igneous zoning is apparent in the CL image and indicates that the grain is an end-fragment of a larger euhedrally zoned crystal. An interpretive schematic structural diagram summarising SHRIMP U-Pb dates, with -+-1cr uncertainties, obtained for this grain is also shown.

source rocks. Wilde et al. (2001) reported a 3180 value of 7.4 + 0.7%0 (2a uncertainty) obtained on an area of grain W74/2-36 from which they obtained the oldest 2~176 date of 4404 + 8 Ma (20- uncertainty). These authors argued that this required the presence of liquid water on the Earth's surface at 4404 Ma. A lower ~180 value of 5.0-+-0.7%0, indistinguishable from values obtained from mantle-derived zircons, was obtained from part of this grain which gave a younger 2~176 date of 4364 + 6 Ma (2a uncertainty).

Chapter 1: The Early Earth

24

(a) W74, grain 2-36

((n r "1o

_Z

o~ o~

.

0

.~_

4300

4320

4340 2~176

4360 4380 Age (Ma)

4400

4420

Fig. 1.2-5. Gaussian Summation probability density plots for >/-t-95% concordant SHRIMP U-Pb analyses older than 4.0 Ga obtained from (a) zircon W74/2-36 from a sample of conglomerate from the Jack Hills, Narryer Gneiss Complex (data from Wilde et al., 2001 ); (b) grain 10, quartzite sample 169075 from the Southern Cross Province of the central Yilgarn craton (data from Nelson, 2002).

Wilde et al. (2001) interpreted this apparently younger area as a magmatic overgrowth on a 4404 -t- 8 Ma, high 6180 core. However, from the images and maps shown in Wilde et al. (2001) and Peck et al. (2001), grain W74/2-36 is a grain fragment that crystallised during a single episode. The range in 2~176 dates obtained for this zircon is therefore likely to be due to episodic loss of radiogenic-Pb from micron-scale domains (cf. Nelson et al., 2000). A more detailed schematic diagram of grain W74/2-36 presented in Peck et al. (2001) shows that the high 6180 value of 7.4 was obtained from a site that gave a younger 2~176 date (4284 + 6 Ma) than the site from which the lower 6180 value of 5.0 was obtained (4364 + 6 Ma). This raises the possibility that the high 3180 value is of secondary origin and therefore does not provide evidence of the presence of liquid water on the Earth's surface at 4404 Ma. Experimental studies (Watson and Cherniak, 1997) indicate that oxygen diffusion in zircon occurs at a significantly higher rate and at lower temperatures than does Pb diffusion (see Fig. 1.2-6). If the observed variation in 2~176 ratios within individual/> 4000 Ma zircons is due to Pb diffusion at high temperatures, it is unlikely that magmatic oxygen isotopic compositions in these zircons would remain unaffected. In a morphological study of > 3900 Ma zircons from the Jack Hills, Pidgeon et al. (2001)

1.2. Earth's Formation and First Billion Years

25

Fig. 1.2-6. Diagram showing curves of diffusion radius (lxm) versus closure temperature (T~ for different cooling rates for Pb (from J.K.W. Lee et al., 1997) and O diffusion (Watson and Cherniak, 1997) in zircon. The observed 207pb/206pb variation within individual ~>4.0 Ga zircons may be attributed to Pb diffusion at a ~m-scale at T c. 850-950~ between 4405-3750 Ma. The shaded area corresponds to the estimated diffusion radius, based on a SHRIMP ion probe spot size of c. 25 ~m diameter, required to account for the observed 207pb/2~ variation in the ~>4.0 Ga zircons that have so far been examined in detail (Nelson et al., 2000; Wilde et al., 2001 ; Nelson, 2002). Oxygen diffusion in zircon under high temperatures and in the presence of H20 will occur far more rapidly than Pb diffusion in zircon. The observed 207pb/206pb (and ~;180) variation within individual >~4.0 Ga zircons is consistent with Earth's surface temperatures of c. 850-950~ at > 3.75 Ga.

also pointed out that most of these ancient grains lack morphological features, such as large central cores surrounded by strongly magmatically-zoned rims, consistent with their crystallisation within peraluminous melts derived from sedimentary precursors. The oldest rocks so far identified on the Earth's surface are 4030 Ma granitic gneisses of the Acasta Gneiss Complex of the Northwest Territories, Canada (Stern and Bleeker, 1998; Bowring and Williams, 1999). Granite-greenstone crust, consisting of thick (~> 10 km) arcuate or linear, synclinal sequences of predominantly mafic volcanic rocks (greenstones) that have been intruded by antiformal tonalite, trondhjemite, granodiorite and monzogranite complexes, is largely unique to the Archaean era (see also, chapters 2 and 4). Early Archaean examples include the 3800-3700 Ma Itsaq gneiss terrane of west Greenland and the 3600-3000 Ma Pilbara and Kaapvaal cratons of Western Australia and South Africa, respectively. Pillow lavas preserved within Isua greenstones (section 2.2) provide clear evidence for the presence of submarine magmatism and of oceans at c. 3700 Ma (Myers, 2001a). Amphibolite-facies rocks from the Isua greenstone belt and from related rocks

26

Chapter 1" The Early Earth

._,g, I/1 t,"1o om

o L Q,.

t~ t'r

3600

3700

3800

3900

4000

4100

20r Pbl 2o6Pb Age

4200

4300

4400

4500

(Ma)

Fig. 1.2-7. Gaussian Summation probability density plot for ~> -t-95% concordant SHRIMP U-Pb analyses older than 4.0 Ga obtained on zircons from the Narryer Gneiss Complex, Murchison granite-greenstone terrane and Southern Cross quartzites of the Yilgarn craton. Data sources: Froude et al. (1983), Compston and Pidgeon (1986), Maas et al. (1992), Nelson et al. (2000), Wilde et al. (2001), Nelson (2002). from Labrador have unradiogenic W isotopic compositions compared to terrestrial values (Schoenberg et al., 2002b), indicating that their 3800 to 3700 Ma sedimentary precursors contained a detrital component derived from meteorite debris. Granite-greenstone crust formation within the Pilbara and Kaapvaal cratons was episodic, with major episodes of c. 10 to 100 My duration resulting in contemporaneous greenstone volcanism on, and granitic intrusion into, pre-existing granite-greenstone crust, separated by periods of magmatic quiescence during which clastic and chemical sedimentary rocks were deposited (Nelson et al., 1999). Although there is evidence suggesting that increased magmatic activity within the Kaapvaal craton may have occurred simultaneously with that in the Pilbara craton, the overall chronological patterns of granite-greenstone crust growth evident for the two cratons are not similar (Nelson et al., 1999), suggesting that by the early Archaean, crust formation was mostly associated with localised processes rather than global-scale convective overturn (cf. Nelson, 1998a; see section 3.4). The greenstone sequences consist of basalt, high-Mg basalt and komatiite with interlayered minor felsic volcanic, chemical and clastic sedimentary rocks. Field and chemical characteristics of greenstone ultramafic and mafic lavas are consistent with their eruption as highly voluminous "flood basalt"-like flows derived by large degrees of partial melting of both shallow and deeper upwelling

1.3. Early P r e c a m b r i a n Stratigraphic R e c o r d

27

mantle sources (see detailed discussions in chapter 4). They are probably therefore the Earth's equivalents of the volcanic plains of the crusts identified on the other terrestrial planets.

1.3.

THE EARLY PRECAMBRIAN STRATIGRAPHIC RECORD OF LARGE EXTRATERRESTRIAL IMPACTS

B.M. SIMONSON, G.R. BYERLY AND D.R. LOWE Introduction

Supracrustal rocks whose features are sufficiently well preserved to enable detailed inferences about Earth's surface environments date back to at least 3470 Ma (e.g., Byerly et al., 2002). The primary features of older supracrustal rocks can be so obscured by deformation and metamorphism that their very origin as sediments is in dispute (e.g., Fedo and Whitehouse, 2002). Based on the lunar record, Ryder (2003) inferred that the impact rate in the Earth-Moon system exceeded c. 15 times the present rate at 3800 Ma but had declined to about 2 times the present rate by 3000 Ma. The oldest impact structure yet found on Earth is the Vredefort structure of South Africa, at 2020 Ma (Grieve, 1998). This means that no impact structures have been identified in terrestrial rocks representing close to 1.5 billion years of geologic history during which a multitude of extraterrestrial objects must have struck the Earth. It is highly unlikely that many impact structures will ever be identified in Archaean and Palaeoproterozoic rocks, given the fragmentary state of their preservation. There is, however, another record of large impacts preserved in the Earth's stratigraphic archive. This record consists of thin layers rich in distinctive sand-size spherules representing silicate droplets formed by the melting and vapourisation of terrestrial target rocks during impacts by large extraterrestrial bodies such as asteroids and comets. At least 11 layers rich in spherules representing a minimum of 7 different impacts have been documented in Archaean to early Palaeoproterozoic successions (Table 1.3-1). All but one of these layers have been found in two relatively small areas with rocks of roughly equivalent ages in South Africa and Australia (Figs. 1.3-1 and 1.3-2). The spherule-bearing layers are generally thin, but they are dispersed over very large areas, so they are unlikely to be erased by regional erosion or tectonism. The spherule-bearing layers will be preserved readily if they were deposited below wave base in continental shelf to slope environments. This means that successions deposited on continental margins can preserve a cosmopolitan record of large impacts, even those which took place in oceanic crust that has since been resorbed into the mantle. Spherule layers may ultimately prove to be our best source of information on the timing and nature of impacts in the early Precambrian. There may also be differences between the impact spherule layers formed in early Earth history versus those formed in more recent times, which could reflect secular changes in Earth's lithosphere and hydrosphere (Simonson and Harnik, 2000). However, we still have much to learn about the known spherule layers, and more examples are needed in order to test the validity of any The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, t).R. Nelson, W.U. Mueller and O. Catuncanu

Table 1.3-1 (part 1). Comparative data on Precambrian spherule-rich layers Name Host formation of of spherule layer layer

Group to which formation belongs

Geographic location of layer

N/A

Graensesci

Vallen

Ketilidian orogen, South Greenland

NIA

Dales Gorge Member Hamersley of Brockman

Hamersley basin, Western Australia

N/A

Wittenoom

Hamersley

Hamersley basin, Western Australia

N/A

Carawine*

Hamersley

Hamersley basin, Western Australia

NIA

Jeerinah*

Fortescue

Hamersley basin, Western Australia

N/A

Monteville

Ghaap

Griqualand West basin, South Africa

S4

Mapepe

Fig Tree

Barberton Greenstone Belt, South Africa

S3

Mapepe

Fig Tree

S3

Ulundi

Fig Tree

Barberton Greenstone Belt, South Africa Barberton Greenstone Belt, South Africa

S2

Mapepe

Fig Tree

S1

Hooggenoeg

Onverwacht Barberton Greenstone Belt, South Africa

S1

Apex Basalt

Warrawoona Pilbara Craton, Western Australia

Barberton Greenstone Belt, South Africa

Approximate Minimum geologic age lateral (in Ga) persistence (in km) z 1.85,

26

Approximate thickness of spherule layer (in cm) 100

Approximate content of spherules in layer by volume 20%

Approximate aggregate thickness of spherules (in cm) 20 - -

N M

Table 1.3-2 (part 2). Comparative data on Precambrian spherule-rich layers Name of layer

Host formation of spherule layer

N/A

Grrensesa

N/ A N/ A

Dales Gorge Wittenoom Carawine* Jeerinah* Monteville Mapepe Ma~e~e Ulundi Mapepe Hooggenoeg Apex Basalt

N/A N/A N/A S4 S3 S3 S2 S1 S1

Diameter of largest spherules observed (in mm) 1.4 2.1 1.1 2.3? 2.0' 1.6 1.6 4.0 2.0 2.5 1 .O 0.8

Grain size of coarsest clastic detritus in layer

Highest Ir concentration reported (ppb)

Most anomalous E ( ~ ~reported c ~ )

References from which data are derived

?J

3

2 P 8' 5

fine sand none? fine sand coarse sand coarse sand coarse sand coarse sand coarse sand coarse sand coarse sand coarse sand coarse sand

0.02 17.9 0.43 1.54 11.4 6.4 450 519 725 3.9 3.0 N.D.

N.D. N.D. N.D. +O. 16 f0.03 +O. 12 f0.03 -0.04 f0.04 -0.32 f0.06 -0.41 f0.08 -0.41 f0.10 -0.09 h 0.03 N.D. N.D.

@.

1

7

2,3,4 2,3,5,6,7 3, 5, 6, 8, 9, 10 4, 10, 11, 12, 13 10, 14, 15 16, 17, 18, 19, 20 16, 19, 20, 21, 22, 23 16, 19, 20, 21, 22, 23 16, 19, 20, 23 19.24 19, 21, 24

4s 6'

Data sources are unpublished data and ( I ) Chadwick el al.. 2001; (2) Simonson, 1992; (3) Hassler and Simonson, 2001; (4) McDonald and Simonson, 2002; (5) Simonson et al., 1998; (6) Simonson et al., 1993b; (7) Woodhead et al., 1998; (8) Simonson and Hassler, 1997; (9) Simonson et al., 2000a; (10) Shukolyukov et al., 2002; (I 1) Simonson et al., 2000b; (12) Simonson et a]., 2001; (13) Simonson et al., 2002; (14) Simonson et al., 1999; (15) Simonson et al., 2000c; (16) Lowe et al., 1989; (17) Kyte et al., 1992; (18) Shukolyukov et al., 2000; (19) Lowe et al., 2003; (20) Kyte et al., 2003; (21) Lowe and Byerly, 1986; (22) Byerly and Lowe, 1994; (23) Byerly et al., 1996; and (24) Byerly et al., 2002. N/A = not applicable. N.D. = no data reported. *= layers in two different formations and groups that nevertheless were probably formed by a single impact (Simonson et al., 2002) ?these layers also contain more irregular, flow-banded particles of former melt up to 2 cm across.

3 c;

3 I;

30

Fig. 1.3-1.

Chapter 1: The Early Earth

1.3. Early Precambrian Stratigraphic Record

31

Fig. 1.3-1. Geographic locations of early Precambrian spherule layers. (a) Representative locations (largely surface exposures) throughout known areas of occurrence of the Carawine and Wittenoom spherule layers (map (2)) and the Dales Gorge and Jeerinah spherule layers (map (3)) in the Hamersley basin (map (1)), Western Australia. (b) The four known locations of the Monteville spherule layer (map (2)) in the Griqualand West basin (map (1)), South Africa; location next to question mark (?) on southwestern edge of basin is a possible fifth occurrence. (c) Geologic map of western part of the Barberton greenstone belt, where spherule layers occur; the stratigraphic units hosting spherule layers are indicated in Figure 1.3-2b and c. Maps in (a) and (b) are adapted from Hassler and Simonson (2001); map in (c) is adapted from Lowe and Byerly (1999a).

32

Chapter 1." The Early Earth

HAMERSLEY BASIN

(a)

Main

outcrop

TRANSVAAL BASIN

area Kuruman

l

Brockman

Iron Fm ~

=,. F,, - 2470+4 F,, F -,, F,, rp

~"Fm v-

Dales Gorge

Mt. McRae . . . . . . + Mt. " . . . . . .

Shale Sylvia

Fm.

Iron Fm. 2521_+3

,~,

S undifferentiated

"22222-

450

m _.%

units b

om/tted

cl. oL.

Wittenoom S

--------

>, m L

"_-----_" (2541+18/-15) :--.-_'---_-'.- 2561_+8

E ,~ ""

Wittenoom

m

Formation

1 Marra

_

_

_ _

_ _

_

!

m M a m b a

Iron Fro.

t

~

!

E~

Iron-Formation

~

_

Fe g m 2597_+5

"--ZZZZ"-'---~;

- Jeerinah S

"---------"-

\ 2 6 2 9 _+ 5 2557+_.49

Revilio _

_

. Fm.

_

_

.

_

_ .

.

_

_

_ .

_

_

oe,"D

.

.

Fm.

Maddina

SL ages

= 6")

_ i i_

Jeerinah

i,o

All radiometric

3

lOOmI

Basalt Layer

_

ell _

Carbonate

Spherule

.

i - l - l i _ _

~

!

Q.

o" Argillite

l - i - l i _ _

i - i_-

m m

_ .

l - i_- l i _ _

are

in Ma.

Basalt

-----------':- 2690 + 16 *VTV~UT UuUvVvU

~.u.u.u,u.v u. continues

down

Monteville S ~ 2650 _+ 8

......

continues

Lokammona

Fm. down

Fig. 1.3-2. Stratigraphic contexts of early Precambrian spherule layers. (a) Stratigraphic columns for main outcrop area of Hamersley basin (left) and Griqualand West basin (right); stratigraphic positions of spherule layers are indicated in space between columns by tick marks adjacent to names of layers (as listed in Table 1); also indicated are available age determinations (in millions of years) and their stratigraphic positions. Columns in (a) adapted from figure 3 of Hassler and Simonson (2001), with new dates cited from Simonson et al. (2000b).

generalisations made about them. In this section, we summarise briefly what is known about early Precambrian spherule layers, discuss the significance of their characteristics for interpreting early Earth history, and speculate about what future studies of these layers might ultimately tell us.

Distinctive Characteristics of Spherule Layers Spherules What sets all of the layers apart are an abundance of spheroidal particles whose distinctive external shapes and internal textures suggest an origin by quenching of silicate melt droplets (Lowe et al., 2003; Simonson, 2003). These particles are referred to collectively as spherules. Most are truly spherical in three dimensions (Fig. 1.3-3), but a substantial minority of spherules depart from a spherical shape. After spheres, ovoids are the commonest shape but lesser numbers of spherules have non-spherical shapes such as teardrops, dumb-

1.3. Early Precambrian Stratigraphic Record

(b)

33

SOUTHERN BGB

4oo

2001~[~ i

~

NORTHERN BGB (lower (Iowe plate)

.....: <---S3, S4

lOO1 .1, ,I

NORTHERN BGB

(upper plate)

i 9

~

!

i 9

:

2

>rv

:-1(3

i

~

" !

:~'.--

uJ ~,~ rr ~ o " ~.

1

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

<--Sl

~:c-

: ~.

i

(.9

=-

,z,-~

~=~-J''~ e~- i P.t-,-~l.k "~ '~ ~'C2 ~_.

~o =- ~

-r

IIIII

N

Fault

km

"' z o o -To

Iny0ka Fault Spherule Layer

= S

Fig. 1.3-2 (continued). (b) Stratigraphic column for southern part of western Barberton greenstone belt with stratigraphic levels of spherule layers indicated on right sides of column as S 1, $2, etc. (c) Stratigraphic columns for northern part of western Barberton greenstone belt with stratigraphic levels of spherule layers indicated on fight sides of columns by arrows labeled $3. Stratigraphic columns in (b) and (c) adapted from Lowe et al. (2003).

bells (Fig. 1.3-3c), and fused spheres (Fig. 1.3-3d). The latter represent pairs of droplets that became joined because they collided in flight while still hot (referred to as agglutinated spherules). Together, spherical and non-spherical spherules form a continuum that is analogous to the fluidal shapes of the variety of tektites and microtektites referred to as splash forms (Glass, 1990). Spherules range in size up to 5 mm in diameter, but those less than a millimetre or two across tend to be the most highly spherical. Related melt particles that range up to at least 2 cm long occur in the Late Archaean Carawine layer (Table 1.3-1), but they have more angular shapes and show different textures internally. These angular particles have more in common with Muong Nong-type tektites than splash forms and may represent the most proximal impact ejecta yet found in Archaean strata (Simonson et al., 2000a). Although they are found in the same layer as spherules, the more angular particles are not considered further here.

34

Chapter 1: The Early Earth

Internally, spherules show a variety of textures indicative of their initial molten state. These include infilled bubble cavities, microlitic textures, radial or spherulitic aggregates of fibrous crystals grown during quenching and/or devitrification (Figs. 1.3-3d, e), and skeletal spinel crystals (Fig. 1.3-3f). With the exception of the spinel crystals, the minerals and glass originally present in the spherules have all been replaced by secondary phases ranging from diagenetic to metamorphic in origin. Some spherules do not show any internal textures diagnostic of their formerly molten state (Figs. 1.3-3a, c). Spherules lacking diagnostic textures could either have consisted originally entirely of glass (like microtektites) or had any such textures obliterated by wholesale replacement. Mineralogically, spherules most commonly consist of quartz, sericite, chlorite, and/or K-feldspar, all of which are secondary in origin. The authigenic origin of the K-feldspar is based on the fact that it replaces entire spherules, has a low sodium content, and commonly occurs in crystals whose shapes are typical of plagioclase microlites (Simonson, 2003). Primary textures are best preserved in spherules that consist of K-feldspar, whereas textures are less likely to be preserved as well in spherules that currently consist of quartz, sericite, and chlorite. Where textures are preserved in the latter, they tend to be outlined by fine inclusions of rutile. In general, Early Archaean spherules have been metamorphosed and deformed to a higher degree than those of the Late Archaean and Palaeoproterozoic. However, textures in the latter spherules can also be obliterated via diagenetic replacement with coarse carbonate crystals (calcite, ankerite, and/or dolomite) and/or pyrite. The textures in spherules are also obscured commonly by surface alteration in outcrops of layers. This takes the form of impregnation by caliche carbonate and oxidation of ferrous iron in minerals such as pyrite and ankerite. Nevertheless, spherules in carbonate-rich layers typically stand out in high relief on weathered surfaces (Fig. 1.3-4a).

Fig. 1.3-3. Photomicrographs of impact spherules in plane polarised light. Long axes of fields of view are 1.85 mm in (a) and (b), 460 ~tm in (c) and (d), 2.5 mm in (e), and 600 ~m in (f). (a) Well-sorted spherules from Dales Gorge layer. The spherules consist of K-feldspar (grey, primarily rims) and weathered (oxidised) stilpnomelane (dark). A cement consisting of stilpnomelane filled the interstitial pores early enough to prevent significant compaction of spherules. Carbonate rhombs (white) cut across both cement and spherules. (b) Well-sorted spherules from Wittenoom layer. These spherules consist of K-feldspar (grey and white) plus some calcite (also white); spherules show more textural heterogeneity and are more compacted than those in a. (c) Dumbbell-shaped spherule from Dales Gorge layer, with a thin rim of K-feldspar and massive core of felted stilpnomelane. (d) Particle from Wittenoom layer, consisting of two K-feldspar spherules that fused in flight; the one with circular cross-section was stronger, and therefore cooler and/or more crystalline at the time of collision; whereas the other spherule contains a teardrop-shaped bubble cavity. (e) Well-sorted spherules with bubble cavities, from an unweathered core sample of Monteville layer. The spherules and most of interstitial material consists of K-feldspar, variably darkened with impurities. Finely felted sericite is also present, especially filling bubble cavities. The small clear crystals between spherules are fine quartzose sand. (f) Chlorite-chert spherule from impact layer $3 in the Barberton greenstone belt, with an abundance of Ni-rich spinel octahedra (black), many hollow.

1.3. Early Precambrian Stratigraphic Record

35

Spherule-rich layers

Spherules occur in discrete layers that form regionally persistent markers beds. For example, the S 1 layer in the Hooggenoeg Formation of the Onverwacht Group persists along strike for tens of kilometres around folds in the Barberton greenstone belt (Lowe et al., 2003) and the Wittenoom layer is preserved over tens of thousands of square kilometres in the relatively undeformed strata of the Hamersley basin (Simonson, 1992). The areas

Fig. 1.3-3.

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Chapter 1: The Early Earth

Fig. 1.3-4. Surface exposures and hand sample of impact spherule layers. (a) Caliche-coated surface normal to bedding on Monteville spherule layer type locality, with numerous white spherules standing out in relief; coin is 16 mm in diameter. (b) Weathered surface of Wittenoom layer at the type locality of the Bee Gorge Member. Light resistant beds at the top and bottom are ambient carbonate lutite layers; intervening darker strata make up the spherule layer. Note the imbricated stack of rip-up slabs of limestone in centre and current-ripple crests climbing obliquely to right above them. The coin is 21 mm in diameter. (c) Sawn slab from impact layer $2 from the Barberton greenstone belt, with a mixture of numerous large siliceous rip-up clasts plus spherules and other finer debris. Major divisions on the scale at bottom are in cm. (d) Wittenoom layer from near Mt. Farquhar in the western part of the Hamersley basin. The spherules occur in cm-thick, discontinuous lenses at level of head of rock hammer (centre). Resistant, dm-scale bed just above the hammer head is dolomite, probably a distal carbonate turbidite. Enclosing both are evenly layered, shaly strata typical of the basinal lutites of the Bee Gorge Member of the Wittenoom Formation.

over which the spherule layers were originally deposited were presumably much larger. The global extent of the Cretaceous-Tertiary (K/T) boundary layer proves that spherules generated by one large impact can carpet the surface of the entire planet (Smit, 1999). It has already been suggested that individual impacts in both the Early and Late Archaean may have each generated contemporaneous spherule layers in South Africa and Australia (Simonson et al., 1999; Byerly et al., 2002). Confirmation that one or both of these layers were dispersed globally must await detailed scrutiny of contemporaneous strata at localities scattered widely around the Earth.

1.3. Early Precambrian Stratigraphic Record

37

Most of the spherule layers recognised so far are anomalously coarse sand-rich beds within mud-dominated successions. Based on sedimentary structures displayed in both the mud- and sand-rich layers, most of these successions were deposited below wave base in deeper water environments. The spherule layers per se commonly display large rip-up clasts from subjacent strata (Figs. 1.3-4b, c), normal grading, and/or cross-stratification (Fig. 1.3-4b). The Late Archaean to Palaeoproterozoic spherule layers are enclosed in muddy strata which are thinly laminated and show few or no signs of deposition from currents or waves (Fig. 1.3-4d; Hassler and Simonson, 2001). The Early to Middle Archaean spherule layers were deposited in a broader spectrum of environments involving a mix of shallower to deeper water and even subaerially emergent environments such as fan-deltas (Lowe et al., 2003). In all settings, the spherule layers are products of anomalously high-energy events. At least in some cases, the source of the high energy appears to have been tsunami waves and attendant currents induced by the impacts that created the spherules themselves (Hassler et al., 2001; Lowe et al., 2003). In other cases, they formed thick layers with coarse clasts deposited by mass movements (Simonson, 1992; Hassler and Simonson, 2001). Regardless of how and where they were deposited, spherule layers are highly distinctive. Under optimal conditions, they can be recognised confidently on the basis of a single exposure one metre wide (e.g., the early Archaean $4 layer; Lowe et al., 2003) or by examination of a single sample from a drill core with a hand lens (e.g., the Late Archaean Jeerinah layer; Simonson et al., 2000b). Evidence f o r an impact origin The study of impact spherule layers is a relatively new development in the field of Precambrian geology, essentially dating from the publication of Lowe and Byerly (1986). Most Precambrian impact structures are Meso- to Neoproterozoic in age, and even these are not very numerous (Shoemaker and Shoemaker, 1996). Given their lower abundance and higher degree of deformation compared to younger rocks, it is unlikely that a record of impact structures as rich as that of the Phanerozoic will ever be found in Archaean to Palaeoproterozoic rocks. However, there is no reason why additional spherule layers could not be preserved in many Precambrian successions. The interpretation of these spherules as products of impacts by large extraterrestrial bodies initially met with widespread scepticism and has been challenged on a number of counts. Most troubling among the arguments raised against the impact interpretation is the failure to isolate any shock-metamorphosed minerals from the spherule layers, a deficiency which is still not fully understood. However, a number of arguments have been marshaled in support of the interpretation of these spherules as impact products. These arguments, recently summarised in Lowe et al. (2003) and Simonson (2003), are recapitulated below. Perhaps the most obvious argument has to do with the textural characteristics of the spherules themselves. It is now well documented that a number of Phanerozoic impacts produced ejecta layers that consist predominantly of highly spherical grains of former melt, generally known as microtektites (if they consist entirely of glass) or microkrystites (if they display primary crystallisation textures). In addition, a minority of the particles in Phanerozoic microtektite layers display distinctive shapes such as teardrops and highly

38

Chapter 1: The Early Earth

elongated dumbbells. These are a close match for the shapes and sizes of particles found in many of the Precambrian spherule layers, whereas there is no close match amongst layers formed by volcanic eruptions. The only volcaniclastic layers in which spheroidal particles dominate are those rich in accretionary and armoured lapilli (Reimer, 1983), which are very different from impact spherules. Accretionary and armoured lapilli are typically larger than impact spherules, and if they are the same size, accretionary lapilli typically show crude concentric growth bands internally (e.g., Lowe, 1999b) rather than the bubble cavities and crystallisation textures of impact spherules. Even where internal textures are altered beyond recognition, the presence of unusual shapes such as dumbbells and teardrops can be used to distinguish impact spherules from accretionary lapilli. Accretionary lapilli formed during impacts have textural characteristics identical to those of volcanic accretionary lapilli (e.g., Warme et al., 2002), but no accretionary lapilli have been reported from any spherule-rich layer in either the Precambrian or Phanerozoic. Spherules of silicate melt are of course widespread in volcanic deposits (e.g., Heiken and Lofgren, 1971; Hay et al., 1979). The key difference between the spherule layers described here and the melt spherules formed by terrestrial volcanic processes such as lava fountaining is that volcanic spherules are always highly subordinate to non-spherical particles such as bits of scoria or Pele's hair. In contrast, the majority of the initially molten particles in impact spherule layers are spherical. This can be determined readily in Precambrian spherule layers, provided they are not too severely affected by tectonic deformation or diagenetic processes such as replacement and compaction. In rare instances, a high proportion of the melt particles in a volcaniclastic layer have spheroidal shapes (e.g., G61inas et al., 1977a), but only in layers close to their point of origin and intimately associated with volcanic flow rocks. This is clearly different from impact spherules, which occur in thin layers that persist regionally to perhaps globally. More direct evidence that extraterrestrial bodies were involved in forming the spherules comes from the geochemistry of spherule-rich samples. A chondritic component was first recognised in spherule layers on the basis of iridium anomalies, enrichments in siderophile elements, and platinum group elements (PGEs) with inter-elemental ratios approaching chondritic values (Lowe et al., 1989; Kyte et al., 1992; Simonson et al., 1998, 2000c). Objections were raised to some of these interpretations, in part because of the extremely high values of PGEs in some samples (Reimold, 2000). However, given the great age of these layers, some blurring of the extraterrestrial signal is to be expected. In the case of some of the late Archaean layers, the departures from chondritic ratios closely match those in a layer of known impact ejecta from the late Precambrian Acraman impact structure of South Australia (Simonson et al., 1998). Another line of evidence pointing to an extraterrestrial origin for the spherule layers are nickel-rich chrome spinels in some of the older Archaean layers (Fig. 1.3-3f; Byerly and Lowe, 1994; Lowe et al., 2003). Although they differ in composition from those associated with K/T and Eocene impact spherules, the spinels in the Archaean spherules are clearly enriched in extraterrestrial components and differ in composition from spinels found in closely associated komatiites. Most recently, extraterrestrial chromium has been detected in Precambrian spherule layers (Table 1.3-1) via the isotopic technique of Shukolyukov and Lugmair (1998).

1.3. EarlyPrecambrian Stratigraphic Record

39

Timing of Impacts

The ages of the older Archaean spherule layers are well constrained by U-Pb zircon dates from both the layers and associated rocks. Identical zircon populations were reported from the oldest spherule layers in both the Barberton greenstone belt of South Africa and the Pilbara craton of Western Australia, indicating their depositional ages are indistinguishable at 3470-t-2 Ma (Byerly et al., 2002). Similar analyses have been done on two of the younger layers in the Barberton greenstone belt and place their depositional ages at c. 3260 Ma (Byerly et al., 1996) for $2 and 3243 -t- 4 Ma for $3. Layer $4 is believed to be very close in age to $3 because it is only 6.5 m stratigraphically above $4 (Lowe et al., 2002). Taking the total span of time represented by these four layers and simply dividing it into equal increments yields a recurrence interval (RI) of roughly 77 My between spheruleforming impacts in the Barberton greenstone belt. Focusing solely on the three youngest layers yields a significantly shorter RI of roughly 10 My between spherule-forming impacts (Lowe et al., 2003). The ages of the Late Archaean to Early Palaeoproterozoic spherule layers are not as well constrained. A minimum of three impacts is required to create the spherule layers in the Hamersley basin (Table 1.3-1). Direct dating has only been attempted on the Wittenoom layer. Pb-Pb analysis of carbonates in and around this layer yielded a date of 2541 + 11~Ma (Woodhead et al., 1998), which is consistent with other dates from surrounding strata (Trendall et al., 1998). The age of the spherule layer in the Jeerinah Formation is probably the best known as it lies within a few meters stratigraphically of a 2629 + 5 Ma tuff dated via the SHRIMP U-Pb zircon method (reference number 139 in Table 1 of Nelson et al., 1999). The age of the Dales Gorge spherule layer, the youngest of the Hamersley layers, can only be approximated as roughly 2490 Ma by interpolating between widely spaced zircon-dated tuff layers (summarised in Nelson et al., 1999). New SHRIMP U-Pb analyses of zircons from tufts in overlying formations (Pickard, 2002) suggest the Dales Gorge layer may be slightly older than this. The age of the spherule layer in the Monteville Formation (Table 1.3-1) is not well constrained either. It is bracketed between tufts whose zircons yielded SHRIMP U-Pb dates of 2650 -4- 8 Ma and 2602 -t- 14 Ma (Gutzmer and Beukes, 1997) and is stratigraphically closer to the older one. Taking outer age limits of 2630 and 2490 Ma for the three spherule-forming impacts in the Hamersley basin yields a RI of c. 70 My for the younger cluster of early Precambrian layers. This value is surprisingly close to the initial estimate from the Barberton greenstone belt. The existence of the Monteville spherule layer is not a factor because it may have been formed by the same impact that produced the spherule layers in the Jeerinah Formation (Simonson et al., 1999, 2002). In addition to the spherule layers in Western Australia and South Africa, one possible spherule layer has been reported from the Ketilidian orogen of South Greenland (Chadwick et al., 2001; Table 1.3-1). This layer has not been studied in the field since the 1960s, and the succession in which it occurs is not well dated. Chadwick et al. (2001) summarised the age constraints and indicated that the maximum and minimum depositional ages for this spherule layer were c. 2130 and 1848 Ma respectively. That places it firmly in the Palaeo-

40

Chapter 1: The Early Earth

proterozoic and significantly younger than all of the other known Precambrian spherule layers. Implications for Earth History

The characteristics of impact ejecta are determined by the properties of two things: the impactor and the target area. Impactor properties that can affect the nature of the ejecta include its composition, velocity, and angle of impact. Properties of the geologic deposits in the target area that can affect the nature of the ejecta include their chemical composition, their state of induration, the existence of any anisotropy, and whether or not they are covered by water. Simonson and Harnik (2000) noted that the handful of early Precambrian impact spherule layers found so far appear to differ in some ways from their Phanerozoic counterparts. If indeed they do, this could reflect secular changes in either the nature of impactors or Earth's surface environments and rocks, or both, through time. The characteristics of early Precambrian spherule layers that may differ from Phanerozoic examples are outlined next, followed by a discussion of the possible implications for early Earth history. The most striking difference between early Precambrian and Phanerozoic spherule layers is the fact that spherules are simply much more abundant in the older layers. Spherules constitute the vast majority of impact-generated material in all of the early Precambrian layers, so the aggregate thickness of spherules calculated independently of other types of detritus is a reasonable proxy for impact size. Precambrian spherule layers vary greatly in thickness (Table 1.3-1), but most of this variation reflects dilution by non-impact detritus entrained and redeposited by waves and/or currents. Even though the average aggregate thickness of spherules is often very difficult to estimate, it appears to be measured in tens to hundreds of millimetres in many Precambrian spherule layers (Table 1.3-1). In contrast, the aggregate thickness of spherules in the distal parts of the KfI' boundary layer averages 2 to 3 mm and never exceeds 100 mm (Smit, 1999). The aggregate thickness of impact ejecta can be much greater at more proximal sites, but such locations are not good analogues for Precambrian spherule layers because the ejecta is not dominated by spherules. For example, spherules are found in an Eocene ejecta layer 80 mm thick on the continental slope off New Jersey at DSDP site 612, but they are outnumbered by shock-metamorphosed solid ejecta and non-spheroidal impact glass particles (Glass, 1989). The great thickness of spherules strongly suggests that early Precambrian spherule layers were created by objects at least as large as the K/T boundary projectile, whose diameter is estimated at 10 km (e.g., Shukolyukov et al., 2000). However, equating layer thickness with projectile size has its pitfalls. The total mass of spherules in the K/T boundary layer can be estimated with confidence because its thickness has been determined all over the world (Smit, 1999). In contrast, early Precambrian layers are only preserved over a small fraction of the Earth's surface, and even within these areas they vary dramatically in thickness. For example, the Jeerinah and Carawine layers were probably formed by the same impact (Simonson et al., 2002), yet the entire Jeerinah layer consists of c. 3 mm of nearly pure spherules at one site whereas the Carawine layer c. 200 km away has an aggregate spherule content of roughly 250 mm (Table 1.3-1). There is as yet no data whatsoever con-

1.3. Early Precambrian Stratigraphic Record

41

cerning Precambrian spherule layers from most of the Earth's surface. Therefore, extreme caution is advisable in estimating the total mass of spherules from a given impact in the face of such extreme variations in thickness. The lateral extent of the spherule layers is also something of an open question. The furthest any early Precambrian spherule layer has been traced continuously is a few hundred kilometres (Table 1.3-1). The 3470 Ma layers in South Africa and Western Australia are most likely products of a single impact (Byerly et al., 2002), and the Late Archaean Monteville and Jeerinah spherule layers in South Africa and Western Australia, respectively (Table 1.3-1), could also have formed from a single impact (Simonson et al., 1999). These correlations, plus the great thickness of spherules noted above, clearly indicate the spherules were dispersed laterally for thousands of kilometres. However, striking geological similarities between these two cratons suggest they were in close proximity to one another at times (see, however, review in Nelson et al., 1999, for a contrasting viewpoint), in which case the observed thicknesses might be typical for part, but not all, of the Earth's surface. On the other hand, geochemical evidence suggests the Precambrian spherule layers represent ejecta dispersed globally. By analogy with compositional contrasts between microtektites and microkrystites in the K/T boundary layer (J. Smit et al., 1992), the high Ir and Cr contents of the Precambrian layers are more typical of distal than proximal spherules, although it is difficult to generalise given the small number of Precambrian spherule layers so far identified. It is still not understood why some Phanerozoic impacts produced spherules in abundance and dispersed them over wider areas whereas others apparently did not. Most early Precambrian spherule layers were probably produced by objects at least as massive as the K/T boundary impactor. Such major early Precambrian impacts were probably more closely spaced in time than was the case for their Phanerozoic counterparts. The recurrence interval (RI) of spherule-forming impacts recorded in both the Barberton greenstone belt and the Hamersley basin is roughly 70 My, which is virtually identical to the Phanerozoic RI of an impactor ~> 10 km in diameter (Chapman and Morrison, 1994). One would expect the RI to be even shorter in the early Precambrian from Ryder's (2003) reading of the lunar impact record, and the early Archaean spherule layers coincide with evidence of a higher rate of impacts from lunar impact spherules (Culler et al., 2000). In fact, the RI of large impacts in the late Archaean to Palaeoproterozoic is likely to become shorter via the discovery of additional spherule layers. Almost all of the Precambrian spherule layers found so far are in two relatively small areas (Fig. 1.3-1) that represent a fraction of early Precambrian time. Other spherule layers are likely to be found as the search spreads to strata in new geographic areas and of different ages. If the Precambrian spherule layers found so far represent thicker accumulations in relatively close proximity to the points of impacts rather than global blankets, even the layers found to date are not representative of all of the large impacts during their respective time windows. This too would mean the RI of large impacts in the early Precambrian was significantly shorter than the inferred maximum of 70 My. Differences in the Earth's hydrosphere could also contribute to contrasts in thickness and frequency of early Precambrian versus Phanerozoic spherule layers. The modern ocean

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Chapter 1: The Early Earth

shields the deep sea floor from all but the largest impactors because the diameter of a projectile has to roughly equal the water depth before rocks on the seafloor can be excavated and dispersed as vapour or melt droplets (Gersonde et al., 1997). What if Archaean oceans were shallower (see discussion, section 3.6) on average than those of the Phanerozoic? Then smaller projectiles could have produced spherules from oceanic crust in the Archaean. This in turn would mean that a greater number of spherule layers would be produced, even if the flux of impactors were the same then as it was in the Phanerozoic. A higher abundance of impact spherules from mafic target rocks in the early Precambrian could therefore signal shallower oceans as opposed to a greater rate of impacts when compared to the Phanerozoic. The chemistry of Archaean seawater (sections 5.2-5.5) could have also played an important role. Silica concentrations were much higher in Precambrian oceans than they are today (Siever, 1992), so Precambrian seafloor environments may have been more amenable to the preservation of impact spherules. This is particularly true of those with low silica contents, as they are the most vulnerable to dissolution in modern seawater (Glass, 1984). In fact, the two basins where most of the spherule layers have been found are known for their silica-rich deposits. The Hamersley basin is well known for the size and extent of its banded iron formations (Trendall, 1983b) (section 5.4), and early silicification is very pervasive in the Barberton greenstone belt (Duchac and Hanor, 1987; Lowe, 1999a). Perhaps unusually high levels of silica played a role in the preferential preservation of spherules layers in these two basins, as did the fact that they contain some of the least altered and deformed supracrustal rocks on Earth in their respective time windows. The internal textures of early Precambrian spherules also appear to differ from those of Phanerozoic impact spherules. Most of the mass in distal impact ejecta is derived from target materials (Koeberl, 1998), so any systematic change in the primary composition of impact spherules over geological time probably reflects changes in the average composition of Earth's lithosphere. Simonson and Harnik (2000) proposed that early Precambrian spherules were more basaltic on average than Phanerozoic spherules. This is consistent with a smaller volume of continental crust in the early Precambrian (e.g., Eriksson, 1995) (chapter 2, section 3.6), which would statistically produce a lower percentage of ejecta from high-silica continental rocks and a correspondingly higher incidence of impacts into oceanic crust. A predominance of impacts into mafic target materials could also help account for the paucity of shock-metamorphosed grains from continental source rocks such as quartz and zircon in early Precambrian layers (Simonson et al., 1998). This is not to say that continental impacts did not produce ejecta in the Archaean. The Late Archaean Carawine layer contains melt particles that resemble Muong Nong-type tektites (Simonson et al., 2000a, 2002). The latter are larger and more layered than normal tektites and formed during a continental impact in Indochina (Glass, 1990). The melt material in the Carawine layer therefore represents a good candidate for early Precambrian continental impact ejecta. Whatever their size, frequency, and target materials, there is clear evidence that the impactors responsible for the Precambrian spherule layers had chondritic compositions. The relative abundances of siderophile elements, particularly PGEs, support this interpretation (Lowe et al., 1989; Kyte et al., 1992; Simonson et al., 1998, 2000c). Even among

1.3. Early Precambrian Stratigraphic Record

43

the few spherule layers known, variations in the relative abundances of chromium isotopes (Table 1.3-1) indicate several different types of impactors were responsible. Meteoritic materials consistently have excess 53Cr relative to terrestrial materials; additionally, carbonaceous chondrites have excess 54Cr relative to other terrestrial and extraterrestrial materials (Kyte et al., 2003). In the convention proposed by Shukolyukov et al. (2000), both of these indicators are combined into a normalised E53 value in which terrestrial materials have a value of zero, whereas ordinary and carbonaceous chondrites depart from zero in positive and negative directions respectively. Most of the older Archaean layers show negative e53 anomalies, indicating they were produced by impactors with compositions like carbonaceous chondrites (Lowe et al., 2003; Kyte et al., 2003). In contrast, some of the late Archaean spherule layers show positive anomalies, indicating they were produced by impactors with compositions like ordinary or perhaps enstatite chondrites (Shukolyukov et al., 2002; McDonald and Simonson, 2002). Data are still scarce and more work is clearly needed to constrain the compositions of the impactors which produced the Precambrian spherule layers.

Scientific Value of Spherule Layers Research on spherule layers by a handful of scientists has already yielded a wealth of tantalising information about impacts in early Earth history, but it has generated even more questions about how and why they form. Only now are earth scientists at large starting to search systematically for these layers, and already they are being discovered in strata that have been examined by geologists many times in the past (e.g., Walkden et al., 2002). Given the large number of impacts that must have happened in early Earth history (Glikson, 1999), it is likely that many additional Precambrian spherule layers will be discovered. In closing, we will speculate a little about what we stand to learn about impacts and their role in early Earth history via the discovery of additional spherule layers and their characterisation. Precambrian spherule layers can provide important insights into the dynamics and environmental effects of large impacts. Craters and related structures provide the best evidence of the location and size of an impact, but they are commonly poor recorders of an impact's environmental effects. In contrast, spherule layers are ideal for assessing environmental effects because each one is found enclosed within sediments deposited immediately before and after the impact. Moreover, Precambrian layers are superior to most Phanerozoic layers for studies of the processes which happen during spherule deposition because physical sedimentary structures are never obscured by burrowing organisms, although they can be deformed tectonically or otherwise altered just as easily as Phanerozoic strata. Most Precambrian spherule layers found so far show evidence of reworking, very possibly by processes induced by the impacts themselves (Hassler and Simonson, 2001), although a minority appear to be direct fallout deposits (Lowe et al., 2003). This persistent evidence of reworking may indicate that the thick spherule accumulations found to date were mostly deposited in the same region as the impact that generated them, rather than on a distant part of the Earth's surface. Lastly, studies of the textures and compositions of the spherules may

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Chapter 1: The Early Earth

shed light on processes in large ejecta clouds. It is unclear at present exactly what type(s) of impact spherules form as ballistic droplets of impact melt versus condensates from rock vapour. Much of what we infer about this topic is based on studies of the K/T boundary layer, which has been examined in great detail but, in the final analysis, is still the product of a single impact. Studies of spherules from numerous impacts are needed to delineate the full range of possibilities, so contrasts between spherules formed by different Precambrian impacts may shed new light on this question. They may also help answer the question of why only certain impacts produce spherule layers. Even if we don't fully understand how they form, impact spherule layers may provide excellent tools for stratigraphic correlation. All of the spherule layers found so far are "marker beds" that were deposited rapidly by anomalously high-energy events, probably in time periods ranging from a few hours to perhaps a few days. Combined with their dispersal over very large areas, this means spherule layers are excellent time-stratigraphic markers. Byerly et al. (2002) have already made a convincing case for the correlation of Early Archaean spherule layers from South Africa to Western Australia, and further work may establish similar correlations among Late Archaean spherule layers. If these and other layers can be identified successfully and correlated on other continental shields as well, they could form the basis for a framework of intercontinental stratigraphic correlation among early Precambrian successions whose time resolution could be on the order of centuries, decades, or even years in strata that are billions of years old. Studies of spherule layers may allow us to detect changes in both impactors and target rocks through geologic time. Spherule layers have already extended the record of terrestrial impacts back in time by almost 1.5 billion years. Provided enough can be found to assemble a database comparable to that of impact structures (Grieve, 1998), spherule layers could be used to test for variations in the rate of large terrestrial impacts through geologic time, for example, via comet showers (Shoemaker, 1998). The fact that spherule layers always occur in stratigraphic context means they could also be used to test for correlations between large impacts and turning points in the history of life (see chapter 6). Even in the absence of skeletal fossils, recent advances in the study of biomarkers might make it feasible to do this in some early Precambrian successions. While the number of spherule layers known to date is clearly too small for a statistical approach to secular variations, some tantalising coincidences have already been documented. For example, spherule layers in the Barberton greenstone belt coincide with critical transitions from volcanic to sediment-dominated successions marked by major unconformities, suggesting that the impacts that generated them induced major lithospheric changes (Lowe et al., 2003). The oldest spherule layer in the Hamersley basin occurs just a few metres stratigraphically beneath what is perhaps the oldest major BIF on Earth. Large iron-formations require large-scale hydrothermal systems in the deep ocean (Jacobsen and Pimentel-Klose, 1988) (various genetic models are discussed by Trendall and Blockley, section 5.4), so the onset of BIF deposition could reflect hydrothermal activity initiated or enhanced by a large oceanic impact. Late Archaean spherule layers in the Hamersley and Griqualand West basins (quite possibly products of the same impact) are also stratigraphically close to the base of the first large carbonate plat-

1.4. Early Precambrian Impact Events

45

forms (Grotzinger, 1989). Perhaps the appearance of the first carbonate platforms reflects an adaptive radiation of new microbial life forms in the wake of a large impact? In summary, the study of impact spherule layers, particularly those of the Precambrian, currently stands in a position analogous to studies of impact craters near the start of the 20th century. There have been relatively few discoveries and detailed studies made so far, but the pace should accelerate dramatically in the decades to come. As it does, we stand to learn a great deal about both projectiles from space and the surface of the Earth they hit.

1.4.

STRATEGIES FOR FINDING THE RECORD OF EARLY PRECAMBRIAN IMPACT EVENTS

D.H. ABBOTT AND J.T. HAGSTRUM Introduction

Recent work has shown that there is a high degree of correlation between the timing of impact events and of strong plumes (e.g., komatiites; Abbott and Isley, 2002a) (see also section 4.3 on komatiites, and sections 3.2 and 3.3 on plumes). There is also a c. 26 My periodicity in komatiite activity at about 2.7 Ga (Isley and Abbott, 2002). This is the same periodicity (26-36 My) as is observed in the Phanerozoic record of large impacts (Stothers and Rampino, 1990; Rampino and Stothers, 1999). These correlations suggest that mantle plumes were strengthened by large impact events. If this is true, the ideal part of geological time to test this idea is within the Archaean rock record. During the Archaean, large impacts were more abundant than in the Phanerozoic (section 1.3), and the strongest mantle plumes were at least 10 times stronger than the strongest Phanerozoic plumes (Abbott and Isley, 2002b) (sections 3.2 and 3.3). The Archaean rock record also contains most of the documented komatiites (section 4.3). So far, only about 7 Archaean impact debris layers have been identified (Lowe et al., 1989; Simonson et al., 1999) (see section 1.3). As we will show, there should have been at least 350 impact events that produced global impact layers during the period from 2.5 to 3.8 Ga. Because continents covered about 27% of the Earth's surface by the late Archaean, most of these impacts would have been within the Archaean ocean basins. In this section, our knowledge of Phanerozoic oceanic and terrestrial impact events has been applied to formulate a search strategy for Precambrian distal impact layers. The search strategy incorporates: (1) characteristic sedimentary facies of distal impact layers, (2) geophysical methods of finding distal impact layers, (3) distinctive spherule and rock types in impact layers, and (4) distinctive minerals in impact layers. Some possible stratigraphic associations that have not yet been found, but that might be characteristic of some distal impact layers, are also proposed. The Precambrian Farth: Tempos and Events Edited by EG. Eriksson. W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Calculating Crater Size Necessary to Produce a Global Impact Layer The rate of formation of impact craters on the Earth's surface can be estimated by tabulating the abundance of impact craters on cratons, and by measuring the abundance and size distributions of asteroids and comets with Earth-crossing orbits. The estimated cratering rate for the last 120 My is about 516 impact craters over 10 km in diameter (Shoemaker, 1983; Shoemaker et al., 1990; Shaw, 1994). Because the recent cratering rate is higher than in the rest of the Phanerozoic, we assume that the recent cratering rate sets a lower bound on the background-cratering rate for Archaean time. Using this cratering rate, at least 5,500 craters over 10 km in diameter must have formed between 3.8 and 2.5 Ga. Because an 11 km diameter crater is required to put material into the stratosphere and to make microtektites (Glass et al., 1979), most Archaean cratering events should have produced regional tektite-bearing impact layers that covered minimum surface areas of c. 7.1 • 105 km 2. Because most of the Archaean Earth's surface was ocean basin, most of the resulting impact layers would have been deposited within the deep ocean basins and subsequently destroyed by subduction (section 3.6 debates the age of the onset of plate tectonics). To ensure that a layer was deposited on the continents and preserved, we must estimate the minimum crater size producing a global impact sedimentary layer. A global impact layer is distinguished by the presence of impact spherules, nearperfectly round spheres of quenched impact melt (Harnik and Simonson, 2000; Simonson et al., 1999). Some impact spherules are crystalline and are called microkrystites (see section 1.3). Other impact spherules are pure glass, and are microtektites (McCall, 2001). Tektites are generally deposited asymmetrically and close to the source crater, whereas microkrystites are generally more widely distributed. Both the Popigai impact event (crater diameter 100 km) and the Chesapeake Bay impact event (crater diameter 80 km) produced global microkrystite-bearing layers of late Eocene age (Whitehead et al., 2000). The ages of the two impact layers are nearly identical and the layers are intermixed in deep-sea cores. The combined impact layer is typically several centimetres thick. In the Phanerozoic, ash layers this thick are usually bioturbated into the surrounding sediments. A minimum layer thickness of c. 5-10 cm is required to prevent dispersal of a rapidly deposited layer by bioturbation. Thus, the presence of two closely spaced impacts increases the possibility that a global layer will be preserved and it is probable that an even larger minimum size for an impact crater might be required to produce a global layer in the Phanerozoic. However, in the Archaean, bioturbating organisms were not present (e.g., section 6.2). Global impact layers only a few millimetres thick should be preserved as intact layers and the layers should have sharp upper and lower contacts. Therefore, we assume that during the Archaean a minimum crater size of c. 80 km would have produced a global impact layer. The cratering rate decreases by about a factor of two for each doubling in crater size (Hughes, 1998). Thus, the cratering rate for 80-kin diameter impact craters is 1/24 or about 1/16th the rate of creation of 10-km impact craters. This predicts that 32 impact craters over 80 km in diameter formed during the last 120 My. For the entire Archaean, there should have been about 350 impact events that produced global layers a millimetre

1.4. Early Precambrian Impact Events

47

or more in thickness. So far, we know of only about 7 Archaean impact layers (Lowe et al., 1989; Simonson et al., 1999) and no Archaean impact craters have been identified. Even with very poor stratigraphic preservation, there must be many more impact layers within Archaean sedimentary rocks.

Important Differences: Submarine, Abyssal and Continental Impacts Impact events that are large enough to produce global layers will have characteristic features that depend on the environment of the impact. If the impact is submarine, it will produce characteristic sedimentary structures over a broad region. If the impact occurs on continental crust, either the continental shelf or subaerial continental crust, it is likely to hit either quartz-rich sedimentary rocks or quartz-bearing igneous and metamorphic rocks. These rocks are all potential sources of shocked quartz. In contrast, an impact into the abyssal ocean will strike a thin cover of pelagic sediment overlying oceanic crust (depending on water depth; section 1.3). The rock types that make up the oceanic crust, MORB (mid-ocean ridge basalt), OIB (ocean island basalt) and oceanic plateau basalt are not quartz normative. These rocks do not contain quartz grains that form shocked quartz. Only oceanic fracture zones or sheeted dykes may contain small amounts of a quartz normative rock type, such as plagiogranite. Typical pelagic sediment has only a small component of wind-blown quartz from the continents. This wind-blown quartz is very fine grained (i.e., < 63 ~m in diameter). The shocked quartz that is unique to impacts has at least two directions of planar deformation features (PDFs) within a single grain (Short and Bunch, 1966). Because multiple directions of PDFs are easier to find within large grains of quartz, these small grains make it difficult to find shocked quartz within abyssal impact layers. Thus, shocked quartz grains are not expected to be characteristic of abyssal impact layers and cannot be relied upon to verify an impact event in the abyssal ocean. Prior to 3.0 Ga, the continental crust had about 20% of its present day volume (Taylor and McLennan, 1985). Typical Archaean continental crust was thicker (c. 49 km) than Phanerozoic crust (c. 40 km; Mooney et al., 1998; Abbott et al., 2000), so the continents covered about 7% of the Earth. By the end of the Archaean, continental crust had about 80% of its present day volume (Taylor and McLennan, 1985) (see also, section 2.8). Including their greater crustal thickness, Late Archaean continents covered about 27% of the Earth, compared to 41% at present (Cogley, 1984). Particularly in the Early Archaean, it is likely that there was significantly less wind-blown quartz in abyssal sediments (section 7.6 discusses temporal control on aeolian deposits). Therefore, it is most probable that Archaean-age impact layers do not contain shocked quartz.

Differentiating Cometary and Asteroid Impacts Anomalies in the amount of Ir and other PGEs are widely used as evidence for impact events (Schmidt et al., 1993; Smit et al., 1997; Chadwick et al., 2001) (see also section 1.3).

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However, the concentration of iridium in an impact layer depends on the nature of the impactor. Most comets contain about 33-50% volatile material that contains almost no Ir. The non-volatile remainder is not as high in Ir as a typical asteroid. Furthermore, typical Earth-encounter velocities for asteroids are 15-25 km/s whereas the velocities of comets are typically between 40 and 70 km/s (French, 1998). Therefore, the dilution of the impactor in the ejecta blanket differs greatly between asteroids and comets. Because comets produce larger craters for a given size of impactor, cometary material constitutes a smaller proportion of the ejecta layer. Because the size of an Ir anomaly is related to the dilution of the impactor within the ejecta layer, this further reduces the average Ir anomaly from cometary impact events. Van den Bergh (1994) estimated that the typical Ir anomaly from a cometary impact would be about 1% of the size of the Ir anomaly from an asteroid impact of similar mass. Because the Ir anomalies from asteroid impacts are difficult to measure, the Ir anomalies from cometary impacts would usually be below background levels and hence undetectable. We do know of impact events that have no discernable Ir anomaly. Only one of the five known strewn fields of tektites has a significant Ir anomaly (Glass, 1980; Koeberl, 1993). Because tektites form during impacts, this may indicate that the majority of impact layers do not have significant Ir anomalies. Another important difference between cometary and asteroid impactors is their crater size distribution. As impact craters become larger, the relative proportion of those produced by cometary impactors increases. Hughes (1998) suggested that around 95 % of the impactors producing craters between 20 and 100 km in diameter are comets. For craters over 100 km in diameter, the proportion of comets should be even higher. This means that comets, not asteroids, produce most global impact layers and that most global impact layers will not have detectable Ir anomalies. Characteristics of Submarine Global Impact Layers: Sedimentary Structures The most important immediate effect of a submarine impact would be the generation of tsunami waves whose size would be directly dependent on the size of the impactor. An impact that could produce a global impact layer from an 80 km diameter impact crater would require an impactor of c. 4 km in diameter (Melosh, 1989). The tsunamis produced in this event would have local heights on continental shelves of 1 km or greater (Artemieva and Shuvalov, 2002). The erosive power of such tsunamis would generate large ripped up blocks, cross-beds, profound disconformities, regional redeposition, and reworked sediments (Hassler et al., 2000; Hassler and Simonson, 2001). Hassler and Simonson (2001) described ripped up blocks of iron-formation that are > 2 m across. These large blocks are inferred to result from an impact that produced a crater c. 210 km in diameter (impactor c. 10 km in diameter). Massive rip-up blocks that are many metres in length are often misidentified as intact outcrops; this initially occurred with the Alamo impact breccia (J. Warme, 2003, pers. comm.). The problem would be even greater in deformed Archaean terranes. Careful structural measurements are required to identify massive rip-up blocks in poorly exposed Archaean cratons with highly metamorphosed

1.4. Early Precambrian Impact Events

49

strata. In addition, geologists should check carefully to see if smaller fragments surround the large blocks. An intact, well-bedded sedimentary formation that suddenly becomes part of an apparently conformable shear zone or a melange could actually be rip-up blocks from an impact event. If the shear zone or melange does not appear to have a higher metamorphic grade or to be more altered than the well-bedded sediments, then it could be an impact layer. In sand-sized sediments at continental shelf depths, tsunamis can produce antidunes. Antidunes are found today in the outflow zones from areas where large lakes of glacial meltwater were impounded (Waitt, 1985). Antidunes migrate upstream against the current, so their apparent downstream edges will be on the opposite side of the structure from the palaeocurrent direction. Antidunes are also extremely large, with 3-15 m heights and 100-325 m wavelengths (Bretz, 1969). Thus, any cross-beds in Archaean continental shelf sequences with amplitudes more than 3 m could be part of an antidune sequence. The erosive power of a 1 km amplitude tsunami is profound. As a result, the underlying sediments will be eroded. At large distances from the impact or in deeper water, the impact will produce a disconformity with a broad age gap in the sediments. In the Phanerozoic rock record, such disconformities can be identified using micropalaeontology. In the Archaean, these disconformities can only be identified if the geochronology is good enough and deposition was rapid. Although disconformities are difficult to identify in the Archaean rock record, erosion from tsunamis produces other features that might be preserved. Tsunamis often redeposit marine sediments within terrestrial sequences that are tens of meters above sea level and tens of kilometres inland (Bourgeois et al., 1993). The easiest tsunami deposits to identify are thin, graded layers of coarse clastic rocks that occur within otherwise continuous sequences of carbonate, chert, or iron-formation. Carbonate, chert, and iron-formation require depositional environments nearly free of terrestrial detritus. Therefore, the sudden appearance of coarse clastic lithologies within fine-grained sedimentary sequences could have resulted from strong wave action that moved distal clastic sediments into the area. During the Phanerozoic, some of these deposits may have been the result of large earthquakes. However, most tsunami deposits from large earthquakes are only a few centimetres in thickness and are typically gravel-sized or finer (Bourgeois and Johnson, 2001; Bourgeois et al., 2001). In contrast, impact tsunami deposits can be metres to tens of metres thick with a basal conglomerate containing rip-up clasts from the underlying beds (Takayama et al., 2000). Thus, any >~ 1 m thick coarse clastic layer within an Archaean carbonate, chert, or iron-formation sequence should be studied to see if it came from an impact-generated tsunami. The uppermost layer of most impact deposits should be an air-fall deposit. If this is the case, the layer will contain a few parts per thousand of impact spherules, characteristic impact minerals (discussed later), and microtektites. Impacts onto sedimentary substrates produce a smaller amount of impact melt (Pierazzo et al., 1997). From our own observations and those of others (Hassler et al., 2000), we infer that spherules resulting from impacts onto (marine) sediments are smaller than spherules resulting from impacts onto bare rock. We calculate that the impact spherules and microtektites from abyssal craters

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Chapter 1: The EarlyEarth

over 80 km in diameter should only be about 130 ~m in diameter. The few Archaean deposits that have been found contain medium to coarse sand-sized (250-1000 l.tm) impact spherules (Simonson et al., 1999) (section 1.3). These spherules can be seen with a hand lens, but they are not typical. An impact crater about 300 km in diameter is the minimum size needed to produce spherules that are visible with a hand lens (c. 250 ~tm particles). Therefore, only about 92 of the estimated 350 Archaean global impact layers should contain spherules that are visible using a hand lens. Thus, a careful consideration of sedimentary structures and how they make sense in their stratigraphic setting could produce an abundant return in newly identified impact layers within rock sequences. Unless new methods are developed to quickly magnify and image small grains in rocks, verifying these impact layers will require intensive petrographic investigation.

Geophysical Methods for Finding Impact Layers and Impact Craters During a large impact, the country rock is melted to produce an impact melt body that is usually quite magnetic. All impact craters larger than 40 km in diameter have large central magnetic anomalies (Pilkington and Grieve, 1992). Impact debris layers have high magnetic susceptibilities, due to their high magnetic mineral contents. For example, Worm and Banerjee (1987) found that the K/T impact layer has about 8-15 times higher magnetic susceptibility than the surrounding sediments. In our own studies, the contrast in magnetic susceptibility between impact layers and the surrounding sediment depends upon the magnetic susceptibility of the surrounding sediment. In red clay sediments, we find that the upper contact of the impact layer with the overlying sediment occurs at magnetic susceptibility values of about 40 • 10 -6 cgs (Fig. 1.4-1). The magnetic susceptibility within the impact layer is at least 40 • 10 -6 cgs and can be as high as 200 • 10 -6 cgs. If the sediment is a carbonate rather than red clay, the susceptibilities of impact layers will be somewhat lower. Because carbonates have a negative magnetic susceptibility, carbonate rocks and fossils dilute the high magnetic susceptibility of impact layers. In high carbonate sediments, the upper contacts of impact layers occur at susceptibility values of about 20 • 10 -6 cgs, with increases to values as high as 100 • 10 -6 cgs within the impact layer. Ash layers also have magnetic susceptibilities above 40 x 10 -6 cgs. Careful inspection of the coarse fractions (> 63 and > 125 ~tm) of a high susceptibility layer will show if the layer is an ash layer or an impact layer. Ash layers

Opposite: Fig. 1.4-1. Left: Magnetic susceptibility versus depth, core RC11-200. Note the increase in susceptibility to values above 40 • 10-6 cgs between 56 and 62 cm. The background magnetic susceptibility is about 10 • 10-6 cgs. Right: Photo of abyssal impact layer in split core. Abyssal impact layer containing several hundred impact spherules is between 56 and 62 cm. Abyssal impact layer is darker than most of the core. Note soft sediment deformation at 65-98 cm depth. This soft sediment deformation could be due to rapid deposition of the impact layer or to impact-related seismic activity.

1.4. Early Precambrian Impact Events

51

RCl1-200

Mag Sus, cgs (1.0E-6) 0 0

0

0

E 'r 0

Q. 4) r

0

0 0

50

100

52

Chapter 1: The Early Earth

often contain biotite or pumice, both of which are typically absent from abyssal impact layers. Preservation of high magnetic susceptibilities for Archaean sedimentary rocks depends on the degree of alteration. Because high magnetic susceptibility is primarily due to the occurrence of magnetite in the impact layer, high-grade sediments containing diagenetic magnetite might be mistaken for impact layers. If the magnetite in impact layers is removed or altered by weathering or metamorphism, the impact layers might lose their high susceptibility signature. Although the high magnetic susceptibility of impact layers might disappear if the sediment was to be metamorphosed or altered, study of drill-core samples of Archaean low metamorphic grade sedimentary rocks could yield new discoveries of impact layers. If the sedimentary facies and structures suggest possible impact events (as discussed earlier), scientists should measure magnetic susceptibility. A 30 m long core can be measured at 2 cm intervals easily in one day, whereas petrographic examination of the same core would constitute a more substantial project. Examining a 30 m core with a hand lens would take many hours and might miss dispersed layers of impact spherules and impact minerals. For this reason, the magnetic susceptibilities of Archaean sedimentary rocks from low-grade sequences that might contain impact layers should be measured.

Characteristic Rock Types and Features: Microtektites, Pseudotachylites, and Spherules Tektites consist of impact melt that may have travelled between 20 and 80 km into the stratosphere (Koeberl, 1992). Larger tektites have aerodynamic forms including teardrop shapes, flattened concave disks, and dumbbells (Fig. 1.4-2). The microtektite in Figure 1.4-2d is teardrop-shaped and shows surface ablation that is characteristic of tektites. However, most tektites from abyssal impacts are spherical microtektites (McCall, 2001; Figs. 1.4-2b, c). Some of the abyssal microtektites may be transitional forms from microfossils to true microtektites (Fig. 1.4-2a). Some microfossils have shapes that approximate teardrops and may be mistaken for tektites. Some weathering of volcanic glass can produce spalling of glass that resembles the sculpturing of microtektites by atmospheric erosion. Thus, microtektites such as that in Figure 1.4-2a often require further verification by thin section work. In contrast, round microtektites (e.g., Figs. 1.4-2b, c) are much easier to verify. Most abyssal microtektites are composed of perfectly clear, nearly pure silicate glass. They are much richer in SiO2 than microtektites from on-land craters (Taylor and Levinson, 1970). Their compositions mean that microtektites will resist alteration and are likely to survive in the geological record. A logical source for SiO2 in abyssal microtektites is silica-rich material such as chert or biogenic silica. Archaean cherts are relatively common, although they have somewhat higher Mn and Fe contents than Phanerozoic chert layers (Veizer et al., 1989). These Archaean cherts are a potential source of silica-rich microtektites in Archaean impact deposits. Simonson has found tektite-like objects in some Proterozoic and Archaean impact layers (Simonson et al., 2000a) (section 1.3). Their microtektites have classic teardrop shapes and dumbbell forms. They occur along with much more abundant microkrystites.

1.4. Early Precambrian Impact Events

53

Fig. 1.4-2. SEM secondary-electron images of microtektites. (a) Transitional club-shaped tektite. RC12-49 704-706 cm. Scale bar is 50 ~tm. (b) Round microtektite with craters. RC11-205 0-4 cm. Scale bar is 50 ~tm. (c) Round microtektite with ablation. MSN151P 0-5 cm. Scale bar is 50 ~m. (d) Teardrop-shaped tektite with ablation. RC 11.-205 0--4 cm. Scale bar is 100 ~tm.

We have documented four different morphological and compositional types of impact spherules from our abyssal impact layers. Depending on the spherule composition, the outer surfaces of these spherules have various sorts of quench textures. The microtektites discussed above are spherules with high silica contents (> 95% SIO2). These spherules are about half the size of the other spherules in our impact layers. We have also found spherules with quench textures that resemble miniature columnar jointing (Fig. 1.4-3). In visible light at 50 times magnification, the spherules are shiny, yellow, and round. In images taken with a scanning electron microscope (SEM), spherules etched in KOH to remove the outer layer show quench textures that resemble columnar jointing. Like columnar joints, the best-developed and most symmetrical features are beneath the surface. Also like columnar joints, the joints do not continue to the bottom of the rock layer (Fig. 1.4-3a). Furthermore, the quench textures of these spherules resemble the quench textures on the surfaces of spherules produced during experimental impacts (Evans et al., 1994). These columnar-textured spherules have an overall composition that resembles a highFe red clay. The high K contents of these rocks indicate that they are terrestrial in origin

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Fig. 1.4-3. Spherules with miniature columnar jointing. All spherules are from RCI 1-200, 55-62 cm. Note that the size of the jointing varies across individual spherules. This variation is inconsistent with a biogenic origin for the jointing. (a) Cross-section of polished spherule. Note that jointing occurs only on the edge and does not extend to the middle of the spherule. (b) and (d) SEM secondary-electron images of columnar spherules. Scale bar is 50 ~m. Spherules in b and d were treated with KOH to remove the outer coating and exposed the jointing. On upper left centre of d, some of shiny outer coating still remains. Note that d has a tektite-like shape. (c) SEM backscattered-electron image. Scale bar is 50 ~tm. Spherule c has cocoliths on its surface that partially obscure the jointing.

(Taylor et al., 2000) and are not cosmic spherules. Due to their high Fe contents, these spherules weather readily, and they would be less likely to be preserved in Archaean sediments than more silica-rich impact spherules. However, in an Archaean chert layer, they might be the most readily identifiable impact spherules. We also have found white spherules with quench textures that resemble miniature spinifex textures in SEM images (Fig. 1.4-4). These spinifex microkrystites have a terrestrial origin, with K contents of a few percent and an overall composition similar to that of a high-K amphibole or pyroxene. For comparison, Figure 1.4-4d shows a thin section of Archaean spinifex texture. The enlargement (Fig. 1.4-4b) of a small part of the spherule

1.4. Early Precambrian Impact Events

55

Fig. 1.4-4. Spinifex textured spherules. (a) SEM backscattered-electron image of spherule from ODP41-1R3W, 145-147 cm. Scale bar is 100 gm. (b) Enlarged backscattered-electron image of radiating aggregate in left centre of (a). Scale bar is 10 gm. (c) Backscattered-electron image of spherule from RC11-203, 1294-1296 cm. Scale bar is 50 ~tm. (d) Spinifex texture in thin section of V-8 komatiite from the Barberton greenstone belt.

in Figure 1.4-4a shows a strikingly similar texture. A major difference is the size, with komatiite-spinifex textures (see also, section 4.3) being typically much larger. However, microspinifex texture has also been found in some komatiites (Puchtel and Humayun, 2001). Therefore, there is no impediment to a wide range of sizes of spinifex texture. Furthermore, spinifex texture is not confined to komatiites, but has been documented in harzburgites, impact carbonates, and in native sulphur (Imai and Geshi, 1998; Nixon et al., 1999; Jones et al., 2000). Thus, spinifex texture is more indicative of rapid cooling and crystallisation than it is of a specific rock composition. The last type of impact spherules we have found contains randomly oriented crystal aggregates (Fig. 1.4-5). The crystals are most likely to be plagioclase feldspar. The microkrystites have a few percent K, which precludes an extraterrestrial origin. Some of the spherules appear to be hollow (Fig. 1.4-5b). Such spherules might be analogues of the originally hollow Archaean spherules with outer crystal layers (Fig. 1.4-5d) found by Simonson and Hassler (2001). One of our crystal aggregates has some large crystals that jut out of the

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Fig. 1.4-5. Crystal aggregate spherules. All but (d) are SEM backscattered-electron images; (b), (c), (e) and (f) are spherules from ODP 41-1R3W, 145-147 cm. (a) Scale bar is 50 ~tm. Spherule from PLDSII1P, 0-2 cm. (b) Scale bar is 35 ~tm. Note that this spherule appears to be hollow. (c) Scale bar is 50 ~m. (d) Thin section of Archaean spherule (after Simonson and Hassler, 2001). The centre of this spherule may originally have been hollow and the cavity in-filled later. (e) Transitional spherule with small crystals and some columnar jointing. Scale bar is 50 ~m. (f) SEM backscattered-electron image of spherule with protruding crystals. Scale bar is 35 ~m.

1.4. Early Precambrian Impact Events

57

Fig. 1.4-6. Colours and surface textures of impact spherules. All photographed in visible light at 50x magnification. (a) Silver coloured metal spherule from V20-36, 61 cm. Spherule is hollow. (b) Yellow and white spherules from JYN5-46P, 14-16 cm. (c) All from ODP 41-1R3W, 145-147 cm. (cl) White spherule (has random crystals in SEM image). (c2) Yellow to white spherule (SEM images in Figures 4a and b). (c3) Black spherule (has columnar texture in SEM image). (c4) White spherule (SEM image in Figure 5b). (d) Red spherule. (e) Clear spherule (microtektite). RC12-228, 153 cm. (f) Green spherule. V29-66, 0-4 cm (columnar texture in SEM image).

spherule (Fig. 1.4-5f). Another crystal aggregate has a texture that is transitional between random crystal aggregates and columnar jointing (Fig. 1.4-5e). In visible light at 50 times magnification, both the spinifex and random types of microkrystites appear white and perfectly round (Figs. 1.4-6b, c4). Because they are similar in size and colour to marine microfossils, these white impact spherules are easy to miss in Phanerozoic sediments. There are two key tests that can be performed using a microscope. Microfossils are hollow and easily crushed, whereas impact spherules are usually solid. If the impact spherules are not solid, the hole in the centre has a diameter that is less than half the overall diameter of the spherule. In contrast, microfossils have a very thin outer layer. Impact spherules can be crushed, but they are usually slightly magnetic; when compressed

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Chapter 1: The Early Earth

with a metal pick, they either stick to the pick or jump away. As there are no microfossils in Archaean sedimentary rocks, these impact spherules should be readily identifiable. We have found many colours of impact spherules: white, yellow, green, brown, orange, black, and metallic silver (Fig. 1.4-6). Each impact layer seems to have characteristic spherule compositions and morphologies. Thus, the four types of abyssal impact spherules outlined above are unlikely to represent an exhaustive compilation of impact spherule morphologies. For example, some impact spherules resulting from the Eltanin abyssal impact contain vesicles and others have included spinel grains (Gersonde et al., 1997; Kyte, 2002), whereas others appear to be pure metal (Fig. 1.4-6a). Although most spherules are yellow or white in colour, some are black, brown, orange or green (see Fig. 1.4-6). As more abyssal impact layers are identified, the range of documented spherule morphologies and colours is likely to grow. Impact events may also produce widespread melting of the country rock (Dressier and Reimold, 2001). The distribution of melting is very heterogeneous, and may produce pseudotachylites. Pseudotachylites are veined rocks resulting from impacts in which all of the veins have exactly the same composition as the surrounding matrix. The darker colour of pseudotachylite veins is due to the finer grain size of the glass that makes up the veins. Pseudotachylites have been recognised and described at the c. 2 Ga Vredefort impact crater, Kaapvaal craton (Colliston et al., 1999; Reimold and Colliston, 1994).

Characteristic Impact Minerals Diamond and shocked minerals Once a probable impact layer has been identified, it must be verified as a bona fide impact layer. To do this, one must find minerals that can only form at impact sites (Dressier and Reimold, 2001). There are three such minerals, each has a characteristic form: shocked quartz, shocked zircon, and impact diamond (Miura, 1994). Shocked quartz and shocked zircon show planar deformation features (section 1.3) that are the result of deformation along crystal planes during impact. The result is absolutely parallel planes that cut across the crystal and occur in characteristic crystallographic directions. The two most common directions are co{1013 } and Jr {1012} (Bohor et al., 1987). To qualify as shocked quartz, the parallel planes should occur in at least two different directions. Diamond resulting from shock metamorphism is very different from the diamond found in kimberlite pipes (Shelkov et al., 1998; Masaitis et al., 1999). Impact diamond is not always white; it can be black and have a metallic lustre under reflected light. Impact diamonds are usually only a few microns in size and have a characteristic hexagonal or cubic crystal form. The hexagonal form of impact diamond, or lonsdaleite, has characteristic convex crystal faces. Impact diamonds have a grain density of about 3.5 g/cm 3, higher than that of graphite, which has a grain density of about 2.23 g/cm 3. Impact diamond is very hard and will scratch all other minerals. Diamond is also nearly insoluble and can be separated out by dissolving the rest of the rock-forming minerals. Raman spectroscopy is most effective for distinguishing black-coloured impact diamond from graphite.

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1.4. Early Precambrian Impact Events

Typical abyssal sediments have about 0.5-1% organic carbon, sufficient to provide the C for impact diamond. Because most of the larger impactors are comets, the abundant C from comets can also be a precursor to impact diamond. Some segregations of nearly pure C occur in our identified abyssal impact layers, but we have not yet located impact diamond. Nevertheless, impact diamond is more likely to form in abyssal impacts than is either shocked quartz or shocked zircon. Native elements

Impact events are extremely reducing (Yakovlev and Basilevsky, 1994). This may be due to the formation of plasma in the impact vapour cloud. Plasmas strip all the electrons from atoms, thus reducing them to their lowest charge state. This means that impact layers could contain reduced mineral forms and native elements. Because they usually have a metallic lustre and no oxygen, native elements can be readily identified. Of the native elements, native carbon (diamond and graphite) has been found at most impact sites (Table 1.4-1). Native S, Fe, Pt and Ni have also been found. Experimental results predict that native A1, Ca, K, Co, Mn and Mg should also form during impacts (Yakovlev and Basilevsky, 1994) but none have yet been found. The most durable and the most abundant of these native elements are native Fe, C, and Ni. Placer diamonds have been found in the Archaean Witwatersrand Supergroup (Nisbet, 1987). Thus, impact diamonds should be preserved within low-grade Archaean sequences, even if those sequences are not reducing. In low-grade black and brown shale sequences, it is likely that native Fe and Ni would also be preserved.

Table 1.4-1. Native elements found in impact layers Element Pt Fe C C C C C C C C C C C S S Ni

Source Crater Morokweng Popigai Popigai Ries Sudbury Chixulub Kara Puchezh-Katunki Zapadnaya Ternovka Boltysh Ilintsky Lappajarvi Sudbury Chixulub Popigai

Forms Globules and segregations Globules and segregations Diamond Diamond Diamond Diamond Diamond Diamond Diamond Diamond Diamond Diamond Diamond Massive or granular Massive or granular Globules and segregations

References Andreoli et al., 1999 Vishnevsky, 1975 Visknevsky et al., 1975 Hough et al., 1995; Rost et al., 1978 Masaitis et al., 1997 Carlisle and Braman, 1991 Masaitis et al., 1999 Khakhaev et al., 1994 Gurov et al., 1996 Gurov et al., 1995 Gurov et al., 1995 Gurov et al., 1995 Langenhorst and Masaitis, 1996 Heymann et al., 1997a Heymann et al., 1997b Masaitis and Sysoev, 1975

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Chapter 1: The Early Earth

Fig. 1.4-7. Platy magnetite grains, feldspar crystals and impact spherules. (a) and (b): SEM secondary electron images. (c) and (d): Photographs in visible light at 50x magnification. (a) Entire magnetite grain. RC 11-205, 0-5 cm. Scale bar is 300 l~m. (b) Surface detail of magnetite grain in a. Scale bar is 10 ~m. (c) Magnetite grains lined up on-end by a magnet beneath the slide. White circles are radiolaria and sponge spicules. V20-36, 362 cm. (d) Clear feldspar crystal (right) next to two impact spherules on a 63 I~m sieve. Spherules are left bottom and left centre. RC11-200, 59 cm. Abyssal impacts We have found extremely large quantities of platy magnetite in ejecta layers from abyssal impacts (Figs. 1.4-7a, b, c). Although magnetite can be an important component of heavy mineral sands, impact-generated magnetite is different in form. Because impact magnetite is deposited from the air, it is extremely platy, with an aspect ratio of about 15:1 (Fig. 1.4-7c). In contrast, detrital magnetite is rounded and abraded during sedimentary transport and will as a result commonly have a 1 : 1 aspect ratio. In addition to its differing morphology, detrital magnetite is concentrated in heavy mineral sands with other heavy minerals such as zircon, sphene, ilmenite, and garnet. It does not occur generally in silty or clay-rich sediments. Abyssal impact layers we have investigated are dominated by clay- and silt-sized grains. In most abyssal impact layers, magnetite grains are the coarsest component (typically 63-500 ~m in width) and there are very few quartz grains. Most of these quartz grains are very small, typically just over 40 I~m in diameter. In contrast, detrital quartz and heavy minerals found with detrital magnetites are typically sand-sized with no clay matrix.

1.4. Early Precambrian Impact Events

61

Large, clear feldspar crystals also appear in some impact layers. The feldspar crystals, typically plagioclase, are generally a few millimetres in length, dwarfing the c. 180 ~tm impact spherules (Fig. 1.4-7d). Some of the feldspar grains are zoned. The feldspar grains are very fragile and often fall apart if touched. Both the magnetite and feldspar grains we have found are often significantly larger than the impact spherules. Thus, they may be the first minerals to be recognised in thin sections of abyssal impact layers. Sedimentary layers with abundant platy magnetite and perfectly crystalline feldspar grains should be checked carefully for impact spherules and microtektites.

Stratigraphic Associations that May Be Characteristic of Impacts The science of finding impact deposits in Archaean sediments is extremely new and we have yet to find all of the possible types of deposit that might be characteristic of Archaean impact events. It is justified therefore to speculate about other sedimentary and volcanic associations that could be characteristic of impact events. If impact sediments are characteristically reduced and contain reduced compounds, distal airfall deposits may often appear as brown to black shale layers. We have also postulated that large impacts may increase the intensity of existing mantle plume volcanism (Abbott and Isley, 2002a). If an increase in melting occurred, the resulting rocks would be komatiites and basaltic komatiites (see also section 4.3). Thus, if a black shale layer is overlain by extrusive komatiitic to basaltic komatiitic rocks, the shale should be examined for impact debris. The reduction produced by impacts will also affect the degree of reduction of ironformations. Iron-formations commonly show lateral changes in oxidation state that have previously been attributed to a change from more oxidising shallow water to reducing deep water conditions (James, 1954) (see section 5.4 for full discussion). However, it is possible that this lateral change in the oxidation state actually reflects the distance of iron-formation deposition from an oceanic impact crater. Sulphide-facies iron-formation could form closest to the crater, and the oxide-facies iron-formation furthest away. If sulphide facies ironformation reflects proximity to an impact event, it is the most likely of all the types of iron-formation to contain relatively large clasts of impact breccia and concentrations of impact spherules. In addition, jadeite has been found in at least one impact layer (Odette, 1969). Jadeite and omphacite form when basalt is metamorphosed under high pressure to form eclogite. Most impacts onto continental crust may not have hit basaltic targets, but abyssal impacts will almost invariably have done so. If the impactor had a high enough energy, it would have penetrated through the pelagic sediment cover and into basaltic basement. Thus, the largest abyssal impacts could produce both jadeite and omphacite from the high-pressure breakdown of basaltic oceanic crust. If these minerals are present in any quantity, they will be quite distinct visually. Jadeite is bright green, and omphacite is pink. Consequently, there is a distinct possibility that Archaean impact layers might sometimes have prominent pink and green minerals within them. Because jadeite and omphacite are not common as detrital minerals, they could possibly serve as markers for impact horizons.

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Chapter l: 7"he Early Earth

Summary Impact debris layers should be common in early Precambrian sedimentary rocks. We estimate that there should be at least 350 global-scale layers that range in age from 3.8-2.5 Ga. Most of these layers would be from impacts into the abyssal ocean. We have studied Phanerozoic abyssal impact ejecta in deep-sea sediment layers in order to identify their distinguishing characteristics. These layers usually contain: (1) platy magnetite, (2) microtektites, and (3) microkrystites. The most common high-pressure mineral from abyssal impacts would be impact diamond rather than shocked quartz or shocked zircon. Distal abyssal impact layers often have sharp contacts, internal laminations, dark colours, and high magnetic susceptibility. The uppermost 10 cm of unusual sedimentary layers with high magnetic susceptibility values should be examined for impact spherules, exotic minerals, and other debris that characterise Phanerozoic sediment layers resulting from abyssal impacts. A set of features within Archaean rock suites that is characteristic of impacts has been delineated. These impact characteristics range from small-scale features like impact spherules, shocked quartz, and impact diamonds to large-scale rip-up blocks and sedimentary structures from tsunami waves. We have also described characteristic mineral associations, rock types, and textural features that should be examined carefully for impact debris and impact spherules. These suggestions result from our studies of Phanerozoic impact events and in particular of Phanerozoic impact events within the ocean basins.

1.5. COMMENTARY D.R. NELSON The short-lived nuclides have provided remarkably detailed chronologies of the formation and differentiation of the planets, as was outlined in section 1.2. They indicate that the Earth's accretion, and large-scale differentiation into metallic core and silicate mantle, was largely completed within c. 20 My of the time of formation of our solar system. However, during the latter stages of the Earth's accretion and probably for at least a further 100 My, it is likely that the planet was reworked on a global scale many times, by collisions with substantial (Moon- and Mars-sized) impactors. Late-accretion collision was the dominant process determining the state of the early Earth and conditions on its surface for it's first 100 My or more, with major collision events destroying or partially resetting any evidence of earlier differentiation processes. The accretionary flux decreased rapidly during the first few hundred million years, with possible short-lived spikes in the flux (i.e., the late heavy bombardment at c. 3.85 Ga) due to the deflection of asteroids or comets from the outer Solar System into geocentric orbit (see Hartmann et al., 2001). The physical state of the Earth and the physicochemical conditions on its surface during the lengthy time interval prior to c. 3.8 Ga largely belong within the realm of speculation. Despite the almost complete absence of any firm constraints, there is little doubt The Pmcambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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63

that physicochemical conditions on the Earth's surface changed more dramatically during this c. 770 My time interval than during any other time. It would, however, be an oversimplification to assume that the Earth evolved in a regular and orderly fashion in one direction towards conditions resembling those of today--from hot to cool surface temperatures, for example--during this interval. In section 1.2 it was argued that, during at least the first hundred million years following the Earth's accretion, its surface temperatures were probably so high that water resided mostly in the atmosphere as steam, rather than in the oceans. An alternative view, based on elevated g 180 values obtained for > 4.0 Ga detrital zircons from the Yilgarn craton of Western Australia, has been advanced by Wilde et al. (2001), Mojzsis et al. (2001), Peck et al. (2001) and Valley et al. (2002). These authors argued that high g180 values determined for some of these ancient zircons indicated that they crystallised from melts derived from high-180/160 sedimentary or low-temperature hydrothermally-altered source rocks. Based on one elevated ~ 180 value (7.4 + 0.7%0, 20" uncertainty) obtained from grain W74/2-36, from which the oldest 2~176 date of 4404 + 8 Ma (20" uncertainty) was also obtained, Wilde et al. (2001), Peck et al. (2001) and Valley et al. (2002) surmised that the Earth must have been cool enough for liquid water to exist on the Earth's surface by c. 4.4 Ga. Some inconsistencies in these arguments were pointed out in section 1.2; it was argued that the high 6180 values of these ancient zircons were of secondary origin and therefore unrelated to the 180/160 characteristics of the melts from which the zircons crystallised. What is currently known of the influence of major impact events on the Earth during the Precambrian was summarised in section 1.3. Although progress has been rapid, section 1.3 reveals that this field of investigation is in its infancy and there is still much to learn. There are no known terrestrial impact structures older than c. 2.0 Ga, so any future assessment of the direct influence of extraterrestrial impacts on the Earth's early evolution and on physicochemical conditions on its surface will be largely dependent on the recognition and examination of impact-related horizons within the Archaean sedimentary rock record. Since the pioneering study of Lowe and Byerly (1986), comparatively few impactrelated sedimentary layers from within Archaean to early Palaeoproterozoic successions have been identified (see Table 1.3-1). As pointed out in Section 1.3, the earliest interpretation of spherules as resulting from meteorite impacts was initially regarded skeptically, although this is now more widely accepted. Such acceptance has arisen largely because of the improved characterisation of the diagnostic features of impact-related sedimentary horizons. Section 1.4 outlined the diagnostic sedimentary, mineralogical and geophysical features of distal impact layers and proposed an improved strategy for their recognition. Armed with this new set of diagnostic criteria, it is hoped that a better appreciation of the important but commonly ignored role of meteorite impacts on the Earth's early geological development can be attained.

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The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology; Vol. 12 (K.C. Condie, Series Editor) Published by Elsevier B.V.

65

Chapter 2

GENERATION

2.1.

OF CONTINENTAL

CRUST

INTRODUCTION

W.U. MUELLER AND D.R. NELSON As with many aspects of tracing and documenting Earth's evolution, controversy arises, and the generation of the continental crust is no exception. Implicit in the formation of continental crust, is the development and formation of granite-greenstone terranes because they generally constitute nuclei of most modern-day continents. Many studies favour an episodic increase in the volume of continental crust from the Early Archaean (Taylor and McLennan, 1985; McLennan and Taylor, 1991; Condie, 1998), whereas Armstrong (1991) advocated that most of the continental crust had been generated prior to 3 Ga, but was lost from the geological record due to recycling into the mantle (section 2.8). The latter point had not gone unnoticed by Dimroth (1985), who demonstrated from mass balance calculations that magma emplacement into Archaean and modern arcs exceeded the growth of continental crust by a factor of two, therefore implying that significant recycling of continental crust back into the mantle must have occurred. Evidently, the results from geochemical studies of sedimentary rocks and from the stratigraphy of Precambrian volcano-sedimentary sequences of preserved but fragmented terranes, still demand a verdict, yet new geochemical studies of a 3.4 Ga norite dyke from the Isua area shows that crustal recycling was a viable process (Nielsen et al., 2002). By what process or processes are large amounts of continental crust produced? This contentious issue revolves around the concepts of mantle plumes (sections 3.2 and 3.3) and subduction. Most of the focus of this chapter deals with the mechanisms or large-scale geodynamic structures that may have contributed to the formation of continental crust during the Archaean. Evidently, depending on the study area, different models are favoured. Mantle plumes may dismember continents efficiently but can also accrete material to the crust by producing volcanic flood provinces. In contrast, horizontal plate movements related to subduction, which include the various phases of the Wilson cycle, generate most of the volcano-plutonic material at active margins and mid-oceanic rifts (Fisher and Schmincke, 1984). Subsequent subduction-related tectonics may accrete newly formed crust onto continental margins, but was this the prevalent mechanism during the Precambrian and especially during the Archaean? Superplume events, advocated to have had major influence on pre-1.9 Ga Earth evolutionary events (Condie, 1998, 2002; Nelson, 1998a), primarily record continental breakup and can be monitored partially by using major magmatic episodes (Nelson, 1998a), black shales (Condie et al., 2001) and banded iron-formations (Isley and Abbott, 1999)

Chapter 2: Generation o f Continental Crust

66

as tracers (Condie, section 3.2). Major upwelling due to plumes contains the notion of large-scale diapirism, yet in many greenstone belts, unequivocal structural evidence for all-encompassing vertical tectonics controlling the belt's evolution remains elusive. A revival of vertical-dominated tectonics, a modification of the "sagduction theme", has been considered for parts of the Pilbara (see section 2.6) and Indian cratons (Chardon et al., 1998). In contrast, the Late Archaean cratons of North America show a new wrinkle with the coeval evolution of plume-generated komatiites and subduction-related arc volcanism. Evidence for plume-arc interaction occurring over a protracted period of c. 30 My is best recorded in the Abitibi greenstone belt (Dostal and Mueller, 1997; Mueller and Mortensen, 2002; Daigneault et al., section 2.4), and Polat and Kerrich (2001) support the concept of subduction-accretion complexes with intervening komatiite plume volcanism for the southern Superior Province. The Yilgarn craton reveals a similar history. In the eastern part of the Yilgarn craton, an elongate rift formed at c. 2705 Ma, and was filled rapidly with komatiitic and basaltic lavas. By c. 2670 Ma, subduction and continent-continent collision had closed the rift and initiated regional deformation. This was followed by widespread granitoid plutonism and metamorphism. Thus, the regional deformation episodes preserved in many Late Archaean granite-greenstone terranes were apparently not due to the ascent and lateral movement of the granitic plutons resulting from melting of older crust by a mantle plume, as argued by Campbell and Hill (1998) for greenstones of the Yilgarn craton, but reflect tectonic plate interactions. Evidently, geodynamic plume-arc events played an important role in the Late Archaean. This does not, however, preclude the possibility that diapirism played a more significant role in the formation of granite-greenstone terranes during the Early Archaean, and that its significance may have diminished with time. The contributions in this chapter display the diversity of opinions on crust formation, granite-greenstone belt evolution and its possible controls.

2.2.

ISUA ENIGMAS: ILLUSIVE TECTONIC, SEDIMENTARY, VOLCANIC AND ORGANIC FEATURES OF THE > 3.7 GA ISUA GREENSTONE BELT, SOUTHWEST GREENLAND

J.S. MYERS Introduction

The Isua greenstone belt is located within a part of the Archaean gneiss complex of West Greenland (Bridgwater et al., 1976) that contains remnants of early Archaean tonalitic gneisses (Fig. 2.2-1). Some rocks within the belt are considered to provide evidence for the existence of life before 3.7 Ga (Rosing, 1999; Pflug, 2001). Early interpretations of the geology concluded that "the sequence exposed in the Isua greenstone belt is essentially of sedimentary origin" (Dimroth, 1982) and that it contains substantial amounts of shallow marine carbonates, clastic quartzites and felsic volcanic rocks (Nutman et al., 1984). More recently it was found that the carbonates are metasomatic in origin (Rose et al., 1996; The Precambrian Earth: 7k'mposand Events Edited by P.G. Eriksson, W. Altennann, D.R. Nelson, W.U. Mueller and O. Catuneanu

2.2. Isua Greenstone Belt: Features

67

Fig. 2.2-1. Simplified map of the Godthfibst]ord region of southwest Greenland showing the main geological units and the location of the Isua greenstone belt. Numbers indicate ages in Ga. Geology from Allaart (1982) and Myers and Crowley (2000).

Rosing et al., 1996), "felsic volcanics" reflect mylonitised tonalite and altered pillow lava, and the chloritic and amphibolitic schists that make up most of the greenstone belt were derived from basaltic pillow lava (Myers, 2001 a). These schists were isoclinally folded and refolded, tectonically interleaved, refolded into open folds, and then cut by dolerite dykes at c. 3.5 Ga. Nevertheless, several recent studies continue to underestimate substantially the amount of deformation and distortion. Primary features continue to be described that are in stark contrast to detailed recent structural and lithological re-interpretations of the relevant outcrops (e.g., Myers, 2001b for a summary). Recently described examples of

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Chapter 2: Generation of Continental Crust

Fig. 2.2-2. (a) Basaltic dyke, interpreted as pillow lava, and dyke contact interpreted as a thrust by Komiya et al. (1999), located on Fig. 2.2-2b. (b) Geological map of the northeastern part of the Isua greenstone belt and cross-section, from Komiya et al. (1999, their Fig. 9a).

2.2. Isua Greenstone Belt: Features

69

Fig. 2.2-2 (continued). (c) Geological map of the same area as Fig. 2.2-2b, by Myers (1998, unpublished). these controversial interpretations are briefly discussed below in the light of the significant new field evidence.

1998-2002 Interpretations Reconsidered Intra-oceanic accretionary complex Based on repetition of inferred oceanic plate stratigraphy, the northeast part of the Isua belt was interpreted by Komiya et al. (1999) as an intra-oceanic accretionary complex. Unfortunately, the origin of many rocks that are key elements of the model were misinterpreted and the effects of multiple episodes of intense deformation were not recognised. Intensely deformed rocks that were derived from basaltic pillow lava were misinterpreted as hyaloclastite and mafic turbidite. Some dykes were mistakenly interpreted as pillow lava and the sharp boundaries of some dykes were inferred to be thrusts (Fig. 2.2-2).

Pillow breccia A small outcrop of basaltic pillow fragments enclosed by quartz (located at 1 on Fig. 2.2-3) was interpreted as a well-preserved pillow breccia by Appel et al. (1998) and Fedo et al. (2001 ). The sides of the outcrop indicate substantial extension of the fragments and quartzfilled vesicles into rods. Fluid inclusions in unstrained quartz from the vesicles were interpreted as remnants of early Archaean seawater (Appel et al., 2001). However, the quartz in these vesicles must have recrystallised during and after repeated episodes of deformation. Therefore the fluid inclusions cannot be primary undeformed features and are probably of metamorphic or metasomatic origin. The inferred breccia may also be of tectonic origin. The outcrop does not appear to be located in a zone of low strain. On the contrary, it is situated adjacent to a zone of

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Chapter 2: Generation of Continental Crust

Fig. 2.2-3. New simplified geological map of a northeastern part of the Isua greenstone belt showing the locations of deformed pillow breccia (1) and polymictic conglomerate (2) in relation to fault zones of chloritic schist. intense late Archaean or early Proterozoic ductile deformation. The fragments are of early Archaean pillow lava, but the brecciation may be a result of deformation that post-dates the episodes of intense early Archaean ductile deformation.

Polymictic conglomerate Deformed fragments of quartz, metachert and amphibolitic schist enclosed in chloritic schist (located at 2 on Fig. 2.2-3) were interpreted as deformed polymictic conglomerate

2.2. Isua Greenstone Belt: Features

71

Fig. 2.2-4. New simplified geological map of a northeastern part of the Isua greenstone belt showing the extent of a narrow belt of complex schist and the location within it, at point denoted as 3, of "graded turbidite", shown in photograph. (Appel et al., 1998; Fedo, 2000; Fedo et al., 2001). This rock lies within a Late Archaean or Early Proterozoic shear zone where this zone crosses a unit of metachert and amphibolitic schist derived from pillow lava. Therefore the polymictic conglomerate appears to be tectonic in origin. Graded turbidite A complex layer of heterogeneous schist, interpreted as metasedimentary rocks (Nutman et al., 1984; Nutman, 1986; Rosing et al., 1996), lies within amphibolitic schist mainly derived from pillow lava. At one locality (3 on Fig. 2.2-4) the layer contains two thin layers of quartz-mica-chlorite-graphite schist that appear to have well preserved depositional grading (Fig. 2.2-4), and they have been interpreted as graded turbidite (Nutman et al.,

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Chapter 2: Generation of Continental Crust

Fig. 2.2-5. Quartz mylonite, previotsly interpreted as clastic quartzite, along an early Archaean fault zone in the southeastern part of the Isua greenstone belt.

1984" Nutman, 1986" Fedo et al., 2001). 13C-depleted carbon in graphite was interpreted as being of biogenic origin by Rosing (1999) and the best oldest evidence of life. Recent mapping (Fig. 2.2-4) indicates that this composite layer is not isoclinally folded and refolded like most nearby layers of chert, BIF and pillow lava. It appears to be an early Archaean fault zone associated with tectonic interleaving after episodes of isoclinal folding and refolding. The layers interpreted as graded turbidite contain rootless isoclinal folds and mylonitic fabrics (Constantinou, 2001) and are highly deformed rocks. The apparent sedimentary structures are thus misinterpretations. Clastic quartzite

A number of thin layers of quartzite were interpreted as deposits of clastic sedimentary origin (Nutman, 1986; Rosing et al., 1996; Fedo, 2000; Fedo et al., 2001). Recent mapping indicates that most of these layers are quartz mylonites in repeatedly deformed fault zones (Fig. 2.2-5). The quartz and quartz-fuchsite rocks probably originated as vein quartz. These fault zones post-date isoclinal folding and refolding of the greenstones and were related to tectonic disruption and intercalation of rock units that predated intrusion of c. 3.5 Ga dykes.

2.2. Isua Greenstone Belt: Features

73

Fig. 2.2-6. Mylonite derived from early Archaean tonalitic gneiss and amphibolite, located at position 4 on Fig. 2.2-4.

Microfossils Microfossils called Isuasphaera isua were described (Pflug, 1978; Pflug and JaeschkeBoyer, 1979) from the greenstone belt, as well as morphologies resembling capsules of iron bacteria called Appelella ferrifera (Robbins, 1987). Although substantially discredited by Bridgwater et al. (1981) and Roedder (1981) (see also section 6.2), these spherical siliceous microstructures were recently reiterated as "cellular structures of the Huroniospora-type (cyanobacteria)" by Pflug (2001). The samples came from metachert in a thick metachert and BIF unit at the northeastern end of the greenstone belt. The rocks were repeatedly deformed during the Early Archaean and then enormously constricted and rodded. Any primary spherical objects must now be extremely elongated, and so these circular microstructures cannot be original. They have recently been interpreted by Appel et al. (2003) as products of pre-Quaternary weathering.

c. 3.64 Ga mylonites Mylonites (Fig. 2.2-6; located at 4 on Fig. 2.2-4) in the northwest part of the greenstone belt were interpreted by Hanmer and Greene (2002) and Hanmer et al. (2002) as a c. 3.64 Ga thrust-nappe stack, and the oldest known evidence of"a modern structural regime .... similar to those of modern mountain belts". More detailed mapping indicates that the mylonites

Chapter 2: Generation of Continental Crust

74

post-date the c. 3.5 Ga dykes. These dykes cut across the adjacent greenstones but are cut by the mylonite stack and are transposed into the mylonitic fabric. Conclusions

New field investigations indicate that what were thought to be the best preserved examples of sedimentary, volcanic and biogenic structures, and an intra-oceanic accretionary complex, may not be of primary, Early Archaean origin. The Isua greenstone belt contains schists and mylonites derived from pillow lava and a small amount of chert and BIE All primary structures and textures are distorted. Understanding the original nature and significance of the components of the greenstone belt requires an understanding of the effects of deformation. In greenstone belts, as in gneiss terranes, superficial simplicity may be caused by multiple episodes of intense ductile deformation (see also section 7.4).

2.3.

GEOCHEMICAL DIVERSITY IN VOLCANIC ROCKS OF THE > 3.7 GA ISUA GREENSTONE BELT, SOUTHERN WEST GREENLAND: IMPLICATIONS FOR MANTLE COMPOSITION AND GEODYNAMIC PROCESSES

A. POLAT, A.W. HOFMANN AND EW.U. APPEL Introduction

Geochemical signatures in Archaean mafic to ultramafic volcanic rocks can provide important constraints on the thermal and chemical characteristics of their mantle source regions (Sun, 1984; Xie et al., 1993; Arndt, 1994; Condie, 1994a; McCulloch and Bennett, 1998; Jochum et al., 2001). Recent geochemical investigations of Archaean greenstone belts have documented a great compositional diversity in volcanic rocks, reflecting diverse source characteristics, petrogenetic processes, and tectonic settings in the early Earth (Arndt, 1994; Polat and Kerrich, 2001b; Wyman et al., 2002a; Polat et al., 2002). However, preservation of these mantle-derived mafic to ultramafic volcanic rocks diminishes from the Neo- to the Palaeoarchaean. They become extremely rare in terranes older than 3.6 Ga (Nutman et al., 1996). Preservation of primary volcanic features, including pillow basalts (see, however, section 2.2), in the oldest known, well-exposed > 3.7 Ga Isua greenstone belt, West Greenland (Fig. 2.3-1), provides an excellent opportunity to study the geochemical characteristics of Palaeoarchaean mafic to ultramafic volcanic rocks. In this study, for descriptive purposes, the western part of the Isua greenstone belt has been divided informally into three litho-tectonic sequences: outer arc, central arc, and inner arc (Fig. 2.3-2). In this contribution, we compare the geochemical characteristics of metamorphosed pillow basalts and associated flows from these three litho-tectonic sequences. Samples of the inner and outer arcs are from only the western part of the belt, and samples of the central arc are from both eastern and western segments of the belt. The objective of this study is to evaluate the significance of the geochemistry of the Isua volcanic rocks in The Precambrian Earth: Tempos and Events Fxtited by P.G. Eriksson, W. Aitcrmann, D.R. Nelson, W.U. Mueller and O. Catuneanu

2.3. Isua Greenstone Belt: Geodynamic Processes

75

Fig. 2.3-1. Simplified geological map of the Godthfibsfjord region, adapted from Myers and Crowley (2000).

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Chapter 2: Generation of Continental Crust

Fig. 2.3-2. Simplified geological map of the western part of the Isua greenstone belt (IGB), showing the location of samples, modified from Nutman (1986). The western part of the belt has been informally divided into three lithotectonic sequences: outer arc, central arc, and inner arc. The Isua greenstone belt has been remapped by Myers (2001a, b); however, a complete map of the belt has not been published yet. Inset shows the location of the Isua greenstone belt in southern Greenland.

terms of Palaeoarchaean petrogenetic processes, mantle source characteristics, and geodynamic setting. According to Nutman et al. (2002), the outer arc sequence is composed of c. 3800 Ma volcano-sedimentary rocks, whereas the central arc sequence is characterised by c. 3700 Ma supracrustal rocks. These dates were obtained from felsic rocks, the origin of which (felsic volcanic versus felsic intrusive) is a matter of debate (Myers, 2002; see also section 2.2); therefore, they may not be directly relevant to the formation of volcanic rocks discussed in this study. No age has been assigned to the inner arc sequence. Geochemical signatures in many Neoarchaean volcanic suites have been attributed to Phanerozoic-style volcanism, such as mantle plume, island arc, mid-ocean ridge, and back-arc volcanisms (Arndt et al., 1994; Condie, 1994b; Dostal and Mueller, 1997; Cousens, 2000; Corcoran and Dostal, 2001). Accordingly, in this study we assumed, in a broad sense, that the geochemical characteristics of the > 3.7 Ga Isua volcanic rocks have similar geodynamic significance to their Phanerozoic counterparts. This reasoning is based on the assumption

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77

that certain groups of elements (e.g., HFSE, REE) will behave consistently for a particular petrogenetic process throughout the Earth's history.

Geological setting The 3.7-3.8 Ga Isua greenstone belt is located in the Godth~,bsfjord region of southern West Greenland (Fig. 2.3-1; Nutman, 1986). The region contains extensive, well-exposed Palaeo- to Neoarchaean intrusive and supracrustal rocks (Nutman et al., 1996; Myers and Crowley, 2000). The Palaeoarchaean intrusive rocks are dominated by the 3.65-3.80 Ga Itsaq Gneiss Complex derived mainly from tonalities and granodiorites (Fig. 2.3-1). In addition, there are numerous Palaeoarchaean ultramafic intrusions within the Itsaq Gneiss Complex (Friend et al., 2002). Most of the Palaeoarchaean rocks were intensively deformed and metamorphosed during a Neoarchaean tectonothermal event, resulting in widespread modification of Palaeoarchaean structures and stratigraphic relationships. Palaeoarchaean volcanic and sedimentary rocks are best exposed in the Isua greenstone belt in the Isukasia area (Fig. 2.3-2). On the basis of U-Pb ages and field relationships, Nutman et al. (2002) and Hanmer et al. (2002) proposed modem plate tectonic-like models for the Palaeoarchaean evolution of the Isukasia area. Friend et al. (1996) showed that the Godth~bsfjord region is composed of Palaeo- to Neoarchaean terranes assembled at about 2720 Ma (Fig. 2.3-1). These terranes are: (1) the Akulleq terrane composed of the Palaeoto Mesoarchaean (3800-3600 Ma) granitoids and supracrustal rocks within the Itsaq Gneiss Complex, and Neoarchaean supracrustal and intrusive rocks; (2) the Akia terrane consisting mainly of Mesoarchaean (3220-3000 Ma) orthogneisses; and (3) the Tasiusarsuaq terrane including Neoarchaean (2920-2800 Ma) granitoids and older supracrustal rocks (Fig. 2.3-1). Neoarchaean intrusive rocks include variably deformed and metamorphosed granites, tonalities, gabbros, and anorthosites. The Ivis~.rtoq greenstone belt contains the largest, well-preserved Neoarchaean volcanic and sedimentary rocks in the region (Fig. 2.3-1). The Isua greenstone belt is about 35 km long and up to 4 km wide. The belt contains the oldest known, relatively well-preserved metavolcanic and metasedimentary rocks on Earth (Rosing et al., 1996; Appel et al., 1998; Fedo, 2000; Myers, 2001a, b). Recent detailed mapping suggests that the belt is composed of several fault-bounded litho-tectonic sequences, including basaltic and high-MgO basaltic pillow lavas, ultramafic intrusions, chert-banded iron-formation (BIF), and a minor component of clastic sedimentary rocks (Myers, 200 l a, b). The thickness and lithological characteristics of these supracrustal sequences vary along the belt. All lithologic units are variably metamorphosed, metasomatised, and deformed (Myers, 2001b). Geochemical studies suggest that the belt experienced several tectonothermal events in the Palaeoarchaean, Neoarchaean, and Proterozoic (Blichert-Toft and Frei, 2001 ; Frei et al., 2002; Polat et al., 2003), resulting in widespread modification of chemical and isotopic compositions. The most pervasive metasomatism occurs in the outer arc sequence of the belt (Fig. 2.3-2; Nutman, 1986; Rose et al., 1996). According to Rose et al. (1996), this calc-silicate-carbonate metasomatism is closely associated with early Archaean ultramafic intrusions.

78

Chapter 2: Generation of Continental Crust

On the basis of recent detailed mapping, Myers (2001 b) has recognised low- to highstrain litho-tectonic domains in the belt. This study also revealed a complex history of polyphase deformation and high grade metamorphism. The domains of least strain contain well-preserved volcanic and sedimentary features, including pillow basalts, pillow breccia, heterogeneous volcanic breccia, and polymictic conglomerates (Komiya et al., 1999; Myers, 2001b; Appel et al., 2001; see also section 2.2). The variably deformed pillow basalts are intercalated with ultramafic units. In contrast, primary volcanic and sedimentary structures in the highly strained domains are intensely deformed and metasomatised. The original stratigraphic relations between mafic (pillow basalts) and ultramafic units have been disrupted and complicated throughout the belt. Most supracrustal units have been recrystallised under upper greenschist to amphibolite facies metamorphism. Mafic dykes cut the early schistosity and lineation (Appel et al., 1998). There are a number of lines of evidence suggesting an intra-oceanic origin for the Isua volcanic rocks. These are: (1) stratigraphic association of chert layers, banded ironformations, and pillow basalts (Rosing et al., 1996; Myers, 2001b); (2) the absence of xenocrystic zircons in mafic-ultramafic volcanic rocks; (3) granitoids and felsic dykes in the region are younger than the volcanic rocks of the belt (Moorbath et al., 1977; Nutman and Bridgwater, 1986); and (4) conglomerates contain no continental detritus (e.g., granitic and gneissic pebbles; Fedo, 2000). Mineralogically, volcanic rocks from the central arc sequence are amphibolites, consisting predominantly of tremolite-actinolite-hornblende-chlorite-talc schist (Rosing et al., 1996; Gruau et al., 1996; Myers, 2001b). Myers (2001b) showed that these amphibolites were derived from pillow basalts. Volcanic rocks from the outer arc sequence are composed mainly of hornblende-garnet-biotite-dolomite-ankerite schist and amphibolites, whereas volcanic rocks of the inner arc sequence consist primarily of amphibolites and chlorite-talc schist (Rose et al., 1996; Gruau et al., 1996; Myers, 2001 b). Like those in the central arc sequence, amphibolites and chlorite schists in the outer and inner arc sequences were derived from pillow basalts (Myers, 2001b). Given the fact that all supracrustal rocks in the Isua greenstone belt have been metamorphosed, the prefix "meta" will be taken as implicit throughout the remainder of this section. Geochemistry Central tectonic unit

Detailed major and trace element characteristics of volcanic rocks of the central lithotectonic sequence have been discussed in Polat et al. (2002). Accordingly, only a summary of these compositional features is presented here (Table 2.3-1). The least altered volcanic rocks are characterised by high Mg-numbers (0.60-0.80), MgO (7-16 wt.%), A1203 (14-20 wt.%), Ni (60-645 ppm), and Cr (60-1920 ppm) contents, but low TiO2 (0.20-0.40 wt.%), Zr (12-30 ppm), Y (6-14 ppm), and rare earth elements (REE) concentrations (Table 2.3-1). AlzO3/TiO2 (45-94) ratios are super-chondritic whereas Zr/Y (1.3-2.5) ratios tend to be sub-chondritic. Collectively, these compositional features represent a coherent mafic to ultramafic suite.

Table 2.3-1. Summary of the ranges of the significant compositional and element ratios for mafic to ultramafic volcanic rocks in the > 3.7 Ga Isua greenstone belt*

Si02 (wt.%) MgO Ti02 A1203 Fe203 Mg# Cr ( P P ~ ) Ni Zr Nb Th La Y Yb (La/Sm)cn (LdYb)cn (Gd/Yb)cn ZrN (ZrISm),, A1203/Ti02 (NbILa),, (ThILa),, (Nb/Th)p

Outer arc sequence least altered 48-53 8.9-19.6 0.50-1.05 6.5-1 1.9 12.0-17.7 0.58-0.77 1 94-29 16 88-952 34-7 1 1.46-2.46 0.25-2.12 1.80-3.95 12.3-21.7 1.042.01 0.7-1.1 1.2-2.1 1.4- 1.7 2.6-3.3 0.7-1.1 1 1.3-1 3.5 0.37-0.82 0.78-5.81 0.14-0.80

Outer arc sequence variably altered 45-6 1 4.8-19.4 0.4W.84 5.9-15.8 9.4-15.4 0.440.78 144-3 170 42-1 157 29-64 0.75-2.16 0.1 1-2.24 0.82-9.77 9.421.2 0.88-2.02 0.41.7 0.40-4.3 0.7-1.9 2.M.6 0.67-1.6 11.4-27.8 0.17-0.97 0.36-6.88 0.12-1.70

Central arc sequence least altered 47-54 6.8-16.1 0.17-0.40 13.9-20.2 8.2-1 1.9 0.61-0.77 60- 1920 60-645 12.1-29.5 0.13-0.80 0.04-0.29 0.31-1.83 6.0-13.7 0.94-1.84 0.56-1.39 0.16-0.79 0.2W.61 1.3-2.5 1.1-1.6 45-94 0.23-0.77 0.57-2.07 0.3 1-0.76

Central arc sequence variably altered 40-54 11.7-24.5 0.15-0.26 12.7-19.2 7.9-14.5 0.77-0.85 277-3452 163-863 7.5-19.1 0.02-0.35 0.01-0.06 0.1W . 6 2 5.4-1 1.O 0.84-1.86 0.30-1.0 0.10-0.35 0.38-0.47 1.2-2.2 1.1-2.0 55-80 0.08-2.57 0.28-1.29 0.1 1-3.34

Inner arc sequence least altered 48-53 4.5-21 .O 0.52-1.14 7.9-14.1 12.2-15.0 0.40.77 33-2268 3 1-826 46-77 1.21-2.75 0.40.79 2.44.3 10.9-27.6 1.1-2.7 1.0-1.7 1.3-3.0 1.4-1.9 2.8-5.1 0.8-1.4 12.5-15.0 0.29-0.60 0.8-1.9 0.22-0.57

Inner arc sequence variably altered 51-53 4.3-21.9 0.32-1.18 5.5-14.6 11.3-15.0 0.39-0.81 3 1-2040 31-1 100 25-8 1 0.67-2.83 0.27-0.62 0.4W.52 8.9-27.5 0.90-2.7 0.3-0.8 0.3-1.2 1.1-1.4 2.8-3.1 0.91-1.32 12.5-16.9 0.6-1.7 1.O-5.6 0.3-0.6

*Data for the central arc sequence from Polat et al. (2002). and data for the inner and outer arc sequences from unpublished data of Polat and Hofmann.

Q

2

2

S

",

B

9 ?

z

p 2. "u 3 Z

80

Chapter 2: Generation of Continental Crust

Chondrite-normalised REE patterns are concave upwards (La/Smcn = 0.56-1.39; Gd/Ybcn = 0.26-0.61). On primitive mantle-normalised trace element diagrams, they are characterised by relative depletion of Nb (Nb/Thpm ----0.31-0.76; Nb/Lapm -- 0.23-0.77), but enrichment of Zr (Zr/Smpm -- 1.1-1.6), relative to neighbouring REE (Table 2.3-1). Inner and outer arcs

Volcanic rocks of the inner and outer arc sequences are chemically similar (Table 2.3-1); therefore, their geochemical features are described together. These are compositionally variable at 4-22 wt.% MgO, 31-1157 ppm Ni, 31-3170 ppm Cr, and Mg-numbers of 0.39 to 0.81 (Table 2.3-1). These compositional features are consistent with mafic to ultramafic compositions. They possess variable SiO2 (45-61 wt.%), TiO2 (0.3-1.2 wt.%), and A1203 (6-16 wt.%) abundances (Table 2.3-1). AlzO3/TiO2 (11-28) ratios range from sub-chondritic to super-chondritic, and Zr/Y (2.0-5.1) ratios are mostly super-chondritic. In addition, they have the following geochemical features: (1) depleted to enriched REE ((La/Sm)cn = 0.3-1.7; (Gd/Yb)cn = 0.7-1.9)patterns; (2) low (Nb/Th)pm (0.12-1.70)and (Nb/La)pm (0.17-1.7) ratios, generating negative Nb anomalies (Table 2.3-1). Discussion Element mobility

Element mobility is a major problem for studying the > 3.7 Ga Isua volcanic rocks, which have undergone sea floor hydrothermal alteration, greenschist to amphibolite facies metamorphism, metasomatism, and polyphase deformation destroying primary textures and minerals (Gruau et al., 1996; Frei and Rosing, 2001; Frei et al., 2002; Polat et al., 2002; 2003). It is important therefore to understand and take account of the effects of alteration on the geochemistry of the Isua volcanic rocks before considering any petrogenetic interpretation. Accordingly, in this section we briefly discuss the possible effects of alteration on the geochemical composition of the Isua volcanic rocks in an attempt to assess the least altered samples for near-primary geochemical signatures, particularly those for HFSE and REE. The effects of metamorphic alteration on the central arc volcanic sequence were evaluated by Polat et al. (2002). Therefore, in this contribution we evaluate the effects of alteration on the volcanic rocks from only the inner and outer arc sequences. Following Polat et al. (2002), several criteria were adopted to assess the effects of alteration on the inner and outer arc volcanic rocks. These are: (1) the magnitude of correlations with the least mobile element, Zr, on binary diagrams; elements having a correlation coefficient (R) < 0.75 were considered as mobile; (2) the presence of significant Ce anomalies on primitive mantle-normalised diagrams; samples possessing Ce/Ce* (asterisk denotes the concentration interpolated from that of adjacent elements on the primitive mantlenormalised diagram) ratios greater than 1.1 and less than 0.9 were designated as variably altered; and (3) the existence of significant carbonate or silica alteration ( > 2 wt.%). Given the fact that all Isua volcanic rocks have been altered to some extent, samples were divided into least altered and variably altered groups (Table 2.3-1). Only the former group was used for petrogenetic interpretation.

2.3. Isua Greenstone Belt: Geodynamic Processes

81

Fig. 2.3-3. (a-d) Zr versus selected element variation diagrams to highlight the effects of alteration on Si, Na, K, and Sr. Most altered samples are located in the outer arc sequence, where carbonate alteration is the most pervasive. (e-h) Zr versus selected element variation diagrams to highlight the limited effects of alteration on Ti, Nb, and REE. Strong correlations indicate that these elements were not disturbed significantly by post-emplacement alteration.

Chapter 2: Generation of Continental Crust

82

Outer arc sequence 100

(a) O 462901 Minor alteration 9462915 Minor alteration [] 462912 Strong alteration <> 462916 Strong alteration

3E 13.. r

10

o iv,

1 II

I I I I I 1 1 1 I I I I I I 1 I I Th Nb La Ce Pr Nd Zr Sm Eu Ti Gd Tb Dy Y Ho Er Trn Yb

Inner arc sequence 100 (b)

O 2000-27 Minor alteration 9 2000-15 Strong alteration

3E a.

"~ 10 o

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Th Nb La Ce Pr Nd Zr Sm Eu Ti Gd Tb Dy Y Ho Er Tm Yb

Fig. 2.3-4. Primitive mantle-normalised trace element diagrams of variably altered samples to highlight the effects post-emplacement carbonate alteration. Primitive mantle normalisation values from Hofmann (1988). The mobility of Si, Na, Ca, Rb, K, Sr, Ba, Ca, Fe, P, and Pb in volcanic rocks of the central litho-tectonic sequence is well documented (Polat et al., 2002). There are extreme variations in these elements in the inner and outer arc volcanic rocks (e.g., Rb = 0.2-98 ppm; Ba = 1..1-258 ppm; K20 = 0.04-3.92 wt.%; Na20 = 0.01-4.25 wt.%), which do not correlate well with Zr abundance, and show very weak correlations with MgO content (Fig. 2.3-3). This variability is reflected in significant enrichments or depletions on primitive mantle-normalised diagrams (not shown). Collectively, the large concentration

2.3. lsua Greenstone Belt: Geodynamic Processes

83

range of these elements at a given Zr value signifies loss or gain of these elements during post-emplacement hydrothermal alteration, metamorphism, or metasomatism. Accordingly, these elements designated as mobile were screened out, and were not used for petrogenetic interpretation. In contrast to these mobile elements, a number of studies (Arndt, 1994; and references therein) have found that in many Archaean volcanic rocks the effects of alteration on REE, high field-strength elements (HFSE), A1, Cr, Ni are minor. Thus, these latter elements are widely considered to be immobile. However, some studies suggest that these elements can also be mobile during intense carbonate or silica alteration, and metamorphism, resulting in non-coherent patterns on normalised trace element diagrams (Kerrich and Fryer, 1979; Arndt et al., 1989; Gruau et al., 1992; Lahaye et al., 1995). Therefore, we have assessed the effect of alteration on these elements in the severely altered Isua volcanic rocks. On diagrams of Zr versus Nb, Nd, Sm, Ti, Y, and heavy rare earth elements (HREE) most samples display systematic correlations, consistent with the relatively low mobility of these elements (Fig. 2.3-3). In contrast, there is a large scatter of Th and light rare earth elements (LREE) (La, Ce), particularly in the outer arc volcanic rocks, consistent with mobility of these elements during post-magmatic alteration processes. The effects of carbonate alteration on the mobility of these elements are well illustrated by the primitive mantle-normalised diagrams where samples with minor (< 2 wt.%) carbonate alteration have coherent REE and HFSE (Nb, Zr, Ti, Y) patterns, consistent with the limited mobility of these elements (Fig. 2.3-4). However, samples with moderate to strong (2-20 wt.%) carbonate alteration display depleted LREE patterns and minor positive Zr and Nb anomalies, suggesting the loss of LREE, and that Nb and Zr were less mobile than REE (Fig. 2.3-4). The behaviour of Th is rather variable. Frei et al. (2002) suggested that the Isua volcanic rocks gained LREE during post-magmatic alteration, but the results of this study indicate that LREE were lost during intense carbonate alteration. The loss of LREE is further endorsed by Sm-Nd isotope systematics of altered samples (Albar~de et al., 2000; Polat et al., 2003). However, our study supports the conclusions of Frei et al. (2002) on the gain of Th and LILE (large-ion lithophile elements; K, Rb, Ba, etc.) in many samples of the outer arc volcanic sequences. Samples which gained significant amounts of silica (4-10 wt.%) in the outer arc sequence are enriched in A1203 and Na20, but depleted in CaO and LREE. It appears that variably altered samples are closely associated with high-strain domains, whereas the least altered samples tend to occur in low-strain domains (Myers, 2001b). In addition, the variably altered samples tend to occur in the vicinity of ultramafic intrusions, which may have played an important role in the generation of some metasomatic fluids in the outer arc volcanic sequence (Rose et al., 1996; Rosing et al., 1996). Crustal contamination Negative Nb and Ti anomalies in the Isua volcanic rocks could possibly reflect some crustal contamination (Fig. 2.3-5). However, SiO2, MgO, Ni, Cr, Co, Th, Ti, and LREE contents in these rocks do not correlate with the magnitude of negative Nb and Ti anomalies. High AlzO3/TiO2 (45-90) ratios and positive Zr (Zr/Zr* = 1.1-2.1) anomalies in the volcanic rocks of the central tectonic sequence could not have resulted from crustal contamination,

84

Chapter 2: Generation of Continental Crust

given that the continental crust has lower values (18 and 0.94, respectively) of these ratios. There is no evidence from the published initial Sr, Nd, Hf, and Pb isotope composition for any crustal contamination by significantly older rocks in the Isua greenstone belt (see Albar~de et al., 2000; Kamber et al., 2001). In conclusion, negative Nb and Ti anomalies and other geochemical features of the least altered Isua volcanic rocks appear to reflect the mantle source characteristics and petrogenetic processes in the Palaeoarchaean, rather than continental contamination.

Petrogenetic interpretation As shown above, alteration, deformation, and crustal assimilation can all be ruled out as the cause of the specific geochemical characteristics of the Isua volcanic rocks. The high MgO, Ni and Cr contents indicate high temperature partial melting. Continuous compositional range (MgO = 7-16 wt.%, Ni = 60-645 ppm in central tectonic unit; MgO = 5-21 wt.%, Ni = 31-952 ppm in outer and inner arc units) is consistent with olivine-controlled fractionation processes. Negative Nb and Ti anomalies on primitive mantle-normalised diagrams of the Isua volcanic rocks are consistent with a subduction zone petrogenetic origin (Fig. 2.3-5; cf. Pearce and Peate, 1995). The geochemical characteristics of the least altered volcanic rocks are consistent with the presence of two geochemically distinct volcanic associations in the Isua greenstone belt. These are (1) a low-HFSE association (e.g., TiO2 = 0.20-0.40 wt.%; Zr = 12-30 ppm; Nb = 0.13-0.80 ppm; Y = 6-14 ppm) in the central arc litho-tectonic sequence, and (2) a high-HFSE association (TiO2 = 0.50-1.14 wt.%; Zr = 34-77 ppm; Nb = 1.2-2.7 ppm; Y = 11-28 ppm) in the inner and outer arc litho-tectonic sequences (Fig. 2.3-2). These two suites have distinct trace element patterns on primitive mantlenormalised diagrams (Fig. 2.3-5). In addition, volcanic rocks from the central arc sequence have consistently lower Zr/Y and (Gd/Yb)pm ratios than those from the inner and outer arc sequences at a given Mg-number value, and MgO and Ni contents (Fig. 2.3-6; Table 2.3-1). Collectively, geochemical differences between the two suites cannot be attributed to metamorphic alteration, fractional crystallisation, degree of partial melting, or crustal contamination processes but could be explained by two geochemically-distinct mantle sources. Primitive mantle-normalised REE patterns suggest that the low-HFSE association was derived from an LREE depleted mantle source, whereas the high-HFSE association was derived from a fiat to LREE enriched mantle source. The depletion of the source of the lowHFSE association may have resulted from a previous melt extraction event(s). The LREE enriched characteristic of the high-HFSE suite can be attributed to the enrichment of their source by subduction zone metasomatism (cf. Pearce and Peate, 1995). Super-chondritic Zr/Sm ratios in the low-HFSE association may reflect the metasomatism of their sources by slab-derived melts (cf. Pearce et al., 1999). Positively fractionated HREE patterns in the high-HFSE association (Fig. 2.3-5) are consistent with melt generation in the field of garnet stability (cf. Sun and Nesbitt, 1978), whereas negatively fractionated HREE patterns in the low-HFSE association suggest that garnet was not a stable phase in the source. The low-HFSE association was referred to as "boninitic" by Polat et al. (2002) because it has many geochemical characteristics similar to those of Tertiary boninites from the west-

2.3. Isua Greenstone Belt: Geodynamic Processes

100 (a)

==

Central

85

9N-MORB

sequence

arc

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0.1 100

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(b)

Inner arc sequence

10 3E

13. t~ 0

iv.

1

0.1

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100 (c)

10 3E

13.

Outer

arc

sequence

!

t~ 0

n,

1

0.1

i

i

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J

I

J

I

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J

I

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i

i

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Th Nb La Ce Pr Nd Zr Sm Eu Ti Gd Tb Dy Y Ho Er Tm YI

Fig. 2.3-5. Primitive mantle-normalised trace element patterns for the least altered samples from different lithotectonic sequences. Primitive mantle normalisation values from Hofmann (1988).

86

Chapter 2: Generation of Continental Crust

Fig. 2.3-6. (a--c) Ti versus Nb and Zr, and Zr/Y versus Gd/Yb variation diagrams, indicating that samples from the central arc sequence (low-HFSE association) and from the inner and outer arc sequences (high-HFSE association) plot separately, consistent with the presence of two geochemically distinct mafic-ultramafic rock associations in the Palaeoarchaean Isua greenstone belt. These two volcanic associations are likely to reflect distinct tectonic segments of a Palaeoarchaean intra-oceanic arc(s).

2.3. Isua Greenstone Belt: Geodynamic Processes

87

ern Pacific Ocean. The primary melts of the boninitic association require high temperature (c. 1300~ low pressure (< 10 kbar) melting of an extremely depleted clinopyroxenepoor harzburgitic mantle source (cf. Crawford et al., 1989; Taylor et al., 1994). According to Polat et al. (2002), the melts of this association resulted from an interaction between an extremely depleted subarc mantle source and slab-derived adakitic melts. The volcanic rocks of the high-HFSE association can be defined as "picrites" given that they have geochemical features similar to those of Phanerozoic picrites (cf. Ramsay et al., 1984; Eggins, 1993; Kamenetsky et al., 1995). Generation of the primary melts of the picritic association requires high temperature (c. 1300~ and high pressure (c. 30 kbar) melting of refractory peridotitic mantle source (cf. Eggins, 1993). In conclusion, the low-HFSE association was derived from a shallower and more depleted mantle source than the high-HFSE association. Conclusions

On the basis of the least mobile elements in samples screened for minimum alteration, two distinct types of mafic to ultramafic volcanic associations have been recognised in structurally separated sequences of the > 3.7 Ga Isua greenstone belt. These are (1) a lowHFSE association in the central litho-tectonic sequence, and (2) a high-HFSE association in the outer and inner arc litho-tectonic sequences. The former association was referred to as "boninitic" by Polat et al. (2002) because it has many geochemical characteristics similar to those of Tertiary boninites from the western Pacific Ocean. The latter association has geochemical features similar to those of Phanerozoic island arc picrites (cf. Ramsay et al., 1984; Eggins, 1993; Kamenetsky et al., 1995). Volcanic rocks with boninitic affinity have recently been reported from the Neoarchaean Abitibi and Frotet-Evans greenstone belts of the Superior Province, suggesting that boninitic volcanism in the Archaean may have been more widespread than is currently recognised (Kerrich et al., 1998; Boliy and Dion, 2002). Studies of Tertiary boninitic rocks suggest that they are the products of high temperature, low pressure partial melting of a hydrous, refractory mantle source above a subducted oceanic lithosphere at intra-oceanic convergent plate boundaries, such as the Izu-BoninMariana subduction zones of the western Pacific (Taylor et al., 1994). Similarly, Phanerozoic island arc picrites are the products of high temperature, high pressure melting of refractory peridotitic subarc mantle sources (Eggins, 1993). If the geochemical characteristics of the Palaeoarchaean lsua boninites and picrites have the same geodynamic significance as their Phanerozoic counterparts, then they likely originated in an intraoceanic subduction zone-like tectonic setting, suggesting that Phanerozoic-like plate tectonic processes were operating as early as 3.8 Ga (Fig. 2.3-7). On the basis of combined structural and geochronological studies, Nutman et al. (2002) suggested that the c. 3800 Ma outer arc sequence and the c. 3700 Ma central arc sequence were probably juxtaposed by Phanerozoic-like plate tectonic processes operating in the Palaeoarchaean. The geochemical characteristics of the low-HFSE and high-HFSE associations are comparable to those in the Tertiary Mariana forearc, and in the Vanuatu and Solomon island arcs, respectively (Ramsay et al., 1984; Eggins, 1993; Taylor et al., 1994). The occurrence of boninites and picrites have been described from certain Phanerozoic subduction zones featuring high

Chapter 2: Generation o f Continental Crust

88

Fig. 2.3-7. Interpreted geodynamic settings for the Isua boninites and picrites. Given the fact that the original temporal and spatial relationships between the boninites and picrites are unknown, it is not necessary to assume that they erupted above the same subduction zone. It is equally possible that they erupted in different oceanic arcs and were juxtaposed by subsequent tectonic processes. geothermal gradients, suggesting that high geothermal gradients in Palaeoarchaean subduction zones may have played an important role in the production of the Isua boninites and picrites. It should be emphasised that analogies with modem geodynamic settings should be used with great caution as templates for geodynamic interpretation, but such comparisons may help in understanding the processes by which Archaean greenstone belts originated. Similarly, it should also be emphasised that the Isua greenstone belt represents only a tiny fraction of the surviving Palaeoarchaean crust; therefore, the geochemical characteristics of the Isua volcanic rocks may not offer a complete picture of the petrogenetic and geodynamic processes operating in the Palaeoarchaean.

2.4.

ABITIBI GREENSTONE BELT PLATE TECTONICS: THE DIACHRONONOUS HISTORY OF ARC DEVELOPMENT, ACCRETION AND COLLISION

R. DAIGNEAULT, W.U. MUELLER AND E.H. CHOWN Introduction

Comparing the tectonic evolution of Archaean greenstone belts with Phanerozoic counterparts has long been a contentious issue (e.g., Hamilton, 1998), yet thrust structures, inferred from inverted sequences, recumbent folds and subhorizontal high strain zones, have been identified on numerous Archaean cratons (Stowe, 1974, 1984; Poulsen et al., 1980; de Wit, 1982; Jirsa et al., 1992; Mueller et al., 1996; Kusky and Polat, 1999) (see also discussion in section 3.6). Similarly, shallow-dipping seismic reflections, interpreted as a remnant subduction zone, have been recognised at depth, north of the Abitibi greenstone belt (Calvert et al., 1995). The combination of outcrop- and crustal-scale shallow dipping The Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mucller and O. Catuneanu

2.4. Abitibi Greenstone Belt Plate Tectonics

89

structures and thrusts is consistent with shortening via horizontal plate motions generated by subduction. However, the amount of tectonic transport is difficult to assess and defining allochthonous or autochthonous terranes remains a problem (Thurston, 2002). Large-scale processes become complex if plume-generated komatiites are involved. Mantle plumes (sections 3.2 and 3.3) impinging on Archaean subduction zones are a particularly appealing geodynamic process used to explain komatiites associated with arc volcanic rocks (Dostal and Mueller, 1997; Hollings et al., 1999; Wyman, 1999a; Polat and Kerrich, 2000a). Plate tectonics are accepted as the principal mechanism for Archaean greenstone belt formation and deformation (Langford and Morin, 1976; De Wit, 1998) (see, however, section 3.6), but plume activity for komatiite-tholeiitic basalt sequences remains an important process (Tomlinson and Condie, 2001). The resurgence of the "sagduction" concept for greenstone belts of the Dharwar Craton (Chardon et al., 1998), shows that vertical tectonics are important, but may be explained readily by oblique convergence (Chadwick et al., 2000). Kloppenburg et al. (2001) explained gravitational forces in the 3.5-3.4 Ga Warrawoona Group and associated granitoid complexes of the Pilbara Craton as the product of an extensional core complex deformational event linked to plate tectonics. Whilst komatiites are considered plume derivatives, their structural signature imposed on Archaean arc and ocean floor terranes remains elusive. Physical features of plumes such as radiating dyke swarms, selective changes in sedimentary thickness, or domal upwelling are excellent criteria in continental break-up settings but are unlikely to be recognised in the arc and ocean plateau environment. In the arc setting with plume impingement, subduction-generated structures are therefore dominant. Focus is placed here on the structural evolution of the 2735-2670 Ma Abitibi greenstone belt, the largest coherent greenstone belt in the world (Card, 1990). The 300 x 700 km Abitibi greenstone belt is a prime example of subduction-dominated processes even though mantle plumes constantly affected Abitibi evolution. The Abitibi greenstone belt is a linear east-trending volcano-sedimentary sequence pierced by plutonic suites, that displays arc formation, arc evolution, arc-arc collision and arc fragmentation (Mueller et al., 1996) and is therefore strikingly similar to modern collisional orogenies. It is the best belt in the world to decipher Archaean arc time-space variations and the deformation geometry produced by Archaean plate tectonics (Daigneault et al., 2002) because of precise U-Pb zircon age determinations (Mortensen, 1993a, b; Ayer et al., 2002), volcanic facies analysis (Dimroth et al., 1982, 1985), sedimentary facies analyses (Mueller and Donaldson, 1992a) as well as a well constrained emplacement history of plutons (Chown et al., 1992, 2002) and detailed geochemistry (Goodwin, 1982). Divisions of the Abitibi Greenstone Belt The Abitibi greenstone belt (Fig. 2.4-1), divided into Southern (SVZ) and Northern Volcanic Zones (NVZ; Chown et al., 1992) represents a collage of two arcs, delineated by the Destor-Porcupine Manneville Fault Zone (DPMFZ; Mueller et al., 1996). The SVZ is separated from the Pontiac sedimentary rocks, an accretionary prism (Calvert and Ludden, 1999) to the south, by the Cadillac-Larder Lake Fault Zone (CLLFZ). The fault zones are

~ ~ l t r a r n a f i c - m a fvolcanic ic r a d t s Mafc ~ lntrus~on

I

Fig. 2.4-1. Division of the Abitibi greenstone belt into southern (SVZ) and northern volcanic zones (NVZ) with external and internal segments in the NVZ. BRS = Black River segment; MS = Malartic segment. Plutons cited in text: Cb = Chibougamau; Fv = Flavrien; Fr = Franquet; La = Lacorne; Lm = Lamotte; Lp = Lapparent; Ma = Marest; Mi = Mistaouac; Mu = Muscocho; Op = Opemisca, Pr = Preissac; Re = Renaud; Wa = Waswanipi. DLC = Dore Lake Complex, BRC = Bell River Complex. Sedimentary basins: Ch = Chicobi, Tb = Taibi, K = Kewagama, Po = Pontiac, Du = Duparquet, Cs = Caste, Om = Opemisca, Ca = Caopatina. Pie diagrams show distribution of prominent rock types. Modified from Chown et al. (1992) and Daigneault et al. (2002).

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91

terrane zippers that show the change from thrusting to transcurrent motion as documented in the turbiditic flysch basins overlain unconformably by, or in structural contact with, coarse clastic deposits in strike-slip basins (Mueller et al., 1991, 1994a, 1996; Daigneault et al., 2002). A further subdivision of the NVZ into external and internal segments is warranted, and based on distinct structural patterns with the intra-arc Chicobi sedimentary sequence (Fig. 2.4-1) representing the line of demarcation. Interestingly, Dimroth et al. (1982, 1983) recognised this difference and used it to define internal and external zones of the Abitibi greenstone belt. Subsequently, numerous alternative Abitibi divisions were proposed (see Chown et al., 1992), but all models revolved around a plate tectonic theme. The identification of a remnant, Archaean, north-dipping subduction zone by Calvert et al. (1995) corroborated these early studies. The 2735-2705 Ma NVZ is ten times larger than 2715-2697 Ma SVZ and both granitoid bodies and layered complexes are abundant in the former. In contrast, plume-generated komatiites, a distinct feature of the SVZ, are only a minor component the NVZ, observed only in the Cartwright Hills and Lake Abitibi area. Komatiites rarely constitute more than 5% of greenstone sequences and the Abitibi is no exception (Sproule et al., 2002). The linear sedimentary basins are significant in the tectonic history because they link arcs and best chronicle the structural evolution and tempo of Archaean accretionary processes. The NVZ is composed of volcanic cycles 1 and 2, which are synchronous with sedimentary cycles 1 and 2, whereas the SVZ exhibits volcanic cycles 2 and 3, with sedimentary cycles 3 and 4 (Mueller et al., 1989; Chown et al., 1992; Mueller and Donaldson, 1992a; Mueller et al., 1996). The chronology of deformational events within the two zones and during Abitibi evolution is striking. The D l - D 4 events in the NVZ and D l - D 8 events in the SVZ overlap and show a deformation continuum from the north to south (Mueller et al., 1996). These events are the result of far-field motions associated with oblique plate convergence (Table 2.4-1). Northern volcanic zone

The 2735-2720 Ma volcanic cycle 1 in the NVZ is explained as an extensive subaqueous 3-6 km thick mafic basalt plain upon which massive sulphide-bearing, 0.2-5 km-thick central volcanic edifices developed (Chown et al., 1992). Notable examples of NVZ felsic centres include the c. 2728 Ma Joutel volcanic complex (Legault et al., 2002), the c. 2728 Ma Normetal volcanic complex (Lafrance et al., 2000), the c. 2725 Ma Matagami complex (Pich6 et al., 1993), and the c. 2730 Ma Hunter Mine caldera complex (Mueller and Mortensen, 2002) (section 4.6). Time-equivalent intra-arc flysch basins, referred to as sedimentary cycle 1, are interstratified with or overlie volcanic cycle 1 rocks (Chown et al., 1992) This volcano-sedimentary event represents the incipient arc-forming phase. The basins, forming east-trending units over 100 km in length, are bounded by layerparallel faults. Orogen-parallel faulting has been identified in the Abitibi greenstone belt (Daigneault and Archambault, 1990; Lacroix and Sawyer, 1995) but also in the Opatica belt immediately to the north (Benn and Sawyer, 1992; Sawyer and Benn, 1993). Volcanic cycle 2 (2720-2705 Ma) corresponds to arc emergence, with 3-5 km thick mafic-felsic volcanic sequences, and is best documented in the Chibougamau area. The

Table 2.4-1. Deformational events in the northern (NVZ) and southern volcanic zones (SVZ), Abitibi greenstone belt Region

Event

NVZ

DI D2

D3 D4

svz

N-S horizontal shortening

Culmination of N-S shortening Dextral shearing

N-S horizontal shortening Thrusting

Dextral shearing Thrusting

Dextral shearing Renewed thrusting Extensional movement

Characteristics

Ages constraints (Ma; selected examples)

Early folds without schistosity E-trending folds and regional schistosity

< 2716 Chibougamau pluton (Krogh, 1982) Plutons intruding D2 folds and regional schistosity: - 2701 Muscocho pluton (Mortensen, 1993a) - 2700 Renaud pluton (Mortensen, 1993a) - 2697 Opemisca pluton (Frarey and Krogh, 1986)

h)

E-trending faults with reverse movements Dextral SE-trending faults and late dextral shearing along E-trending fault

Plutons emplaced during dextral transpression: 2696 Colombourg pluton (Mortensen, 1993b) - 2692 Franquet pluton (Frarey and Krogh, 1986) - 2695 Waswanipi pluton (Davis et al., 2000) -

E-trending folds and regional schistosity NVZ-SVZ accretion, thrusting along the DPMFZ Caste accretionary prism Formation of the Duparquet pull-apart basin along the DPMFZ SVZ-Pontiac accretion Pontiac accretionary prism

Formation of the Granada pull-apart basin along the CLLFZ Fold and thrust within the clastic Granada (Timiskaming) and Pontiac flysch deposits Exhumation of Pontiac and Malartic segment along CLLFZ and DPMFZ

2695 Early phases of the Lacorne pluton (Steiger and Wasserburg, 1969) - 2694 Caste sedimentary rocks (Davis, 2002) - 2681 Porphyry stock, Duparquet (Mueller et al., 1996) - 2689 Porphyry stock, Duparquet (Mueller et al., 1996) - 2685 Pontiac sedimentary rocks (Davis, 2002) - 2682 Lac Fournikre pluton cross-cutting Pontiac sedimentary rocks - 2673 Granada felsic volcaniclastic rock (Davis, 2002) - 2672 Porphyry intruding Granada basin (Davis, 2002) -

Dextral movement along the Cadillac Fault (sensu stricro) NE-trending Z-folds and Pressure-solution NE cleavage

2

4

B b

g 9

- 2660 Preissac pluton ( D u c h m e et al., 1997) -

Dextral shearing

w

2647 Lamotte pluton ( D u c h m e et al., 1997) 2643 Lamone pluton (Machado et al., 1991)

2

J

v

2.4. Abitibi Greenstone Belt Plate Tectonics

93

mafic sequence is considered a broad subaqueous shield volcano surmounted by subaqueous to emergent felsic centres cored by synvolcanic plutons. The volcanoes were eroded down to the plutonic roots (e.g., Chibougamau pluton, 2718 Ma; Krogh, 1982), and the diverse sedimentary cycle 2, reflecting fluvial, shallow marine and deep-water settings, contains abundant plutonic clasts in the conglomerates that attest to the erosion of synvolcanic plutons (Mueller and Dimroth, 1987). Shoshonite volcanism (Picard and Piboule, 1986a; Dostal and Mueller, 1992) is associated with this stage and supports the inference of evolved arc development.

Structural styles and timing of deformation in the NVZ The NVZ shows heterogeneous deformation with alternating domains of high and low strain. Low strain zones are characterised by a distinct fold pattern, whereas high strain zones are found generally at the interface between different rock assemblages, forming regional fault zones or forming contact strain aureoles around synvolcanic plutons. Preregional deformation (D1) with km-scale north-trending F I folds was deduced from inversion of structural facings (Daigneault et al., 1990; Chown et al., 1992). Although a pervasive schistosity is absent, these locally prominent folds are considered either uplift or subsidence effects between early plutons or synvolcanic fault-related phenomena (Daigneault et al., 1990; Chown et al., 1992). The timing of this event is poorly constrained but must have occurred prior to 2710 Ma. The principal deformational event (D2) in the Abitibi greenstone belt is characterised by a pervasive schistosity due to extensive north-south shortening that developed after 2710 Ma. The internal segment features a tight upright regional east-trending fold pattern and a steeply dipping schistosity (Fig. 2.4-2). Axial traces of major F2 synclines occur within cycle 1 and 2 sedimentary basins, but are also recognised in selected mafic basalt sequences. These folds display a prominent "slaty" axial planar schistosity (Sp; Fig. 2.4-2). In contrast, the four major F2 anticlines are domal structures traced along plutons. The volcanic rocks between the domal anticlines exhibit complex deformation patterns with local north-south folds (D1) or foliations wrapping around the plutons that form triple junction foliation trajectories (Fig. 2.4-2). The external segment of the NVZ is defined by the first appearance of the shallow north-dipping thrust in the linear Chicobi sedimentary sequence (Lacroix and Sawyer, 1995). The different structural pattern is attributed to the absence of large granitoid plutons forming domal anticlines. The external segment displays a series of upright synclines and anticlines that progressively become overturned in the southern portion near the contact with the SVZ (Fig. 2.4-2). Mesoscopic F2 folds are generally isoclinal with steeply plunging axes. The principal east-trending schistosity, axial planar to both the mesoscopic folds and the regional synclines, is coeval with the dominant planar fabric in the internal segment of the NVZ. Layer-parallel faults and shear zones (D3; 2705-2698 Ma), a common feature of modern orogenies, represent east-trending pan-Abitibi discontinuities (Fig. 2.4-2) that commonly occur at the interface between rocks of different mechanical behaviour. For example, the Chapais syncline in the Chibougamau region (Fig. 2.4-2) shows the layer-parallel Kapunapotagen fault separating south-facing sedimentary cycle 2 rocks from north-facing

7-

---.

m m

Deformationzones Dunlnant d~pparallelstretch~ngItneahon Dantnant slnke-41pstretch~ngllneatlon Norlheasl-trend14fault Northern boundary delcfmahcmzone

Fig. 2.4-2. Structural map of the Abitibi greenstone belt. The internal NVZ is characterised by four major domal anticlines and the folds become south vergent near the DPMFZ in the external NVZ, and near the CLLFZ in the SVZ and Pontiac Group flysch deposits. Folds cited in text: WS = Waconichi syncline, CS = Chibougarnau syncline, CA = Chibougamau anticline, ChS = Chapais syncline, LDA = La Dauversikre anticline, DS = Druillette syncline, MMA = Mistaouac-Marais anticline, BA = Bernetz anticline, US = Urban syncline. Major deformation and fault zones cited in text: Ca = Cameron; CB = Casa-BCrardi; Ch = Chicobi; CLLFZ = Cadillac-Larder Lake Fault Zone; Fb = Faribault; GI = Gwillim; Kp = Kapunapotagen; La = Larnarck; LS = Lac Sauvage; Mc = Macamic; Nr = NormCtal; DPMFZ = Destor-Porcupine-Manneville Fault Zone. See Figure 2.4-1 for explanation of patterns.

9

-n 9 S

$

2.4. Abitibi Greenstone Belt Plate Tectonics

95

volcanic cycle 1 rocks. Truncation of a limb of a syncline is a common Abitibi feature. The D3 faults, prominent 1--4 km thick shear zones, exhibit a subvertical mylonitic foliation with a dominant dip-parallel stretching lineation. Movement along layer-parallel faults is difficult to discern, but can be deduced from contrasting lithological associations, missing lithological units and different structural styles on either side of the fault. These faults cross-cut the regional schistosity, truncate F2 fold hinges (Daigneault et al., 1990), and act as d~collement surfaces between blocks with different schistosity trends. However, some deformation zones display no evidence of non-coaxial flow, such as shear-sense indicators. On the contrary, the deformation pattern around objects or inclusions such as pressure shadows, displays a remarkable symmetry that can be interpreted as a dominant component of coaxial flow. This component can be interpreted as strain concentration between contrasting units and may represent only the final deformation increment of an earlier non-coaxial history. Fault movement is not systematic with both south-over-north (Kapunapotagen and Faribault faults; Daigneault et al., 1990) and north-over-south movement (Normetal fault; Lafrance, 2003; Fig. 2.4-2) being present. Late dextral shearing in the NVZ (D4; 2702-2692 Ma) best documents the signature of oblique convergence. Dextral shearing along southeast-trending and layer-parallel east-trending fault zones (Fig. 2.4-2) are the consequence of northwest-southeast shortening. The southeast-trending fault zones, more than 100 km long, cross-cut east-trending faults, regional folds and the pervasive S2-schistosity. Associated, 1-5 km wide deformation zones display a strong mylonitic fabric with subhorizontal stretching lineations and dextral shear-sense indicators. Offsets of up to 5 km are indicated. Subhorizontal stretching lineations and dextral shearing are also observed in the east-trending faults, overprinting and reorienting D3 dip-parallel stretching lineations. Metric-scale asymmetric Z-folds with northeast-trending axial planes affecting the mylonitic foliation are locally observed in the east-trending faults. These steeply plunging folds are commonly associated with a northeast-trending crenulation cleavage and record a late component of dextral movement along the east-trending faults. The 2 km-thick northeast-trending Fancamp deformation zone (FF; Fig. 2.4-2) displays a northeast-trending fold pattern and secondary crenulation cleavages with prominent vertical stretching lineations, and no shear sense indicators that are compatible with the northwest-southeast shortening (Legault et al., 1997). The timing of deformational events in the belt is of critical importance, and this can be elegantly resolved by understanding the plutonic emplacement history (Chown et al., 1992, 2002; Fig. 2.4-3). Synvolcanic plutons predate deformation and are responsible for the Dl folding phase. In contrast, southeast-trending dextral faults (D4) as well as the late horizontal dextral component in east-trending faults with late asymmetric Z-folds constrain the youngest dextral shearing event. The 2702-2692 Ma syntectonic plutons (Chown et al., 1992; Table 2.4-1), which cross-cut prominent De features include the Muscocho (2701 Ma), the Opemisca (2697 Ma) and Renaud plutons (2700 Ma; Figs. 2.4-1 and 2.4-2). The Colombourg pluton (2696 Ma) and the Franquet stock (2692 Ma) show magmatic fabrics compatible with D4 fault solid-state fabrics emplaced during southeast-Macamic and Cameron fault movement, respectively (Chown et al., 1992; Daigneault et al., 2002). Peak syntectonic pluton emplacement is placed around 2698 Ma

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Chapter 2: Generation of Continental Crust

Fig. 2.4-3. Time-space sequence of volcanic, plutonic and deformational events for the NVZ and SVZ of the Abitibi greenstone belt. Ages are compiled from Goutier et al. (1994), Davis et al. (2000), Davis (2002), Mortensen, (1993b) and Mueller et al. (1996).

2.4. Abitibi Greenstone Belt Plate Tectonics

97

(Fig. 2.4-3). Because southeast-trending faults cut regional folds, the regional schistosity and early east-trending faults, the prominent D2-D3 event of the NVZ occurred over 12 My between 2710 and 2698 Ma.

Southern Volcanic Zone (SVZ) The SVZ has a western Blake River segment (2703-2698 Ma; Mortensen, 1993b), considered an oceanic island arc composed of tholeiitic basalts and mafic-felsic volcanic calcalkaline rocks, and a complex eastern Malartic segment (2714-2701 Ma; Pilote et al., 1999) composed of komatiites to tholeiitic basalts, and andesitic to felsic calc-alkaline rocks. The Malartic segment is divided into the c. 2714 Ma Malartic Group and c. 2705-2701 Ma Louvicourt Group (Scott et al., 2002). The former has komatiites, basalts and felsic debris, interpreted as a submarine plain or plateau, whereas the latter displays subaqueous komatiites, pillowed andesites and lobate dacite-rhyolites, and is interpreted as an arc (Dimroth et al., 1982; Desrochers et al., 1993; Scott et al., 2002). Field relationships and age determinations support the notion of coeval plume- and subduction-generated volcanism. The c. 2705-2697 Ma volcanic cycle 3 is closely associated with synorogenic flyschty~pe sedimentary basins (sedimentary cycle 3, 2700-2685 Ma; Mueller and Donaldson, 1992a; Davis, 2002), occurring at the northern and southern margins of the SVZ as well as separating the two SVZ segments. The Lac Caste Formation, which occurs at the NVZSVZ interface, is an inter-arc turbidite sequence, which is the eastern prolongation of the Kewagama and Porcupine sedimentary rocks (Fig. 2.4-1). The Pontiac sedimentary rocks limiting the SVZ to the south, are composed of turbidites, minor conglomerate and pelagic background lithologies (Dimroth et al., 1982). The Pontiac has been interpreted as an accretionary wedge complex (Hodgson and Hamilton, 1989; Card, 1990). The diachronous strike-slip basins (sedimentary cycle 4, 2690-2670 Ma; Fig. 2.4-1), developed along the major E-trending faults within pre-existing flysch basins (Mueller and Corcoran, 1998), are the final volcano-sedimentary increment of oblique collision.

Structural styles and timing of deformation in the SVZ The SVZ, in comparison with the NVZ, exhibits different deformation styles. The deformation history is divided into several events concentrated along the two main terrane boundaries, the northern Destor-Porcupine-Manneville (DPMFZ) and southern CadillacLarder-Lake (CLLFZ) fault zones (Fig. 2.4-1). In order to understand the structural history of the SVZ, a separate nomenclature of deformation events is presented (SVZ DI-Ds), even though there are overlapping deformational events (Table 2.4-1; Fig. 2.4-3). Isoclinal folding, and prominent east-west striking foliations due to regional northsouth horizontal shortening, affecting both SVZ segments, define the Dl event. The Blake River segment has a relatively low strain central domain with higher strain zones at the margins. Fold axial traces around the synvolcanic Flavrian pluton have weak axial-planar foliation in contrast to a well-developed east-trending foliation at the southern and northern margins close to the faults. The Malartic segment is more heterogeneous. Folding is

98

Chapter 2: Generation of Continental Crust

developed locally, but panels of homoclinal rock sequences with a distinct regional foliation trend, separated by a series of layer-parallel faults that act as d~collement surfaces (Desrochers and Hubert, 1996) characterise the general behaviour. The accretion of the SVZ with the NVZ is characterised by thrusting (D2) exposed in the DPMFZ (Fig. 2.4-1). This fault zone, which exhibits different signatures along strike, features early thrusting in the eastern Manneville sector. Moderately to shallow-dipping mylonitic fabrics with a southwards vergence (30-40 ~ dip), are well developed in the Lac Caste sedimentary rocks and in the early phases of the Preissac-Lacorne batholith (Daigneault et al., 2002). An east-trending overturned regional synclinal fold within the deformation zone associated with these shallow-dipping fabrics, the presence of north plunging, dip-parallel stretching lineations, and a north-dipping mylonitic foliation is consistent with horizontal shortening. The transition from thrusting to dextral strike-slip (D3) is recorded in the western portion of the DPMFZ. The clastic sedimentary Duparquet basin, straddling the NVZ-SVZ boundary, has the hallmarks of a divergent fault-wedge basin, a variant of a pull-apart basin (Mueller et al., 1991, 1996). The sedimentology of these late molasse basins is an expression of the tectonic influence on basin evolution (see section 7.3). Structural elements consistent with dextral shearing are: (1) northeast-trending en-~chelon folds, (2) a northeast-trending cleavage developed oblique to the basin elongation, (3) subhorizontal stretching lineations, and (4) dextral shear-sense indicators. Duparquet basin formation between 2690-2680 Ma is based on U-Pb zircon ages of basin-related porphyry stocks (Mueller et al., 1996). The accretion of the SVZ with the Pontiac flysch deposits represents a distinct D4 event that is temporally related to D2 and D3. The Pontiac flysch deposits, an inferred accretionary prism (Card, 1990; Ludden et al., 1993) contain detrital zircons as young as 2685 Ma (Davis, 2002), that constrain the beginning of SVZ thrusting. The generally shallow dipping strata display a well-developed shallow to moderately inclined schistosity dipping to the north, with dip-parallel, north-trending stretching lineations and a series of overturned folds (Benn et al., 1994; Calvert and Ludden, 1999; Daigneault et al., 2002). These signatures are typical of a south vergent fold and thrust belt. A D5 event is responsible for the formation of the Granada basin that straddles the SVZPontiac boundary. This predominantly clastic marine basin with local 2673 Ma volcanism (Table 2.4-1), as well as coeval 2672 Ma porphyry stocks, shows strike-slip movement and resultant basin formation that is c. 5-10 My younger than Duparquet basin evolution. Unlike the Duparquet basin, Granada sedimentary rocks display a complex history of thrusting, extension and late dextral shearing (Daigneault et al., 2002). Renewed D6 thrusting affected the basin and is characterised by shallow-dipping deformation zones with moderately north-dipping schistosities and north-plunging stretching lineations. The schistosity is axial-planar to the Granada syncline (Goulet, 1978), which is a south verging overturned fold that resulted from shortening during southwards tectonic transport (Fig. 2.4-2). The logical consequence of shortening and dextral transpression is late extensional movement along the CLLFZ and the DPMFZ (DT) that accommodated continued stacking in the pre-existing basins. Both major fault zones experienced a late component of

2.4. Abitibi Greenstone Belt Plate Tectonics

99

extensional movement along earlier fabrics formed during D6 thrusting. A component of normal shearing is best recorded in the Granada basin, in which ubiquitous north-sidedown shear-sense indicators are observed. This movement along both fault zones resulted in uplift and exhumation of the Pontiac sedimentary and Malartic volcanic rocks. Metamorphic mismatches occur, with medium-grade amphibolite rocks south of the CLLFZ adjacent to subgreenschist to greenschist facies rocks north of the fault zone (Daigneault et al., 2002). Similarly, medium-grade amphibolite rocks south of the DPMFZ (Malartic segment) were juxtaposed with greenschist facies volcanic rocks north of the DPMFZ in the NVZ. Emplacement of late garnet-muscovite-biotite granitic suites between 2660 and 2642 Ma (Ducharme et al., 1997; Feng and Kerrich, 1991) was contemporaneous with this event (Daigneault et al., 2002; Table 2.4-1). Final dextral shearing (D8) in the SVZ affected east- and southeast-trending fault zones (Daigneault et al., 2002) and is exemplified by local and regional asymmetric northeasttrending Z-folds. The structural signature is evident in areas where D6 thrusts and D7 extensional structures are cross-cut. For example, an east-southeast-trending deformation zone in the Malartic segment cross-cuts and folds amphibolitic fabrics related to extension (Daigneault et al., 2002). In other areas identification is more difficult. The Cadillac fault (sensu stricto) in the CLLFZ, for instance, is interpreted as a result of the final dextral shearing event (Daigneault et al., 2002). The fault has subhorizontal stretching lineations that overprint the earlier down-dip stretching lineations and prominent dextral shear-sense indicators. Generally, a well-developed, northeast-trending, pressure-solution cleavage associated with asymmetric metre-scale Z-shaped folds facilitates recognition.

Discussion The timing of numerous deformation events clearly demonstrates a complex diachronous evolution of the belt (Fig. 2.4-4), but also shows that elements of certain deformation events can only be identified locally. In the Abitibi greenstone belt, an Archaean tectonic evolutionary scheme combining plate tectonics and subtle plume tectonism is proposed, but several aspects require additional consideration, including: (1) diachronous evolution of deformation events, (2) subduction zones and accretionary prisms, and (3) plume tectonism.

Migration of a deformation front: the notion of time and space The 20 My deformation history in the NVZ (2710-2690 Ma) encompasses dominant north-south shortening (NVZ-Dz-D3) that changed into dextral shearing (NVZ-D4; Fig. 2.4-5a). Early NVZ-DI remained a synvolcanic phenomenon. Whilst deformational events NVZ-DI-D3 affected volcanic and sedimentary cycles 1 and 2 in the NVZ, major subduction- and plume-generated volcanism in the eastern Malartic (2714-2701 Ma) and western Blake River (2703-2698 Ma) segments accounted for complex ocean floor and arc construction in the SVZ (Fig. 2.4-3). Deformation influenced the cycle 3 flysch basins linking them (2700-2685 Ma). The NVZ-D4 dextral shearing (2702-2690 Ma; Fig. 2.4-5b), with prominent southeast-trending faults in the NVZ, coincides with SVZ

100

Chapter 2: Generation of Continental Crust

Fig. 2.4-4. (a) General evolutionary model for the Abitibi greenstone belt showing major deformation events (modified from Daigneault et al., 2002); Cs = Lac Caste Formation; Gr = Granada Formation, Du = Duparquet Formation, Cd = Cadillac Group, Po = Pontiac Group, DPMFZ = Destor-Porcupine-Manneville Fault Zone, CLLFZ -- Cadillac Larder Lake Fault Zone. (b) Time framework with age determinations linked to major events in the Abitibi greenstone belt.

shortening and thrusting events ( S V Z - D I - D 2 ) that lead to N V Z - S V Z accretion. After the two zones docked, shortening could not be further accommodated by thrusting, so that

Opposite: Fig. 2.4-5. Palaeogeographic-tectonic evolution of the Abitibi greenstone belt between 2705 and 2661 Ma, displaying plume-arc interaction. Note the southwards migration of the deformation front in time and space. (a) NVZ-D 2 and -D 3 shortening events are contemporaneous with SVZ volcanic activity, that displays the subduction-generated Noranda caldera (Blake River segment) coeval with plume-induced komatiites of the Jacola Formation (Malartic segment). (b) The NVZ-D4 dextral shearing is responsible for southeast-trending dextral faults and dextral reactivation of east-trending reverse faults that created space for syntectonic plutons. In the SVZ, folding and thrusting were prominent (SVZ-DI and -D2). (c) The last stage displays diachronous dextral shearing events with the formation of Duparquet (SVZ-D3; 2690-2680 Ma) and Granada pull-apart basins (SVZ-D5; 2680-2670 Ma) along the major crustal-scale structures which were also related to thrusting events SVZ-D4 and-D6.

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Chapter 2: Generation of Continental Crust

SVZ-D3 transcurrent motion ensued with the development of the Duparquet strike-slip basin (2690-2680 Ma) along the DPMFZ, the NVZ and SVZ interface. A terrane docking event was completed. The SVZ deformation history (Fig. 2.4-4, Figs. 2.4-5b, c), spanning c. 58 My (26982640 Ma), represents a continuation of structural events migrating southwards that produced a fold and thrust front. Thrusting, which affected the flysch deposits (e.g., Lac Caste sedimentary rocks; SVZ-D2) connecting the SVZ with the NVZ, subsequently deformed the southern Pontiac flysch deposits (SVZ-D4). Evidence of a renewed strike-slip event (SVZ-Ds) is chronicled by the 2680-2670 Ma Granada basin that straddles the SVZ and the Pontiac accretionary prism along the CLLFZ. In contrast to the Duparquet basin, renewed thrusting (SVZ-D6) influenced Granada basin geometry. Interestingly the extensional exhumation phase (SVZ-D7) between 2660 and 2640 Ma is well documented in the Granada basin (Daigneault et al., 2002) and the Malartic segment where komatiites and late granitic pluton phases are abundant. A final dextral shearing SVZ-D8 event produced the Cadillac Fault in sensu stricto. This chronological review shows a systematic time space evolution of an arc, arc-arc collision and arc fragmentation, which is recorded both in the volcano-sedimentary and in the structural history. Subduction zones and accretionary prisms The CLLFZ and the DPMFZ, with various deformation styles, record the protracted history of the belt. The turbiditic flysch basins, locus of these fault zones, represent accretionary prisms and trench-related subduction zones with a typical fold and thrust belt geometry. The structural style within the Pontiac and the Caste flysch basins are compatible with a southwards-vergent subduction zone as supported by Lithoprobe seismic reflection data (Calvert and Ludden, 1999). The restricted volume of Caste sedimentary rocks in comparison with the extensive Pontiac flysch terrane (Fig. 2.4-1) is explained by their advanced stage of subduction and by recycling of sedimentary rocks. The DPMFZ and CLLFZ are interpreted as terrane zippers or suture zones representing the expression of relict subduction. The northern DPMFZ subduction zone with the Caste accretionary prism was active during the NVZ deformation, with a transfer to the southern CLLFZ subduction zone during SVZ deformation (Fig. 2.4-4, Figs. 2.4-5a, b). The two-mica garnet granites (e.g., Preissac-Lacorne batholith) in the Caste and Pontiac accretionary prisms (Feng and Kerrich, 1991) are S-type granites generated by partial fusion of the sedimentary rocks (Calvert and Ludden, 1999), lending further support for a relict subduction zone. Plume influence during Abitibi evolution Plume influence on the deformation pattern can be questioned for the Abitibi greenstone belt (Fig. 2.4-5a). The late exhumation phase D7 in the SVZ is the most likely tectonic event that can be connected to plume influence. This does not preclude a plume interaction during early deformational events, but only the shortening and shearing components related to plate tectonic processes could be demonstrated (Fig. 2.4-5c). During the waning stages of plate tectonic processes, the upwelling related to the long-live plume influence becomes a driving force that could have been responsible for extensional movement along

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pre-existing major faults and for the exhumation of the Pontiac terrane (Fig. 2.4-4). This result is compatible with the higher metamorphic grade observed in the Pontiac terrane and in the SVZ. Conclusion

The Abitibi greenstone belt evolved over c. 100 My with volcano-sedimentary sequences developing between 2735 and 2670 Ma and late plutonic activity occurring between 2670 and 2640 Ma. The belt displays the salient features of arc evolution, arc-arc collision and arc fragmentation with the recognition of strike-slip basins. Identifying the various plutonic suites is important because radiometric age determinations of plutons best chronicle the deformation history. The stratigraphic relationships and detrital zircons show that sedimentary basins of cycles 1 and 3 had a protracted history from inception to subsequent stages of deformation. Sedimentary cycle 4 strike-slip basins are well constrained and are restricted to the major faults; they best display the diachronous development of basin-forming events. All the deformation history is recorded in the sedimentary basins, especially along the major fault zones. The DPMFZ and the CLLFZ, two terrane boundaries, are considered suture zones representing relict subduction zones. Oblique convergence can explain the observed complex fault pattern. Individual fault-bounded blocks are not isolated terranes but rather part of an ongoing sequence of deformational events. The structural events in the Abitibi greenstone belt display the classical features of a modern orogenic belt with the constant interplay between thrusting and strike-slip motion, as well as final extension which is generally due to overstacking in modern sequences. Alternatively, exhumation in Archaean greenstone belts could also be readily explained by plume upwelling. In the areas where exhumation is a prominent feature, komatiites are an important constituent of the sequence. Although the tectonic influence of plumes is difficult to quantify, extensional structures, active during the terminal stage of arc evolution after plate forces had dissipated, may have been related to plume activity. The time-space sequence of volcanic and plutonic activities with the southwards-migrating deformation front, however, is more compatible with a plate-tectonic process dominated by subduction and oblique collision.

2.5.

GRANITE FORMATION AND EMPLACEMENT AS INDICATORS OF ARCHAEAN TECTONIC PROCESSES

T.E. ZEGERS Introduction

An essential step in the generation of continental crust is the production, transport and emplacement of granitoid magmas. Granitoid rocks are a major component in all Archaean 771e Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, I).R. Nelson, W.U. Mueller and O. Caluneanu

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terranes and so understanding the process by which they were formed is a key to understanding the process behind the formation of granitoid-greenstone terranes. Discussion has centred on the question of how granites were emplaced. Rival opinions have either favoured uniquely Archaean (solid-state) diapiric processes that are the result of buoyancy forces in the absence of far-field tectonically induced stresses (Hickman, 1984; Choukroune et al., 1995; Collins et al., 1998; see section 2.6), or far-field induced deformation, possibly in a present day-style tectonic setting, during compression (Bickle et al., 1993; de Wit et al., 1987a), extension (Zegers et al., 1996), or strike-slip deformation (Zegers et al., 2001). Experimental work on granitic melt production (Rapp, 1997; Wyllie et al., 1997), and detailed studies of the physical processes that control granitoid extraction, transport and emplacement (Petford et al., 2000, and references therein) have also considerably enhanced our understanding of all aspects of granitoid formation. The purpose of this contribution is to integrate these current ideas on granite formation with field observations in Archaean terranes, to review the potential geodynamic processes that ultimately led to the formation of Archaean continental crust. Archaean Granitic Rocks: General Features and Time Trends

Granite-greenstone terranes typically contain more than 60% granitic rocks of varying age, composition, and degrees of deformation and metamorphism (de Wit and Ashwal, 1997b). In some, mostly Early Archaean cratons, composite granite batholiths form ovoid structures, surrounded by volcano-sedimentary sequences. Typical examples are the Pilbara and Zimbabwe granite-greenstone terranes. However, the majority consist of elongate alternating belts of granites and greenstonesmthe Late Archaean Yilgarn craton in Australia and the Superior Province in Canada provide examples of these (Fig. 2.5-1). Detailed geochronology and geochemistry of granites and greenstone sequences has shown that many felsic to intermediate volcanic rocks are the extrusive counterparts of granites (e.g., Zegers et al., 1998b, for an overview). Within each granite-greenstone terrane the majorelement composition of granites shows a secular change from tonalite-trondhjemitegranodiorite (TTG) to granodiorite-granite-monzogranite (GGM) to the highest K20 syenite-granite (SG) suites (Bickle et al., 1989, 1993; Feng and Kerrich, 1992; Zegers et al., 1998b). Syenite-granite suites are typically post-tectonic, whereas TTG and GGM suites are generally pre- to syntectonic (see Fig. 2.5-1). The internal structure of batholiths is variable and often complex. Batholiths consist of both (migmatic) gneisses and relatively undeformed granites. In general, field relationships suggest that pre- to syntectonic granites intruded originally as subhorizontal sheets or laccoliths in both ovoid and linear granitegreenstone terranes (de Wit et al., 1987a; Chown et al., 1992; Zegers et al., 1996; Collins et al., 1998; Kloppenburg et al., 2001). In ovoid granite-greenstone terranes, subsequent doming resulted in the ring-sheet structures of batholiths comparable to onion-rings. Posttectonic granites transect the original sheeted or onion-ring structure and are often associated with major late strike-slip shear zones (Chown et al., 1992; Van Kranendonk and Collins, 1998; Zegers et al., 1998a).

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Fig. 2.5-1. Simplified overview of geological events in the eastern segment of the Early Archaean Pilbara craton, the Late Archaean Abitibi belt (Superior Province), and Kalgoorlie segment of the Yilgarn craton. The different granitic components are shown in the time line with respect to deformation events and the deposition of volcanic rocks and clastic sediments (greenstones). Early Archaean terranes typically evolved over a long period (600 My), whereas Late Archaean terranes formed within 200 My. Below are schematic maps of two terranes to illustrate the difference between ovoid (Pilbara) and linear (Yilgarn) granite-greenstone terranes.

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Granites and gneisses of the TTG suite are characteristic of Archaean terranes and volumetrically important in some. In the Archaean where TTG suites are absent or volumetrically minor (Yilgarn), the dominant monzogranite suites are derived from reworking of TTG crust (Champion and Smithies, in prep.). Therefore the production of large volumes of TTG melt represents the first and essential step to generating continental crust. TTGs are high Na20/K20 felsic magmas with an expanded silica range consistent with an origin from basaltic precursors. Two subgroups of TTGs are recognised: the high-A1 and low-A1 series, largely reflecting the depth of partial melting (Barker and Arth, 1976). The most common high-A1 subgroup, characterised by high A1203, high Sr, steeply fractionated REE patterns, low HREE and absence of Eu anomaly, is consistent with generation by partial melting of hydrated basalt in the high-pressure garnet stability field (Rapp, 1997; Wyllie et al., 1997; Fig. 2.5-2). Conversely, the low-A1TTG subgroup is characterised by lower A1203, Sr and LREE/HREE, higher HREE and the presence of a negative Eu anomaly, consistent with generation under lower pressures in the plagioclase stability field. The younger GGM suites and SG suites are usually interpreted as the result of crustal melting of TTG, perhaps together with mixing with magma derived from a mantle source (e.g., Bickle et al., 1989; Feng and Kerrich, 1992; Collins, 1993; Bedard and Ludden, 1995). A distinctive, but volumetrically minor type of post-tectonic granite are those of the highMg diorite, or sanukitoid suite (Stern et al., 1989). The chemistry of these rocks is consistent with either melting of a mantle source that was previously metasomatised by TTG-like melt, or with peridotite contamination of TTG-melt during ascent through a mantle wedge (Smithies and Champion, 2000, and references therein). Geodynamics of Tonalite-Trondhjemite-Granodiorite-Granite Formation

The conditions under which tectonic or geodynamic processes operated, and how they led to granite formation differed in the Archaean. The common basis for the differences between the Archaean Earth and present-day Earth arises from the higher heat production from radiogenic isotopes in the Archaean. How much higher the heat production was is still a matter of debate, but estimates range from 2 to 6 times the present-day heat production (Pollack, 1997). Field evidence from komatiites (Abbott et al., 1994; Arndt et al., 1998, and references therein), and thermal modelling (Pollack, 1997) suggest that this led to a mean mantle temperature that was at least 150~ hotter at c. 3300 Ma. This has important consequences for geodynamic processes in the early Earth. Mantle viscosity drops by approximately one order of magnitude for each 100~ temperature increase (Karato and Wu, 1993), resulting in more vigorous mantle convection. In such a hot mantle, decompression melting also starts deeper, causing a significantly thicker basaltic crust to be formed, which is underlain by a thick and possibly stable stratified harzburgitic mantle residue. The enhanced compositional stratification as a result of the thicker, hotter and therefore less dense oceanic crust, is such that gravitational instability, necessary for subduction, may not be reached in geologically realistic time scales (Davies, 1992a; Vlaar et al., 1994). Archaean oceanic crust was possibly similar to very thick present-day oceanic plateaus,

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not only in terms of thickness, but also in terms of stratification, structure and composition (Kusky and Kidd, 1992; Condie, 1997b; Polat et al., 1998). Whether the continental crust was hotter than today is a point of contention. Heat-producing elements K, Th, and U are concentrated in continental crust with respect to the mantle, leading to higher crustal heat production in the Archaean, and therefore higher steady state geotherms (Sandiford, 1989a; Kramers et al., 2001). However, steady state geotherms would not have exceeded the solidus because the geochronological record shows that granitic rocks were not produced continuously by lower crustal melting. Even a slightly hotter geotherm would result in higher temperature metamorphism and in a shallower brittle-ductile transition, as discussed by Marshak (1999). This would limit the amount of crustal thickening that could occur before gravitational collapse (Dewey, 1988; Bailey, 1999). In the section below, the effects that these specific Archaean conditions might have had on geodynamic processes and granite formation are examined.

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Tonalite-trondhjemite-granodiorite (77"G)formation The bulk Archaean continental crust is much more silicic than oceanic crust. As direct derivation of silicic rocks from a mantle source is extremely uncommon, the silicic composition of continental crust must be the result of partial melting of a mafic precursor. However, while TTGs are generally regarded as the product of melting of hydrated basalts, differentiating a thick oceanic plateau-like crust into an upper crust of TTG composition and a residual lower crust of mafic granulite or eclogite, is not sufficient to produce Archaean continental crust. The volume of mafic residue required from production of TTG exceeds the volume of residual mafic lower crust present in Archaean terranes. This means that part of the residue must have been recycled back into the lithospheric mantle. Mass balance calculations reconciling the compositional differences between continental crust and depleted mantle (Taylor and McLennan, 1995; Rudnick et al., 2000) show that this recycled component is most likely of eclogitic composition. Recent studies of xenoliths from several Archaean cratons demonstrated that eclogitic components have compositions complementary to TTG suites (Ireland et al., 1994; Jacob and Foley, 1999; Rollinson, 1997; Barth et al., 2001; Shirey et al., 2001). Also, seismological observations of Archaean cratons suggest the crust is thinner (c. 35 km) than in post-Archaean terranes, and lacks the basal high velocity layer attributed to garnet-bearing granulite or eclogite (Durrheim and Mooney, 1994). Two general geodynamic models for TTG generation are consistent with the abovementioned considerations. The first model invokes shallow subduction (see also section 3.5) of a thick and hot oceanic lithosphere (Martin, 1986; Drummond and Defant, 1990; Davies, 1992a; Martin and Moyen, 2002; Fig. 2.5-3), whereas the second model involves in situ crustal differentiation and delamination (Glikson, 1972; Anderson, 1979; Kr6ner, 1985a; Vlaar et al., 1994; Zegers and van Keken, 2001; Fig. 2.5-4). Although mantle plumes (sections 3.2 and 3.3) may have been an important factor in Archaean geodynamic processes, "plume tectonics" cannot provide a general model for the production of TTG melts and for the recycling of eclogites. As pointed out by Davies (1992a), plumes are complementary to subduction processesmtheir upwelling may be an Archaean equivalent to present day mid-oceanic ridge processes, but they are not an Archaean equivalent for downwelling subduction-like processes. The shallow subduction model involves the production of TTG by partial melting of relatively hot and buoyant subducting oceanic crust (Martin, 1986). TTG melt rising from the subducting slab interacts with the mantle wedge and can thus be regarded as Archaean analogues to adakites (Martin, 1999), or slab-melting occurs at such shallow depth that an overlying mantle-wedge is absent (Martin and Moyen, 2002). The geothermal gradient along the Benioff plane, and hence the mantle temperature, should be relatively high to reach the temperature and pressure conditions necessary for partial melting (see subduction geotherm in Fig. 2.5-3). However, such a high mantle temperature would inevitably result in a thicker oceanic crust and thicker and more depleted and buoyant lithospheric mantle (Davies, 1992a; Vlaar et al., 1994). This conflicts with the subductable, hence gravitationally unstable thin and cool oceanic crust envisaged in the slab-melt models. The majority of observations consistent with this scenario come from Late Archaean terranes,

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Fig. 2.5-3. (a) Schematic representation of a shallow subduction model for the Superior Province with slab melting to produce TTGs. Adapted after Calvert and Ludden (1999). (b) Pressure/temperature diagram showing an estimate of the geothermal gradient along the Benioff zone (Martin, 1986). This estimated gradient transects the solidus and melt contours (Rapp, 1997; Wyllie et al., 1997) to create the conditions for partial melting of hydrated basalt in the garnet stability field, leading to TTG melt. Dashed lines are TTG melt contours (see Fig. 2.5-2). (c) Interpreted Lithoprobe seismic section of the Superior Province (Calvert and Ludden, 1999). The shallow subduction model for TTG melt formation is most consistent with observations from Late Archaean terranes such as the Superior Province. Note the complex amalgamation of rock units with variable affinity and composition.

in particular from the Canadian Superior Province, but also from the 3.0 Ga central Pilbara granite-greenstone terrane. Such observations include the linear large-scale structure of diachronously accreted terranes (Calvert and Ludden, 1999), seismic evidence for slab-like features (Fig. 2.5-3c) (Calvert et al., 1995), and TTG compositions consistent with mantle wedge interaction. Boninite-like rocks, sanukitoids and high-Nb basalts provide geochemical evidence for a subduction-modified mantle (Smithies and Champion, 2000; Polat and Kerrich, 2001b; see also section 2.3). In addition, TTG intrusion occurred during deformation (Chown et al., 1992), as expected in a shallow subduction setting where slab-retreat does not occur (Jordan et al., 1983; Royden, 1993).

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Fig. 2.5-4. (a) Schematic representation of lower crustal delamination in a thick (45 km) oceanic plateau-like crust. The lower crust converts to eclogite or garnet granulite, with a density higher than the depleted harzburgite mantle, and delaminates from the middle crust. This results in geological events typical of delamination (Kay and Kay, 1993): crustal melting, in this case forming TTG melt, uplift and extension. Adapted from Zegers and van Keken (2001). (b) Pressure temperature diagram showing the stability field of eclogite or garnet granulite with density > 3.3 g/cm 3 (shaded area), solidus and melt contours are the same as in Fig. 2.5-2. Geotherms are calculated for t = 1, 10, and 25 My after delamination, indicating that up to 30% partial melt, to produce TTG, is possible within 20 My (Zegers and van Keken, 2001). (c) Schematic crustal section in the Eastern Pilbara granite-greenstone terrane after delamination and TTG intrusion and volcanism at c. 3400 Ma; upper 20 km is based on field observations, geochronology and restoration of subsequent deformation features (Zegers et al., 1996, 2001).

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However, a shallow subduction model for TTG generation cannot explain many of the unique features of Early Archaean cratons. Subduction requires a relatively thin oceanic crust, leaving a relatively thin residual lithospheric mantle behind. The presence of an anomalously thick and depleted mantle beneath Archaean cratons, which at least partly formed during the Middle Archaean (Boyd and McCallister, 1976; Pearson, 1999), is therefore not directly consistent with subduction. To resolve this paradox, models have been proposed, in which the crust and lithospheric mantle formed separately and were juxtaposed at a later stage by a thrust-like mechanism (Herzberg, 1999; Moser et al., 2001). The absence of the lower crustal high-velocity zone, an ovoid large-scale structure, and the absence of igneous rocks that show clear evidence of having interacted with an enriched mantle source (i.e., mantle wedge), and the low Mg-numbers of most TTGs compared to modem-day adakites (Smithies, 2000), are other features of Early to Middle Archaean terranes not directly compatible with a subduction in situ. The combination of these features is more consistent with a delamination model (Fig. 2.5-4). The in situ differentiation and delamination model (Zegers and van Keken, 2001) involves a thick oceanic plateau-like mafic crust, in which the lower part converts into denser eclogite or garnet-rich granulite (Fig. 2.5-2). Depending on the initial geotherm this may involve partial melting, leaving a garnet-bearing residue, and producing high-A1TTG melt. Eclogite and garnet granulite have higher densities than the underlying depleted mantle, resuiting in the delamination or convective thinning of the lower crust, and (1) recycling of this material to the lithospheric, and possibly (2)convecting mantle (Anderson, 1979; Vlaar et al., 1994). The lower crust is replaced by hot mantle, which undergoes decompressive melting. Conductive heating and intrusion of mantle melt, melt a large section of the mafic crust that previously lay directly above the delaminated crust producing TTG. TTG melt derived from the lowest part of the crust (below garnet-in line, Fig. 2.5-4) would leave garnet and hornblende in the residue (Wolf and Wyllie, 1993), and generate high-A1TTG consistent with Nb/Ta and Zr/Sm trace element data from Archaean TTGs (Foley et al., 2002). In contrast, TTG melt derived from the crust above the garnet-in line would have a low-A1 signature. Early models, in which delamination played a role in continental crust formation (e.g., Glikson, 1979; Hoffman and Ranalli, 1988; Kr6ner and Layer, 1992), have been dismissed because it was assumed that both crust and lithospheric mantle delaminate, which is contradicted by the Middle Archaean ages obtained for the Kaapvaal cratonic mantle (Pearson, 1999). However, the eclogitic/garnet-granulite lower crust can delaminate and fall, or drip, through the depleted mantle, leaving the buoyant part of the mantle in place (Zegers and van Keken, 2001). Eclogitic components found in the xenoliths from the Kaapvaal lithospheric mantle (Shirey et al., 2001) may partly represent delaminated lower crustal eclogite that reached neutral buoyancy in the lithospheric mantle. The delamination model is consistent with many of the observations of Early to Middle Archaean cratons, such as the eastern segment of the Pilbara craton and Barberton greenstone belt. In particular, it explains the recorded crustal extension (Nijman and de Vries, section 2.7; Zegers et al., 1996) and core-complex development (Zegers et al., 2001) during uplift and intrusion and extrusion of TTG magmas (Fig. 2.5-4c). This combination of

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events is characteristic of delamination in modern terranes (Kay and Kay, 1993). The delamination model is also inherently consistent with the formation of the thick and depleted harzburgite mantle in the Archaean as the complementary residue of basaltic to komatiitic melt (Herzberg, 1999) to produce the initially thick (> 35 km) oceanic plateau crust. Because the model does not involve the formation of a mantle wedge, it is also consistent with the low Mg-numbers of many TTG granites and the lack of other Early Archaean igneous rocks that show evidence for derivation from an enriched mantle source. Although the delamination model supports many of the observations of the earliest TTG events in Early to Middle Archaean granite-greenstone terranes, a present-day analogue is lacking, and some of its features challenge our understanding of geological processes. The initial thickness of the oceanic plateau must have exceeded 35 km for the lower part to have been in the eclogite stability field. This is thicker than the thickest present-day oceanic plateau (the Ontong Java Plateau which is about 35 km) but is not inconsistent with previous estimates, which suggest that Archaean oceanic crust may have exceeded 40 km (Hoffman and Ranalli, 1988; Davies, 1992a; Vlaar et al., 1994). In addition, for partial melting to occur, this mafic crust must have been, at least locally, hydrated to a depth of between 15 and 35 km. The study of modern oceanic plateaus shows that hydration to great depth can be achieved (1) by stacking to great thickness (> 14 km) of subaqueous basalt flows, and (2) by rotation of lava piles to steeply dipping sheets, as in Iceland (Saunders et al., 1996). Alternatively, the thick hydrated mafic sequence may be the result of obduction as proposed by de Wit et al. (1992). Another aspect is the rheology of the depleted mantle during delamination. The large volume of melt extraction from the primitive mantle results in a relatively cool, viscous and buoyant subcratonic lithosphere (Jordan, 1988). Although the viscosity of the Archaean subcratonic lithosphere was higher than the surrounding mantle, its rheology would still have allowed ductile deformation, necessary for delamination, to have taken place (van Thienen et al., in press). Any geodynamic model for the formation of early continental crust must consider the production of a sufficiently large volume of TTG, and its derived granites, in the presentday continental crust of Archaean age. Whereas in the subduction model, the supply of TTG melt is unlimited by the cycling of fertile oceanic crust through the partial melting zone, the TTG volume in the delamination model is limited by the initial thickness of the oceanic plateau. Theoretically a mass balance approach can be used to test if sufficient TTG melt can be produced in the delamination model. However, some of the most crucial factors are not well constrained. Using the eastern part of the Pilbara craton as an example, the following variables have to be taken into account: (1) The TTG magmatic rocks produced at c. 3450 Ma, or derived from remelting of 3450 Ma TTG (Smithies et al., in press). The Pilbara crust is 32 km thick (Durrheim and Mooney, 1994), with a c. 2 km lower crustal mafic high velocity zone and a c. 5 km mafic volcanic rocks in the upper crust. This leaves a c. 25 km of rocks of felsic or intermediate composition. An upper estimate for the present TTG volume is therefore a layer of 25 km. If the lower 20 km is intermediate (equal parts mafic gneiss and TTG gneiss) instead of felsic in composition, this number may decrease to about 15 km as a lower estimate. (2) The melt percentage is constrained by the hornblende-out line

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(Fig. 2.5-2) and the solidus. Reasonable partial melt percentages to produce TTG compositions are 10-40% (Rapp, 1997). As an upper estimate, starting with a 45-km thick mafic crust and assuming that melt is derived from both the delaminated portion and the overlying crust, an average of 30% melt from 35 km of mafic crust would produce 9.5 km of TTG. This is not sufficient to reach the lower estimate of total TTG volume (15 km). Therefore, a Middle Archaean delamination event from 45 km initial mafic crust is not sufficient to produce the present stable Archaean continental crust. Either the initial oceanic plateau was thicker than 45 km (by stacking, or by increased melt extraction from the mantle to form a thicker oceanic plateau), or the crust was thickened after the production of the initial TTG-producing delamination event. Compressional structures, such as thrusting and folding, and medium pressure metamorphism associated with crustal thickening are well documented in the Pilbara and Kaapvaal cratons for the c. 3200 Ma period (Bickle et al., 1980; Boulter et al., 1987; de Ronde and de Wit, 1994; Nijman et al., 1998b; Dziggel et al., 2002). Therefore at least part of the discrepancy between TTG volume produced by delamination at c. 3450 Ma and the present volume of TTG can be explained as a result of crustal thickening post-dating TTG production. Once a large volume of TTG melt was produced and the residue was recycled back into the mantle, either by subduction or by delamination, the first Archaean continental crust was formed. The structure and composition of this early continental crust depends on the process responsible for TTG production. If we accept the subduction and slab melt model, then we would expect a complex and deformed array of accreted oceanic crust/plateau and volcanic arcs of TTG composition (Calvert and Ludden, 1999; Fig. 2.5-3c). Accepting the in situ delamination model, we would expect a relatively simple and superficially littledeformed crust consisting of a refractory lower crust of mafic gneiss, a middle gneissic crust composed of TTG, amphibolite and residual material from melt extraction, and an upper crust containing TTG lacoliths, oceanic plateau, plateau basalts, and gabbro sills (Fig. 2.5-4c). Granite formation After the initial formation of Archaean continental crust and before final stabilisation (Fig. 2.5-1) considerable volumes of granite (GGM and SG) melt were produced, primarily by melting of pre-existing TTG, but possibly also mixed with a mantle-derived component. Studies have shown that most modern granites form by dehydration melting rather than by fluid-present melting (Thompson, 1999). In the case of Archaean continental crust of TTG-basalt composition, crustal melting would be governed by the breakdown of hydrous minerals such as amphibole and biotite (Patifio Douce and Beard, 1995; Singh and Johannes, 1996). In general, the curves for these minerals have steeply positive or curved negative slopes, and breakdown requires temperatures in excess of 700~ (Fig. 2.5-5). The timing of granite formation in Archaean granite-greenstone terranes (Chown et al., 1992; Zegers et al., 1998b; Nelson et al., 1999) in episodes of c. 40 My indicates that the steady state geotherm did not exceed the solidus. Therefore, to produce crustal melt, the geotherm must have been raised by the addition of heat.

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Chapter 2: Generation of Continental Crust

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mafic gneiss \

2.5. Granite Formation and Emplacement

115

Several tectonic events or settings have been proposed where the addition of heat to the middle and lower crust may lead to melting (Brown, 1994; Thompson, 1999; Petford et al., 2000). Melting may occur in crust thickened by compression (Fig. 2.5-5a), by addition of heat to the lower crust during underplating and intrusion of basaltic melt derived from a mantle plume (Fig. 2.5-5b), by extensional thinning of the crust and intrusion of basaltic melt produced via decompression melting of the upper mantle (McKenzie and Bickle, 1988; Fig. 2.5-5c), by delamination of (previously thickened) lower crust and lithospheric mantle (Houseman and Molnar, 1997; Fig. 2.5-5d) and through intracrustal heating due to high concentrations of radiogenic elements in the crust (Chamberlain and Sonder, 1990; Fig. 2.5-5e). The presence of plate-tectonic stresses is an integral component of all these crustal melting mechanisms, except for the cases of melting influenced by a mantle plume or radiogenic heating. In modern terranes, crustal thickening in combination with delamination and addition of basaltic melt to the lower crust are regarded as the most effective ways to produce crustal melt. Crustal extension and radiogenic heating are not considered to be effective under most modern-day conditions (Thompson, 1999; Petford et al., 2000). Under Archaean conditions, however, the situation may have been very different, with melting by radiogenic heating, delamination and extensional collapse likely to have been most important. Radiogenic heat production was considerably higher (2 to 6 times) during the Archaean, and in situ crustal melting was much more likely to have been effective. Ridley

Opposite: Fig. 2.5-5. Schematic and largely qualitative representation of different processes that can add heat to the lower and middle crust to increase the geotherm, so that dehydration melting can occur. Solid thick lines are the water saturated solidus for the haplogranite system (Singh and Johannes, 1996), the predicted dehydration line (straight line, C and W; Clemens and Wall, 1981) and experimental dehydration line (curved line, S and J; Singh and Johannes, 1996) for a biotite-plagioclase-quartz assemblage. The initial geotherm as shown (solid line, IG) is modelled (by Kramers et al., 2001) for the Zimbabwe craton at 2.6 Ga. The geotherm after the heating is dashed. (a) The effect of crustal thickening, and subsequent erosion and uplift on the geotherm and, pressure-temperature paths (short dashes; after Thompson, 1999). Note that it may take a considerable amount of time (up to 120 My) after crustal thickening before crustal melting occurs; shown by the short dashed particle path. (b) A mantle plume under Archaean continental crust. Heating is the result of heat conduction from the underplated mantle melt and heat advection by intrusion of basaltic melt into the lower mantle. The time lag is expected to be small (Petford and Gallagher, 2001). Melting is concentrated in the lower portion of the crust. (c) The effect of crustal extension and resulting mantle upwelling. Heating is the result of intrusion and underplating of mantle melt. The amount of melt is expected to be small, because the temperature increase is expected to be minimal. (d) The effect of delamination of lower crust and upper mantle. Heating is the result of the replacement of a relatively cool lithospheric section by hot asthenospheric mantle, resulting in mantle melt, underplating and intrusion. Melting is expected to be considerable because heat is added to mid-crustal levels. (e) Addition of heat to the crust by decay of radiogenic elements. In the Archaean this crustal heat source was considerably higher than in modem continental crust. Heating was most likely concentrated in the middle crust where most TTG granites were concentrated.

116

Chapter 2: Generation of Continental Crust

and Kramers (1990) and Ridley (1992) modelled the effects of crustal heat production in an Archaean crust composed of TTG. They suggest that K-rich granites in many Archaean granite-greenstone terranes may be the result of lower crustal melting in an otherwise undisturbed crust. A plume source for crustal heating has also been suggested for several Archaean belts, including craton-wide granite intrusion at c. 2700 Ma in the Yilgarn craton (Campbell and Hill, 1988) and c. 3000 Ma GGM granites in the Pilbara craton (Collins et al., 1998). Crustal thickening and subsequent extensional collapse have been proposed as additional ways to produce crustal melt. In many cases, GGM granites are broadly co-eval with, or directly post-date compressional deformation (Kusky, 1993; de Ronde and de Wit, 1994; Sawyer and Barnes, 1994; Nelson, 1997; Zegers et al., 1998b). Dirks and Jelsma (1998) suggested that crustal melting in the Zimbabwe craton was a direct consequence of crustal stacking. Delamination of the lower crust and part of the mantle lithosphere after crustal thickening has been proposed to explain the occurrence of the late granites in the Superior (Moser et al., 1996) and Yilgarn cratons (Qiu and Groves, 1999). Such processes may also be a direct result of higher heat production in the crust and mantle, resulting in a higher geotherm and therefore a weaker lower crust, more prone to extensional collapse (Bailey, 1999) and delamination. But gravitational instability of the lower crust, and consequent delamination, is also enhanced by the low density of the underlying depleted harzburgite mantle, produced by enhanced melt extraction as discussed below.

Emplacement Recent studies of the physical processes of granite extraction, transport and emplacement have been reviewed by Petford et al. (2000) and have resulted in new insights into the emplacement of Archaean granites. It is now thought that granitic melt has a relatively low viscosity in the range of 108-103 Pas. The low viscosity and high volume increase during melting, may lead to deformation-enhanced segregation that can occur at very low melt fractions (< 5%). The traditional idea of viscous magma ascending through the continental crust as diapirs has been largely replaced by models of magma transport through dykes and along pre-existing faults and shear zones. Emplacement is thought to occur as subhorizontal magma sheets, or laccoliths, consistent with the tabular three-dimensional shape of plutons. As elegantly shown by Vigneresse et al. (1999), subhorizontal intrusion occurs under all stress conditions, including extension and strike-slip conditions, as a result of reorientation of the stress field when granite dykes reach shallower levels of the crust. The level at which magma sheets form is controlled by rheological contrasts rather than by density contrasts. Applying these new insights to an Archaean terrane where granite was emplaced within an originally basaltic crust during numerous intrusive episodes, and under a range of stress conditions, invites the following simplified scenario (Fig. 2.5-6). TTG melt was transported in structurally controlled feeder dykes and intruded as flat lying sheets during extension (eclogite-delamination model) or compression (flat-subduction model). The level of intrusion was most likely determined by the brittle-ductile transition and by pre-existing discontinuities. TTG intrusion formed a rheological and thermal boundary, due to concen-

2.5. Granite Formation and Emplacement

117

Fig. 2.5-6. Conceptual model for the formation of domal geometry in Archaean terranes by episodic intrusion of granites into an initially basaltic crust. (a) Intrusion of TTG granites as flat-lying sheets at the brittle-ductile transition. Transport of TTG melt though subvertical dykes. Partial melting (X) in the lower crust. Mafic melt residue remains in the lower crust, or is delaminated if the residue is rich in garnet. (b) Subsequent intrusion of sheeted granites above and under the pre-existing TTG granites. The TTG plutons act as a rheological and thermal boundary (schematic isotherms are shown as dashed lines), resulting in concentration of intrusions at this level. After erosion to deeper levels the granite complexes will have a ring-sheet geometry.

tration of heat-producing elements. As a result, subsequent granites intruded as sheets under or directly above the existing TTG intrusions. As illustrated in Figure 2.5-6, this would inevitably produce a domal geometry over increments of granite addition. Depending on the regional deformation, the domal geometry may be modified during and after intrusion of granites. Intense compressional deformation and thrusting may have superimposed the large-scale linear fabric (in map view) seen in most Late Archaean terranes. Core-complex formation during extension would enhance the domal structure (Zegers et al., 2001). The domal structure of ovoid granite-greenstone terranes should therefore be attributed to the

Chapter 2: Generation o f Continental Crust

118

lack of intense compressional deformation after granite intrusion, rather than to a unique Archaean diapiric emplacement mechanism. Conclusions

The complexity of multi-stage granite formation and intrusion in Archaean terranes cannot be contrained in any single geodynamic model. As in modern belts, many different geodynamic processes operated, resulting in a complex sequence of geological events during which granite formation and intrusion played an important role. Although the physical and chemical requirements for granite formation remained unchanged through time, the environments in which these requirements were met were different during the Archaean. Higher heat production in the crust and mantle influences the geodynamic process through which Archaean granites formed, either directly as a consequence of a higher crustal geotherm, or indirectly as a consequence of a different crustal and mantle composition, and stratification.

2.6.

DIAPIRIC PROCESSES IN THE FORMATION OF ARCHAEAN CONTINENTAL CRUST, EAST PILBARA GRANITE-GREENSTONE TERRANE, AUSTRALIA

A.H. HICKMAN AND M.J. VAN KRANENDONK Introduction

Many of the world's Archaean granite-greenstone terranes are characterised by regionalscale "dome-and-basin" patterns in which 30-150 km diameter, circular or ovoid domes are separated by irregularly shaped synclines, or keel-like structures (Macgregor, 1951). The domes typically expose cores of granitoid rocks and orthogneiss, whereas the synclines are developed in supracrustal successions of low-grade metamorphosed volcanic and sedimentary rocks ("greenstones"). Archaean terranes with this type of structure include the Zimbabwe and Kaapvaal cratons of southern Africa, parts of the Pilbara and Yilgarn cratons of Western Australia, the Dharwar craton of India, and parts of the Superior craton in Canada. A similar pattern is also developed in some Archaean and Proterozoic terranes, such as the Quadrilatero Ferrifero portion of the Brazilian shield (Marshak et al., 1997; Hippert and Davis, 2000). The prevalence of this type of structure in the Archaean terranes implies an important relationship to the generation of Archaean continental crust, but the origin of the pattern is still highly controversial. The 3.52-2.83 Ga East Pilbara granite-greenstone terrane (EP) in the Pilbara craton of Western Australia provides one of the world's best examples of an Archaean domeand-basin pattern. After cratonisation at c. 2.83 Ga, the EP has not been subjected to any significant deformation or metamorphism, and is therefore an ideal area to study processes involved in the generation of Palaeoarchaean and Mesoarchaean continental crust. This contribution describes the geology of the EP, and explains the dome-and-basin pattern 7"hePrecambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

2.6. Diapiric Processes

119

as the result of gravity-driven deformation resulting from the overturn of low-density rocks (granitoids) that have been buried under denser rocks (greenstone cover). The tectonic model is broadly similar to diapiric models used in other dome-and-basin terranes (Macgregor, 1951; Huddleston, 1976; Drury, 1977; Stephansson, 1977; Fyson et al., 1978; Gorman et al., 1978; Glikson, 1979; Gee, 1979; Mareschal and West, 1980; Borradaile, 1982; Anhaeusser, 1984; Bouhallier et al., 1993; Jelsma et al., 1993; Chardon et al., 1996; Choukroune et al., 1997). As of the early 1980s "uniformitarian", Phanerozoic-style, plate tectonic models have been increasingly applied to dome-and-basin terranes, with the result that diapiric models are now more controversial. Some recent workers (Zegers et al., 1996, 2001; Blewett, 2000, 2002; Kloppenburg et al., 2001; see also section 2.5 above) have argued that horizontal tectonic processes were also important in the early crustal evolution of the EP, although this has been challenged in other recent publications (e.g., Van Kranendonk et al., 2002, in press). The extensive literature on Archaean terranes provides good evidence that apparently similar dome-and-basin patterns formed by various tectonic processes. Certainly, some types of dome-and-basin pattern can be generated during horizontal deformation (e.g., Chardon et al., 2002), but in this contribution we contend that a wide range of criteria are available to distinguish these from EP-type dome-and-basin patterns that resulted from entirely diapiric deformation. These criteria provide compelling evidence that (a) the diapiric tectonic model best explains the crustal evolution of the EP, and (b) that diapirism is likely to have been an important process in the generation of Archaean continental crust. East Pilbara Dome-and-Basin Pattern

The dome-and-basin pattern of the east Pilbara (Fig. 2.6-1) was originally attributed to cross-folding (Noldart and Wyatt, 1962). Hickman (1975, 1983, 1984) rejected this interpretation because geological remapping of the east Pilbara in the 1970s revealed no evidence of regional-scale fold interference. His alternative diapiric model for the crustal evolution of the EP (Hickman, 1983) was based on structural and stratigraphic evidence, with the timing of events being constrained by the available geochronology. Hickman (1984) observed that most of the domes are circular or ovoid in plan, rimmed by steeply concentric tectonic foliations, and encircled by steeply dipping ring faults and shear zones with dome-side-up movement. The intervening greenstone synclines have no prevailing regional orientation, are locally cuspate or star-shaped in plan, and contain a dominant penetrative foliation that instead of being parallel to the axial planes of the synclines (horizontal compression) is arcuate, and parallel to the broadly concentric tectonic foliation of the granitoid domes (Hickman, 1983, p. 165). Hickman (1984) concluded that the EP domes and synclines were formed by successive events of solid-state diapiric deformation, magmatism, erosion and deposition over a period of approximately 600 million years. Recent detailed mapping (Van Kranendonk et al., 2002) indicates that the EP domeand-basin pattern is formed by a cluster of fault-bounded domes (Fig. 2.6-1 ) in which each dome consists of a granitoid core and an attached greenstone envelope. The greenstone "synclines" are formed from the combined greenstone envelopes of adjacent domes, and

120

Chapter 2: Generation of Continental Crust

in all greenstone belts these inwards-facing (younging towards the centres of the greenstone belts) autochthonous packages are separated by subvertical faults along syncline axes (Van Kranendonk, 1998). Hickman (2001) referred to these as axial faults, some of which occur in pairs to form narrow, deep grabens containing the youngest greenstones. The axial faults locally separate completely different stratigraphic levels of the greenstone succession, and can involve relative vertical movements of up to 10 km. Diapiric Doming

Diapiric models for the crustal evolution of the EP are based on the structural geology, stratigraphy, geochemistry, and geochronology of the terrane. Diapirism is primarily a gravitational response to the development of an inverted density profile in the upper crust, but may only be possible if accompanied by thermal weakening of the mid- to lower crust (Collins et al., 1998; Weinberg and Sandiford, 2001). Evidence for both these conditions is provided in the stratigraphy and tectonothermal history of the EP. Stratigraphy

The supracrustal succession of the EP comprises five groups assigned to the c. 3.522.94 Ga Pilbara Supergroup (Tables 2.6-1, 2.6-2), which is a thick succession of metamorphosed (mainly greenschist facies) volcanic and sedimentary rocks (Hickman, 1983). The maximum preserved thickness of the Pilbara Supergroup in any single area is 15-20 km, although this is less than the depositional thickness because basal stratigraphy was invariably excised by granitoid intrusion, and nonconformities mark local erosion of some parts of the succession. Variable stratigraphic thicknesses between and within greenstone belts (Hickman, 1983, his plate 2) are the consequence of successive events of tectonothermal activity and erosion over the 600 My depositional history of the succession. The Pilbara Supergroup contains erosional nonconformities (mostly local) at c. 3.46 Ga (Apex Basalt on Duffer Formation), c. 3.435 Ga (Panorama Formation on Duffer Formation), c. 3.425 Ga (beneath the Strelley Pool Chert), c. 3.325 Ga (beneath the Wyman Formation), c. 3.315-3.308 Ga (Budjan Creek Formation on Wyman Formation),

Fig. 2.6-1. Geological map of the East Pilbara granite-greenstone terrane (EP), showing generalised stratigraphy, way-up evidence, granitoid suites, and major structures. The Kurrana terrane, Mosquito Creek basin, and the Central Pilbara tectonic zone are separate terranes of the Pilbara craton. The dome-and-basin pattern of the EP is made up of randomly distributed domes, most of which contain a granitoid core (complexes of granitoid suites spanning up to 600 My) and a steeply dipping greenstone (Pilbara Supergroup) envelope. The domes are separated by major faults in the axial regions of greenstone synclines. Abbreviations: Granitoid complexes and domes: CA, Carlindi; CD, Corunna Downs; M, Mount Edgar; MU, Muccan; SH, Shaw; WA, Warrawagine; Y, Yilgalong; YU, Yule. Granitoid domes: TA, Tambourah. Greenstone domes: MP, McPhee; NP, North Pole; S, Strelley Granite.

2.6. Diapiric Processes

Granitoids

Pilbara Supergroup

~ - - ~ 2.852.83 Ga ~-~

2.952.93Ga

~~

3.263.24 Ga

~

3.32 3.30 Ga

~3.49

Fig. 2.6-1.

121

3.41 Ga

Gorge Creekand De GreyGroups 3.23 2.94 Ga --'~ Coonterunah,Warrawoona,and Sulphur Springs Groups3.52 3.24 Ga

Structure ----~ LallaRookhWesternShaw StructuralCorridor Geologicalboundary Fault or shearzone Way up

Table 2.6-1. Stratigraphy, and history of granitoid intrusion and deformation in the East Pilbara granite-greenstone terrane from 3.52 to 3.24 Ga Group Formation Age ( ~ a ) Sulphur Springs 3.24 Kangaroo Caves

3.26

Thickness (km)

Lithology

Granitoid intrusion

Deformation

0- 1.5

Basalt-andesite to rhyolite, and chert

3.25-3.24 Ga intrusion of monzogranite in the centres of some domes, and above the level of the Warrawoona Group

Doming and NE-SW trending rift systems developed in the westem half of the terrane

Kunagunarrina

0-2.4

Leilira

0- 1.0

Komatiite and komatiitic basalt Wacke and felsic volcanic rocks Voluminous 3.32-3.30 Ga granodiorite and monzogranite, and minor tonalite intruded into centres of rising domes, and into upper levels of Warrawoona Group

Major diapiric deformation, with resulting local erosion and nonconformities

Unconformity 3.31 Budjan Creek

0- 1.5

Regional unconformity Warrawoona 3.32 Charteris Basalt

0-2.0

3.32

0- 1.0

> 3.34

Wyman Euro Basalt

Conglomerate, sandstone, shale and felsic volcanic rocks

40, ;E:

h,

2.0-9.0

Komatiitic basalt and tholeiite Rhyolite, minor tuff and sandstone Komatiite, komatiitic basalt and tholeiite

$

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3

9

9 $. n

5

-. Table 2.6- 1 (continued). c. 3.42

2.

Strelley Pool Chert Panorama

0.1

Apex Basalt

0-3.0

3.46

Towers

0-0.5

3.47 3.47

Duffer Mount Ada Basalt

0-5.0 2.0

3.48

McPhee

0.1

c. 3.49

North Star Basalt

2.0

c. 3.49

Dresser

0- 1.0

c. 3.45

0.1 - 1.0

Chert, minor carbonate rocks and sandstone Felsic volcanic rocks Komatiite, komatiitic basalt and tholeiite Chert, komatiitic basalt and tholeiite Andesite to rhyolite Komatiitic basalt and tholeiite Carbonated ultramafic lava and chert Tholeiite, minor komatiitic basalt and komatiite Chert and tholeiite

3.49-3.41 Ga intrusion of 7TG sheets and laccoliths beneath the Warrawoona Group. Local penetration of the supracrustal succession forming felsic volcanic centres of Duffer and Panorama Formations

b

a2

Local extension and erosion

Local diapiric deformation and erosion

Local extension Local deformation

Local unconformity Coonterunah Double Bar

0- 1.9

3.52

Coucal

0- 1.4

Table Top

0-3.5

Tholeiite, minor komatiitic basalt Basalt, felsic volcanic rocks, and chert Tholeiite, minor komatiitic basalt

@

Table 2.6-2. Stratigraphy, and history of granitoid intrusion and deformation in the East Pilbara granite-greenstone terrane from 3.24 to 2.85 Ga Group Age (Ga) c. 2.85 De Grey c. 2.95

< 3.05

Formation

Thickness (km)

Lithology

Granitoid intrusion

Deformation

Post-tectonic syenogranite and monzogranite Lalla Rookh Sandstone and Cooragoora

1.8-3.0

Sandstone and conglomerate

Cattle Well

0-2.5

Sandstone, wacke, shale and felsic tuff

0.1

Shale and banded ironformation Komatiitic basalt and tholeiite Sandstone, wacke, siltstone and shale Banded iron-formation, chert and shale Sandstone and shale Banded iron-formation, chert and shale

Regional unconformity Gorge Creek Pyramid Hill Honeyeater Basalt and Coonieena Basalt Cundaline Paddy Market and Nimingarra Iron Corboy Pincunah Hill

1-1.5 0-0.8 1 .O

0- 1.5 1.O

Disconformity, and local unconformity sulohur Sorings Grow. Warrawoona Grow. or ore-3.24 Ga granitoids > 3.24 Ga

2.95-2.93 Ga potassic granitoids widely intruded into domes in the western half of terrane

Major phase of diapiric deformation resulting in deep erosion of rising domes. Strike-slip faulting in western half of terrane

Diapiric deformation and erosion of Gorge Creek group

2.6. Diapiric Processes

125

c. 3.235-3.200 Ga (beneath the Gorge Creek Group), and c. 3.05-2.94 Ga (regional unconformity beneath the De Grey Group) (Tables 2.6-1,2.6-2). The two most extensive nonconformities occur above the Warrawoona Group and beneath the De Grey Group (Hickman, 1990), which are regional breaks, and coincide with major tectonothermal events involving deformation, granitoid intrusion, and metamorphism (Tables 2.6-1, 2.6-2). In most parts of the EP, the Gorge Creek Group unconformity directly overlies the Warrawoona Group, but the local intervention of the c. 3.280-3.235 Ga Sulphur Springs Group, provides evidence of another unconformity that was related to a tectonothermal event at c. 3.24 Ga (Van Kranendonk et al., 2002; Huston et al., 2002). The ages of all the nonconformities closely coincide with periods of felsic volcanism and granitoid intrusion (Tables 2.6-1, 2.6-2). Several workers have suggested that parts of the Pilbara Supergroup succession were tectonically thickened by subhorizontal thrusting and recumbent folding. Bickle et al. (1980, 1985) and Bettenay et al. (1981) interpreted early Alpine-style thrusting to have caused overthickening of the crust leading to density instability and subsequent solid-state diapirism. Krapez (1993) and van Haaften and White (1998) reinterpreted U-Pb zircon data (Thorpe et al., 1992a; McNaughton et al., 1993) from the lower Warrawoona Group east of the Mount Edgar Granitoid Complex (Fig. 2.6-1) to argue that c. 3.30 Ga units are tectonically interleaved with c. 3.47 Ga units, and that this belt is therefore a litho-tectonic complex rather than the normal autochthonous succession. Subsequent, more extensive SHRIMP U-Pb zircon geochronology (Nelson, 1999, 2000) indicates that all the volcanic rocks in this belt are older than 3.45 Ga, and that isotopic ages decrease progressively upwards through the succession according to the original stratigraphic interpretation by Hickman (1983). Recent detailed geological mapping of the EP (Van Kranendonk et al., 2002), supported by SHRIMP U-Pb zircon geochronology, has revealed no stratigraphic repetitions by subhorizontal thrusting or recumbent folding in any of the greenstone belts. The detailed mapping has supported the interpretation that the Coonterunah and Warrawoona Groups form an upwards-facing autochthonous succession, and that this is locally over 15 km thick. The composition and geochronology of this succession (Fig. 2.6-2) indicates that the lower Pilbara Supergroup (Table 2.6-1) was constructed through repeated ultramafic-mafic-felsic volcanic cycles of 11-37 My duration (Van Kranendonk et al., 2002), consistent with derivation from eight successive mantle plume events (Fig. 2.6-2). Arndt et al. (2001) also favoured volcanism related to mantle plumes (sections 3.2 and 3.3), suggesting that the lower Warrawoona Group was probably erupted in an oceanic plateau setting. These authors commented that the geochemistry of upper Warrawoona Group (data from the Euro Basalt west of the Corunna Downs Granitoid Complex) indicates some interaction with continental crust, and that evidence of crustal contamination increases upwards through the Pilbara Supergroup. Other geochemical data and inherited zircon data (Gruau et al., 1987; Thorpe et al., 1992b; Bickle et al., 1993; Nelson, 1998b, 1999, 2000, 2001a; Green et al., 2000) indicate some crustal contamination at all levels of the CoonterunahWarrawoona Group succession, consistent with volcanism on a continental substrate. Van Kranendonk et al. (2002) suggested that melts were extracted from a continuously and vigorously convecting hot mantle, and that felsic magmas were generated by melting of mafic crust.

126

Chapter 2: Generation of Continental Crust

Duration of cycle (My)

Age (Ga)

Schematic stratigraphic section

Felsic and sedimentary formations

Basaltic formations

>~ Group cz "o ::z: >

Illlllllllllllllllllll KANGAROOCAVESFM. Ma

I~~~13240 15-

2 [-.- ".V. " ' : - Cycle7

c.3255 Ma

N ~0

BUDJAN CK. FM. 3308 Ma

"~ <_11

_

WYMAN FM 3325 3319 Ma Ivvvvvvvt ~.y.v.y.v.vy.] 33463335 Ma

z27 Cycle 6

rvvvvvv,4 rvvvvvv,,4 rvvvvvv,vl rvvvvvv,,4

2

[VVVVVVV] iv ~ - ~ v - ~ v IV V V V V V V IV v v v v v v

Approximate 1JKm verticalscale | throughout / column

I,V,,v,,v,,v,,v,,v,,v,

o-

?

- ~ =-/" "~"

1

STRELLEYPOOLCHERT

r/P111MMMI11111PANORAMAFM. /~

Cycle 5

<-37-

3.4

/

Cycle 4

14 -

~, ~ V V V ~,~/t 34583426Ma Vvvvvvv'I Vvvvvvv'l ~/\/ \/ \/__\~\/x/ ,] TOWERSFM. ~ DUFFERFM. 3474 3463 Ma

1.---4 => >

McPHEEFM. ,~,~,~,~,~,~, ~ 3477Ma

Cycle3

t 1919-

3.5-

il

vvv~vvvl V V ',,,,"'vv v I c.3490Ma VVVVVV~] i 1111111111111111 ~' 1~/~/~/9 ~/~/"1 3496Ma vvvvvvl

f

3508 Ma COUCALFM. 3515 Ma

B

Cycle 1

~ a . - 1-

>--15m

Ultramaficlavas

~TTT~ Felsicvolcanics

~

Felsicvolcaniclastics

Basalt

~

~

Coarseclastics

Chert

~ ~

Granite Unconformity

Fig. 2.6-2. Generalised stratigraphic column of the lower Pilbara Supergroup, showing eight ultramafic-felsic volcanic cycles between 3.52 and 3.24 Ga. Note that clastic sedimentary rocks are rare in the succession except above the first regional unconformity caused by erosion of the c. 3.315 Ga domes.

2.6. Diapiric Processes

127

In the diapiric model (e.g., Hickman, 1984; Van Kranendonk et al., 2002, in press) the EP domes developed episodically, but progressively, throughout the deposition of the Pilbara Supergroup. Doming commenced at c. 3.46 Ga beneath felsic volcanic centres (Duffer and Panorama Formations, Table 2.6-1) that developed where intrusion of synvolcanic granitoids (generally as sheets beneath the Warrawoona Group) had locally penetrated and removed part of the upper greenstone layer (Collins, 1989). Evidence supporting this interpretation is provided by sedimentary facies and palaeocurrent data indicating that felsic volcanic centres of the Duffer and Panorama Formations were located in the positions of subsequent granitoid domes (DiMarco and Lowe, 1989). The second doming event, which was demonstrably diapiric and of regional importance, occurred at c. 3.315 Ga. During this event, the early EP domes were amplified by extensive granitoid intrusion (note the widespread distribution of 3.32-3.30 Ga granitoids in Fig. 2.6-1). After 3.315 Ga almost all granitoid intrusion was through the domes (see below), causing further amplification of these structures and uplift of overlying greenstones. A regional tectonothermal event at c. 3.24 Ga, interpreted to be related to a mantle plume (Van Kranendonk et al., 2002), was accompanied by widespread granitoid intrusion (see below), and in the western half of the EP, rifting, deposition of the Sulphur Springs Group, a third doming event, and subsequent widespread erosion. Erosion following the second and third doming events resulted in the major unconformity between the Warrawoona and Gorge Creek Groups. Doming at c. 3.315 and c. 3.24 Ga was followed by dome erosion and by deposition of the Gorge Creek Group in intervening basins. This interpretation implies that stratigraphic units younger than the Warrawoona Group may never have been deposited across the crests of the 3.315 Ga domes, although more palaeocurrent data from the Gorge Creek Group is required to support this conclusion. In the northeastern part of the EP, the c. 3.20-3.00 Ga Nimingarra Iron Formation of the Gorge Creek Group (Table 2.6-2), unconformably overlies deeply eroded granitoids on the northern limb of the Muccan dome (Fig. 2.6-1; Dawes et al., 1995). Hickman (1984) interpreted this northeastern part of the EP as a zone of major early uplift and erosion between deposition of the Warrawoona and Gorge Creek Groups. This concentration of deposition (including dense units of iron-formation and mafic volcanic and intrusive rocks) in the basins between the domes, combined with continuing erosion of the domes, must have accentuated lateral density heterogeneities in the upper crust, increasing the potential for further gravitational sinking of the basins and diapiric uplift of the domes. Therefore some amplification of the domes probably occurred during deposition of the Gorge Creek Group. A fourth doming event at c. 2.95 Ga coincided with a major tectonothermal event on the northwestern margin of the EP (Smithies and Champion, 2000). This doming was accompanied by sinking and synclinal folding of the Gorge Creek Group basins (immediately above the 3.315 Ga synclines in the underlying Warrawoona Group), and resulted in much deeper erosion of the domal granitoid complexes. Thick clastic sedimentary formations of the De Grey Group were deposited between most of the domes, but subsequent erosion has removed evidence for all but four of these basins (Fig. 2.6-1). Strike-slip faulting in the Lalla Rookh-Western Shaw Structural Corridor (Fig. 2.6-1) at c. 2.94 Ga was associated

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Chapter 2: Generation of Continental Crust

with synkinematic deposition and upright folding and faulting of the De Grey Group (Van Kranendonk and Collins, 1998).

Granitoid Complexes The granitoid complexes in the cores of the EP domes are composed of numerous granitoid intrusions and variable amounts of migmatised older gneissic components (Fig. 2.6-1; Hickman, 1983; Collins, 1989; Williams and Collins, 1990). Hickman (1984)observed that the concentric tectonic foliation in some granitoid cores crosses granite-granite contacts within the complexes, establishing that doming was partly later than granitoid intrusion. Recent SHRIMP U-Pb zircon geochronology (e.g. Van Kranendonk et al., 2002) indicates that granitoid intrusion closely coincided with deformation events over a period of at least 550 My. The granitoid complexes of the EP can be divided into three main types. In the western EP, type 1 is exemplified by the Yule Granitoid Complex (Fig. 2.6-1), a broad (c. 120 km diameter), younger magmatic complex composed chiefly of 2.95-2.83 Ga granitoids that locally intrude the highest stratigraphic components of the greenstone succession (De Grey Group). These granitoids consist predominantly of monzogranite and syenogranite generated by partial melting of older granitoid crust (Champion and Smithies, 2000). Type 2 granitoid complexes, best represented by the Shaw, Mount Edgar, and Muccan Granitoid Complexes, are steep-sided structural domes (50-70 km diameter) composed mainly of 3.49-3.24 Ga granitoids, with minor younger granitoids. The intrusive contacts of type 2 complexes do not penetrate above the lower Warrawoona Group, and many of these contacts are strongly sheared. The granitoids of type 2 complexes consist of 3.49-3.41 Ga tonalite-trondhjemite-granodiorite (TTG), 3.32-3.30 Ga granodiorite to monzogranite, and 3.25-3.24 Ga monzogranite. Volumetrically minor components of the Shaw and Mount Edgar Granitoid Complexes are sheets of highly fractionated 2.85-2.83 Ga monzogranite. The western part of the Carlindi Granitoid Complex has the features of a type 1 complex, whereas its eastern half more closely resembles type 2. A third type of complex is seen in the Corunna Downs Granitoid Complex which consists almost entirely of low strain 3.32-3.30 Ga granodiorites and monzogranites that have intruded up into the upper part of the Warrawoona Group. Greenstones around the granitoid core (35 km diameter) of the Corunna Downs Dome dip steeply outwards, but are not strongly sheared, consistent with the present exposure of the dome being at a higher structural level (closer to the crest) than in type 2 complexes. This interpretation is supported by the presence of mushroomshaped folds in the surrounding greenstone belts, similar to those around natural salt diapirs (Van Kranendonk et al., in press). The Warrawagine and Yilgalong granitoid complexes are partly concealed by cover sequences, including the Hamersley basin, but they closely resemble type 2 complexes. As suggested above, the differences between type 2 and 3 granitoid complexes are interpreted to be different depths of erosion and different amounts of granitoid intrusion and uplift. Large intrusions of the younger, more potassic granitoids of type 1 complexes are mainly restricted to the western part of the EP, reflecting the prox-

2.6. Diapiric Processes

129

imity of these complexes to c. 2.95 Ga thermotectonism on the northwestern margin of the EP. The 3.49-3.41 Ga TTG suite is silicic and sodic, and was derived through high-pressure melting of a mafic crustal source (Smithies, 2000; Champion and Smithies, 2001). Bickle et al. (1993) showed that there was a component of > 3.6 Ga crustal material (including felsic crust) in the suite, although this does not necessarily imply that this older crust was the source. The TTG suite does not penetrate above the Panorama Formation, and typically has contacts parallel with the lowermost formations of the greenstone belt succession. This suggests intrusion as horizontal sheets and laccoliths beneath the greenstone belt prior to the first major regional doming at c. 3.315 Ga. The 3.32-3.30 Ga suite typically penetrates the upper levels of the Warrawoona Group and occupies large volumes of the eastern granitoid complexes (Fig. 2.6-1). We infer that the suite was intruded vertically through the rising domes, and leveled off in the crests of these sructures in the greenstone belt. Heat from this intrusion caused contact metamorphism, and migmatisation and ductile remobilisation of the 3.49-3.41 Ga TTG suite during doming. Collins (1993) showed that the 3.32-3.30 Ga and younger, more potassic suites were derived through progressive episodes of melting of the 3.49-3.41 Ga TTG suite. The 3.25-3.24 Ga suite intruded the central sections of the Mount Edgar, Muccan, Yule, and Carlindi Granitoid Complexes, and formed a subvolcanic laccolith (Strelley Granite) south of the Carlindi Granitoid Complex (Fig. 2.6-1). At 2.95 Ga, extension in the Central Pilbara Tectonic Zone immediately northwest of the EP, resulted in intrusion of high-temperature mantle melts (Smithies and Champion, 2000). The thermal anomaly generated by this event caused widespread melting of granitoid crust and amplification of the 120 km-diameter Yule Granitoid Complex, and the western part of the Carlindi Granitoid Complex. The 2.95-2.93 Ga suite intrudes the entire greenstone succession, leaving only isolated greenstone remnants. About 100 My later, at 2.85-2.83 Ga, an essentially post-tectonic suite of fractionated granitoids intruded four of the granitoid complexes (Fig. 2.6-1). The repeated intrusion of successive granitoid melts into the centres of domes, and solid-state reactivation of older granitoid phases accompanied by greenstone belt sinking (Collins et al., 1998), accentuated lateral density contrasts in the upper crust (originating with 3.49-3.41 Ga magmatism) and promoted diapiric amplification of the domes.

Structural Geology Domal granitoid complexes of the EP are separated by greenstone belts of varying widths, orientations and shapes (Fig. 2.6-1). Particularly striking is the abnormally wide exposure of supracrustal rocks referred to as the North Pole Dome, from which six greenstone belts radiate between six encircling domal granitoid complexes. The EP contains two greenstone domes, the North Pole Dome and the McPhee Dome, in which the belt envelope has been only partly eroded. Regional Bouger gravity data (Blewett et al., 2000) indicate that although the upper crust in these areas contains less supracrustal rocks than in the main synclines, this crust is significantly denser than beneath the granitoid complexes. Two more

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Chapter 2: Generation of Continental Crust

domal granitoid complexes (Warrawagine and Yilgalong) are present in the eastern EP, but are partly concealed by unconformable, post-2.77 Ga cover sequences. Most of the EP domes are broad, steep-sided structures, generally circular to oval in plan, and measuring over 50 km in diameter. Gravity data indicate that the type 2 and 3 granitoid complexes of the eastern EP are almost vertical cylinders to at least 14 km depth (Wellman, 2000), matching the depth predicted by scaling of numerical models of diapirs (Mareschal and West, 1980). The gravity data, the steepness of the marginal shear zones and ring faults (Figs. 2.6-1, 2.6-3), and the large vertical displacements (granite-side-up) on these faults (Collins and Van Kranendonk, 1999), indicate that much of the doming was by movement on these faults (Hickman, 1975, 1984). The same process has been proposed for the domes of the Quadrilatero Ferrifero in Brazil and the Penokean orogen in Michigan (Marshak et al., 1997). Differences in stratigraphy and metamorphic grade across the EP faults indicate several kilometres of vertical movement (Van Kranendonk et al., 2002). The granitoid domes have been eroded to different structural levels, reflecting differing amounts of cumulative upwards movement over the long history of their formation. The most deeply eroded domes are those exposing type 2 granitoid complexes characterised by strongly sheared granite-greenstone belt contacts, a steeply dipping tectonic foliation parallel to these contacts, and a relatively high proportion of pre-3.41 Ga granitoids. Domes with predominantly intrusive granite-greenstone contacts (containing types 1 and 3 granitoid complexes), and characterised by weak tectonic foliations, are exposed at higher structural levels. The Shaw and Mount Edgar Domes both possess sheared contacts in the south and intrusive contacts in the north, suggesting that uplift was greatest on their southern margins. Minor structures related to the domes include a tectonic foliation broadly parallel to granite-greenstone belt contacts (Fig. 2.6-3), steeply plunging mineral lineations that are radial off the domes (Fig. 2.6-3), and structures produced by vertical stretching that include elongate clasts (Hickman, 1984) and L-tectonite fabrics in diapiric triple points (Collins et al., 1998). The tectonic foliation is parallel to, and most strongly developed adjacent to, the marginal shear zones and ring faults, consistent with it being a penetrative shear strain fabric intimately related to doming. Exceptions occur where separate doming events, with slightly different centres of uplift, have resulted in non-parallel tectonic foliations (Collins, 1989). An excellent example of discordant foliations is present in the southwestern Mount Edgar granitoid complex where an early northeast striking foliation is abruptly truncated

Fig. 2.6-3. Simplified geology of the Mount Edgar and Corunna Downs domes, and the eastern margin of the Shaw dome, showing generalised trends of tectonic foliations in granitoids, shear zones, major faults, lineations, and way-up evidence in greenstones. Axial faults are visible in the centre of the Warrawoona syncline and in the syncline west of the Corunna Downs granitoid complex. Ring faults and shear zones occur close to granite-greenstone contacts and in greenstones on the limbs of the domes. Note the stratigraphic breaks across axial faults, the relatively young stratigraphic units in axial grabens, the hook-shaped folds on the northwestern and northeastern sides of the Corunna Downs dome, and the radial pattern of lineations around the Mount Edgar dome.

2.6. Diapiric Processes

Fig. 2.6-3.

131

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Chapter 2: Generation of Continental Crust

by a ring fault parallel to a shear zone (Fig. 2.6-3). The northeast striking foliation defines a broad synformal structure in the western half of the complex, and the presence of numerous volcanic enclaves in this synform suggests that it might underlie a completely eroded roof pendant of a synclinal greenstone belt (Hickman, 1975). Steeply plunging mineral lineations are radial on the flanks of the Mount Edgar Dome (Fig. 2.6-3), and indicate opposite senses of shear on opposite sides of the dome. This is a critical feature of diapirism, in which stretching is down-dip and multidirectional on a regional scale, and contrasts with stretching under horizontal extension that results in uniaxial stretching lineations with the same sense of movement across domes on a regional scale (e.g., as in metamorphic core complexes). In the EP, greenstone belt synclines ("basins" of the dome-and-basin pattern) are not simple fold structures, but contain axial faults that separate inwards-facing (younging away from the granitoid complexes) monoclinal supracrustal envelopes of adjacent domes (Figs. 2.6-1 and 2.6-3; Van Kranendonk, 1998; Hickman, 2001). Bedding in the axial regions of the synclines is typically vertical and sheared, with local development of steeply plunging L-tectonites, whereas away from these central areas bedding is generally less steeply inclined, less sheared and is associated with a subparallel tectonic foliation with variable development of lineations. Examples of these features are well documented from the Warrawoona Syncline (Collins and Teyssier, 1990; Collins et al., 1998). Most major and minor folds within the greenstone belts plunge towards syncline triple junctions, consistent with the expected geometrical relationship to the domes. However, it is emphasised that the EP domes are products of several periods of diapiric deformation and granitoid intrusion, and all are relatively complex asymmetric structures. Structures related to diapirism have locally been rotated by, or overprinted by, late-stage (c. 2.94-2.89 Ga) folds, faults, foliations, and lineations formed in response to horizontal compression (e.g., Hickman, 1983). The eastern margin of the Shaw Dome (Fig. 2.6-1 ) contains east-plunging lineations (Fig. 2.6-3), whereas late-stage strike-slip deformation of the Western ShawLalla Rookh Structural Corridor on the western margin (Fig. 2.6-1; Van Kranendonk et al., 2002) has obliterated almost all earlier structures, and is characterised by shallow-plunging lineations and sinistral kinematic indicators. West of this structural corridor, Hickman (1983, p. 169) recognised large-scale late-stage horizontal deformation of the Pilgangoora Syncline, with associated minor structures such as kink bands and conjugate folds. Van Kranendonk (2000) attributed these structures to c. 2.94 Ga regional northwest-southeast compression. Blewett (2002) recognised the same regional event, assigning it to D4c (2.935-2.880 Ga) in his structural scheme. Thus, careful structural analysis may locally be required to distinguish dome-related structures from late-stage structures, particularly in, and adjacent to, reactivated shear zones. Other Tectonic Models

We interpret the dome-and-basin pattern of the EP, and the crustal evolution of the EP between 3.52 and 2.83 Ga, as the result of a combination of repeated cycles of plumerelated volcanism, crustal heating, gravity-driven vertical deformation, dome erosion, and

2.6. Diapiric Processes

133

deposition between domes. This diapiric model explains the major features of the domeand-basin pattern as due to vertical tectonics, with limited late-stage modification by regional horizontal compression. Plume-related volcanism implies some extension of the lithosphere, and some syndepositional extensional faulting has been documented (Nijman et al., 1998b; see also section 2.7). These structures are local and small-scale, but other recently published tectonic interpretations have attributed the EP dome-and-basin pattern to regional crustal extension (e.g., Zegers et al., 1996) or compression (e.g., Blewett, 2002; see section 2.5). Zegers et al. (1996, 2001) and Kloppenburg et al. (2001) interpreted the tectonic development of the Shaw, Mount Edgar and Corunna Downs Domes using a metamorphic core complex model. Blewett (2000, 2002) attributed the dome-and-basin pattern to complex interference between multiple phases of folds formed under regional horizontal tectonics, and structural amplification (steepening and deepening of cross-fold structures), possibly by diapirism. A detailed assessment of the numerous observations, interpretations, and arguments that were presented to support these alternative models was given in Van Kranendonk (2000) and Van Kranendonk et al. (2001, 2002, in press). A general assessment of the extent to which each of the three main types of tectonic models (diapirs, core complexes, and cross-folds) is consistent with EP geology can be made using the compilation of diagnostic criteria in Table 2.6-3. All the expected geological characteristics of the diapiric model are met by EP geology, whereas many features normally associated with core complexes and cross-folds are absent from the EP. Based on Phanerozoic examples, neither the core complex model nor the cross-fold model is compatible with the detailed geometry of the EP domes as described above. Additionally, the diapiric model provides the most complete explanation of the 600 My crustal evolution of the EP by inter-relating magmatism, deformation, erosion and deposition. Discussion a n d S u m m a r y

The first diapiric model applied to the EP (Hickman, 1975, 1983, 1984) interpreted diapirism as a progressive, mainly solid-state process that commenced at c. 3.45 Ga and culminated at 2.95 Ga. Collins et al. (1998), Collins and Van Kranendonk (1999), Van Kranendonk and Collins (2001), and Van Kranendonk et al. (in press) modified this model by recognition of c. 3.32-3.30 Ga magmatic diapirism, resulting from partial convective overturn (see also, section 3.4) of the upper and middle crust. In this model, domical granitoid complexes developed partly through ductile mobilisation of the early synvolcanic granitoid rocks (TTG suite), and partly through the emplacement of successive generations of plutonic suites (magmatic diapirs) into the cores of progressively evolving domes. Bickle et al. (1989) interpreted the 2.95-2.93 Ga granitoid suite to have been derived from partial melting of the older granitoids, and Collins (1993) suggested a similar origin for the 3.32-3.30 Ga suite. Doming is interpreted to have formed in response to sinking of dense, upper-crustal greenstone belts into a thermally-softened, more buoyant granitoid middle crust during episodes of partial convective overturn (Collins et al., 1998). Weinberg and Sandiford (2001 ) emphasised the importance of the 2 to 3 times higher heat production in

Table 2.6-3. Characteristic features of diapirs, core complexes, and cross-folds, noting presence in the East Pilbara granite-greenstone terrane Diavirs J ~ a n d o mdome-and-basin geometry; synclines have no preferred regional orientation

Metamorphic core complexes (MCC) Commonly a linear chain of domes that are elongate parallel to orogen boundary

Cross-folds Regular fold interference geometry of anticlines and synclines, typically semi-orthogonal

J ~ o m e scored by buoyant substrate (granitoids)

Domes cored by deformed, metamorphosed sialic basement of various protoliths of various ages (= thrust-accretion assemblage) Shallow synclines of largely undeformed cover rocks and typically including large volumes of synkinematic clastic sedimentary rocks Domes are simple, generally ovoid structures with shallow-dipping margins (< 30°), commonly half-domes (brachidomes)

Anticlines cored by structurally lowest components of folded sequence Synclines cored by structurally highest components of folded sequence

JDeep, steep synclines occupied by dense overburden (greenstones) and syndiapiric clastic sediments JDomes most commonly have simple, generally ovoid plan outlines with steeply dipping margins, but chaotic internal fold geometry; may form isolated domes or interconnected bodies depending on level of erosion J I ~cross-section, domes vary from symmetric plugs or symmetric mushrooms with single or double fold flaps to highly asymmetric bodies with, or without, fold flaps. JDomes may contain synclinal roof pendants of cover rocks

Domes generally shallow dipping, broad structures and commonly asymmetric in crosssection, with one shallow and one more steeply dipping limb Domes may be partially covered by klippen of cover rocks

JDiameter 3-120 km (salt diapirs; 1-20 km)

JDiameter 4-170 km

JDiameter up to 100 km

J1n plan view, domes contain circular shear zones (ring faults) around granitoid cores, with granite-side-up kinematics; shear zones commonly (but not exclusively) at granitegreenstone contacts. Shear zones transpose stratigraphy, but do not significantly alter it

Flat- to moderate-dipping, younger-on-older faults with upsection transition from ductile to brittle cataclasis. Main, shallow-dipping extensional decollement separates ductilely deformed metamorphic core below from cover affected by brittle deformation. Decollement typically best

Faults are associated with separate sets of folds typically show crosscutting relationships

Anticline crests are generally simple ovoids with predictive changes in minor fold asymmetry (M, S and Z shapes) across hinge zones Sine wave cross-sectional geometry

-

W

P

Table 2.6-3 (continued). Diapirs Shears commence as shallow structures and rotate into steep structures

JLocal development of reverse (greenstoneside-up) faults in greenstone synclines (cf. Dixon and Summers, 1983) Identified in EP: near vertical axial faults in syncline hinges associated with high strain; locally in pairs bounding central grabens; major vertical movement between supracrustal envelopes of adjacent domes J ~ o m e s characterised by strong steep foliations on steeply dipping margins, chaotic folds in core, and weak flat foliations in crest

Jsteep synclines will contain dome-parallel foliations and plunging lineations that point in towards zones of maximum sinking in extreme cases. Contact between core and envelope in granite-greenstone terranes is commonly a mixed transitional zone of granitoid injection and greenstone belts rafting; no abrupt change in metamorphic grade except across major faults; structural fabric shared bv core and envelo~e

Metamorphic core complexes (MCC) developed on shallower limb and characteristically a microbreccia. Extensional decollement may juxtapose youngest part of cover against metamorphic basement, through excision of kilometres-thick sections of stratigraphy. Decollement starts as high-angle normal fault that rotates into shallow dips Commonly gentle synclineshasins occupied by steeply dipping cover rocks deformed by sets of listric normal faults, forming a megabreccia

J(Local) Possible development of reverse shear zones (i.e., greenstoneside-up) along contacts between rheologically distinct units during flexural slip

Domes characterised by flat foliations throughout (rarely > 30°),decreasing in intensity downwards, away from the decollement; if tightened by later compression, domes will display a simple, dome-parallel foliation pattern

Depending on folding mechanism, anticlines will most commonly show sets of cross-cutting subvertical foliations, but may contain shallow foliations in upper parts of more competent members and downwards-fanning steep foliations in pinched cusps of folded units

Shallow synclines contain undeformed cover rocks and synkinematic clastic sedimentary rocks deformed by listric normal faults. Sharp lithological, structural and metamorphic discontinuity at contact between metamorphic core (commonly high-pressure) and lower-grade cover (commonly unmetamorphosed); structural fabrics not shared by two domains

Depending on folding mechanism, synclines will most commonly show sets of cross-cutting subvertical foliations, but may contain shallow foliations in lower parts of more competent members and upwards-fanning steep foliations in pinched cusps of folded units

-

LA cn

Table 2.6-3 (conrin~ted). Diapirs J ~ a d i a t i n lineations ~ may occur around symmetrical domes, whereas asymmetric domes will show asymmetric lineation patterns

Metamorphic core complexes (MCC) Lineations commonly aligned along one trend, even in multiple domes: partially radiating pattern locally developed in Goodenough Dome, D'Entrecasteaux Islands

Cross-folds Usually the first fold set may contain a set of lineations parallel to fold axes, and these will be folded by the second generation folds producing unidirectional, but doubly-plunging fold axes and lineations

J ~ t r a i nwill vary within and around domes, including areas of weak vertical flattening above dome crests and high strain zones of pure vertical stretching in zones of diapiric uplift (core of domes) and greenstone sinking (core of synclines): progressive changes in strain ellipsoid from S r L, L = S, to pure L-tectonites; solidstate granitic diapirs will have isotropic cores and sheared margins J ~ i n o rfold asymmetry on dome margins varies from S to Z shaped in left to right cross-sectional view, opposite to that in folds

Strain most intense in decollement, characterised by intense vertical flattening, unidirectional mineral elongation lineations (i.e., S >> L), and kinematic evidence for a unidirectional vergence of the cover across domal crest

Depending on the buckling mechanism, strain will vary in predictive fashion across folds

No synextensional folds in cover sequence. Previous episode(s) of nappe stacking.

Changes in minor fold asymmetry from Z, to M, to S shape in left to right cross-sectional view across anticline hinges

J ~ o m e smay contain multiple phases of synkinematic material locally intruding cover rocks; magmatism during doming produces characteristic granites-in-granites map pattern Identified in EP: synkinematic granite magmatism associated with each pulse of uplift; progressively more highly fractionated granites

Domes may contain multiple phases of synkinematic granitoid rock (including S-type), emplaced as sill-like masses below decollement, but never penetrate into greenstone belt cover Greek MCC: predominant subalkaline basaltandesite-dacite series, minor shoshonite, alkali basalt, and MORB

Amagmatic; may be preceded or followed by granitoid magmatism

%3 $

-

h,

2 3

". Amagmatic

6

3

z

m 2

2

Y

P

P

3. $.

Table 2.6-3 (continued). Diapirs

Metamorphic core complexes (MCC)

Cross-folds

Jvariety of metamorphic p-T-t paths depending on position around, and within, rising hot diapir; thermal anomaly associated with dome produces contact metamorphic pattern JGenerally prolonged formation, ongoing with clastic sedimentation

Isothermal decompression paths in basement complex and common very high pressure (blueschists and eclogites); generally unmetamorphosed cover rocks Generally rapid formation: 4 My for D'Entrecasteaux Islands 14-30 My for Cordilleran MCC 2 6 3 5 My for Sifnos-Syros, Greece

Folded isobars, with highest-grade exposed in anticline crests

J ~ a t u r a lexamples formed in a variety of tectonic environments, but can occur in the absence of plate tectonics

Formed in active extension, following collisional orogeny and therefore deform earlier thrustslnappes. Cordilleran MCC in the hinterland of an orogen, behind the foreland foldthrust belt Prior stratigraphy thinned by exhumation of core complexes; Jsynextensional deposition of unmetamoruhosed cover in normal fault-bound basins

Form during two (or more) compressional deformation episodes

Jstratigraphy controlled by rising diapirs; interplay between doming and sedimentation

Deformational episodes may be separated by 1-2 My or by eons

Folded stratigraphy; possible clastic deposition in associated intermontane or foreland basins

References: Ramberg (1967); Schwerdtner and Clark (1967); West and Mareschal (1979); Coney (1980); Davis (1980); Mareschal and West (1980); Dixon and Summers (1983); Anhaeusser (1984); Hickman (1984): Schwerdtner and Van Kranendonk (1984); van Berkel et al. (1984); Ramsay and Huber (1987); Suarez et al. (1987); Chen et al. (1990); M.P.A. Jackson et al. (1990); E.J. Hill et al. (1992); Weinberg and Schmelling (1992); Baldwin et al. (1993); Bouhallier et al. (1993); Hill and Baldwin (1993); Jelsma et al. (1993); Brun et al. (1994); Kisters and Anhaeusser (1995); Weinberg and Podladchikov (1995); Chardon et al. (1996, 1998); Jolivet et al. (1996); Ratcliff and Weber (1997); Pe-Piper (1998); Collins and Van Kranendonk (1999); Alsop et al. (2000); Trotet et al. (2001); Van Kranendonk et al. (2002, in press). J ~ e n o t e sfeature present in EP.

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the Archaean (see also, section 3.6) crust due to radioactive decay of K, U, and Th. This implies that Archaean continental crust was possibly much weaker than modern continental crust, and more susceptible to gravity deformation. From 3.52 Ga onwards, the Pilbara Supergroup was deposited on a continental plateau, or possibly an oceanic-type plateau adjacent to continental crust. From c. 3.46 Ga, volcanic formations of the upper Warrawoona Group and higher stratigraphic units were deposited on a thick crust of granitoids (TTG) and greenstone belts that had been locally deformed and subsequently eroded. Five autochthonous groups, deposited over the period 3.52-2.94 Ga, are separated by erosional nonconformities, and change from being almost entirely volcanic in the lower three groups to become largely sedimentary in the upper two groups. The lower groups are interpreted to be the products of a sequence of mantle plume events, whereas deposition of clastic sedimentary rocks in the upper two groups was largely a consequence of the erosion of progressively rising domes. Diapiric doming in the EP was a response to differential loading of the dominantly granitoid crust as thick successions of mafic and ultramafic volcanic rocks were deposited in the Coonterunah and Warrawoona Groups. Gravitational instability increased progressively as this dense supracrustal succession became thicker (locally between 15 and 20 km thick), and as more TTG sheets and laccoliths were intruded beneath the early greenstones (Hickman, 1984; Van Kranendonk et al., 2002, in press). Uneven emplacement of granitoids and eruption of felsic volcanic rocks instigated lateral variations in lithology (density) and thickness in both the supracrustal rocks and the mid-crustal sheeted sill complex. Synvolcanic doming commenced at this time (DiMarco and Lowe, 1989). Doming of the granitoid layer was synchronous with adjacent sinking of the greenstones into a thermally softened, more buoyant granitoid middle crust. Positioning of the domes appears to have closely coincided with the distribution of felsic volcanic centres. In the first major diapiric event, at c. 3.32-3.30 Ga, partial convective overturn of the upper and middle crust was accompanied by magmatic diapirism across at least six of the eight granitoid complexes (see distribution of 3.32-3.30 Ga granitoids on Fig. 2.6-1). A plume-related tectonothermal event at 3.24 Ga resulted in widespread granitoid intrusion and doming, particularly in the western half of the EP. During the final major diapiric event at c. 2.95 Ga, diapirism in the eastern half of the EP was accompanied by far more limited granitoid intrusion, and doming was essentially solid-state. In the western EP, the 2.95 Ga doming was associated with voluminous 2.95-2.93 Ga magmatic granitoid intrusion (Fig. 2.6-1) related to pulses of crustal extension and contraction on the northwestern margin of the EP (Smithies and Champion, 2000). EP granitoid rocks were emplaced into the Pilbara Supergroup during five main episodes at c. 3.49-3.41 Ga, 3.32-3.30 Ga, 3.25-3.24 Ga, 2.95-2.93 Ga, and 2.85-2.83 Ga. The 3.49-3.41 Ga TTG suite is silicic and sodic and was derived through high-pressure melting of underlying mafic crust. The younger, more potassic suites were derived through progressive episodes of melting of the 3.49-3.41 Ga TTG suite. At 2.95 Ga, extension in the Central Pilbara Tectonic Zone immediately northwest of the EP resulted in intrusion of high temperature mantle melts (Smithies and Champion, 2000) and a thermal anomaly that led to widespread melting of granitoid crust and the voluminous emplacement of

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monzogranite into the Yule and Carlindi Domes. At 2.85-2.83 Ga a post-tectonic suite of fractionated granitoids (Sweetapple and Collins, 2002) intruded four of the granitoid complexes. Globally, and in the EP, such granitoids are related to stabilisation of continental crust, and crustal concentration of the heat-producing elements U, Th, and K with partial remelting of earlier granitoids (Sweetapple and Collins, 2002). This contribution has reviewed the evidence that diapiric processes produced the Archaean dome-and-basin pattern of the EE Further examination of other Archaean domeand-basin terranes should determine the extent to which the EP diapiric model has global significance to the generation of Archaean continental crust.

2.7.

EARLY ARCHAEAN CRUSTAL COLLAPSE STRUCTURES AND SEDIMENTARY BASIN DYNAMICS

W. NIJMAN AND S.T. DE VRIES Introduction

Geodynamic models of Early Archaean crustal evolution are commonly based either on comparisons with Phanerozoic plate collisions, island arcs or metamorphic core complexes (e.g., De Wit, 1991, 1998; Kr6ner and Layer, 1992; Zegers, 1996; Blewett, 2002; Sugitani et al., 2002), or on solid-state diapirism, crustal delamination, or mantle-plume activity (e.g., Campbell and Griffiths, 1992; Choukroune et al., 1995; Collins et al., 1998; Hamilton, 1998; Zegers and Van Keken, 2001; see sections 2.5 and 2.6). Basin architecture, an important link between sedimentary and deformational records of crustal evolution, is largely unknown for much of the Early Archaean (cf. Eriksson et al., 2001b). Arkosic or quartz-arenitic depositional systems that record orogenic settings appear in greenstone belts only after 3.3 Ga (cf. Krapez, 1993; Heubeck and Lowe, 1994; Eriksson et al., 1997; Hofmann et al., 2001). Is this simply a consequence of their generally poor state of preservation, or was the architecture of the sedimentary basins formed on Early Archaean greenstone basement fundamentally different from that of younger basins? Unconformities provide an important means for dividing the Early Archaean structural history into sequences of events (Krapez, 1993; Buick et al., 1995c; Nijman et al., 1998b), but can also indicate the position of a sedimentary basin margin. Recognition of their structural control can help to distinguish between local and regional nonconformities. Knowledge of the depth of the Early Archaean basins is a necessary prerequisite in most geodynamic models. In the absence of palaeontological parameters, enough tools are left to distinguish fluvial from shelf environments, submarine from alluvial fans, or to assess water depth (Lanier and Lowe, 1982; Buick and Dunlop, 1990; Nijman et al., 1998a; Eriksson and Simpson, 2002). Misinterpretations, however, cannot be avoided, in particular with respect to low-energy shallow-water versus deep-water environments. There is a growing awareness that many early Archaean cherts were originally volcanogenic sediments (e.g., Sugitani, 1992; Lowe, 1999a). Carbonates do occur, but often as alteration mineralisation. The source of the silica and timing of the silicification The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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have been interpreted very differently. In this respect, hydrothermal circulation appears to have played a major role (e.g., Barley, 1984; Sugitani, 1992; De Ronde et al., 1994; Nijman et al., 1998a; Appel et al., 2001). Because hydrothermal processes are linked with the sedimentary environment (cf. Westall et al., 2001 c, for the biological implications), and with deeper seated processes in the crust, the controlling structural and geodynamic regime becomes an important factor. Sedimentology, therefore, may supply important tools for solving geodynamical problems of the Early Earth, particularly in combination with structural and geochemical analysis. Such an approach has resulted in an hypothesis concerning the generation of the first sedimentary basins as major, possibly ring-shaped, crustal collapse structures, like the coronae on Venus. This is the topic of this contribution.

Early Archaean Extensional Growth-Fault Structures The Coppin Gap greenstone belt of the East Pilbara Nijman et al. (1998b) described the interrelationship between sedimentation and deformation in the Coppin Gap greenstone belt (CGB), a synclinally folded greenstone belt in the eastern part of the Pilbara Craton (see Fig. 2.7-1a). Figure 2.7-2 illustrates the main stratigraphic and structural features of the belt. These are: (1) the belt is composed of stacked shear zone-bounded slices of Early Archaean ultramafic, mafic, intermediate, and felsic volcanic and intrusive rocks, with minor BIF and other chert sediments; (2) U-Pb zircon dating indicates that the Archaean chronostratigraphic succession is not fundamentally disturbed (for details, see Fig. 2.7-2); and (3) extensional normal faults (D1) with west block-down thickness differentials occur at the scale of the entire greenstone belt ([ 1] in Fig. 2.7-2), and, at a smaller scale, as arrays within composing stratigraphic units ([2] in Fig. 2.7-2); they merge downwards and westwards in east-over-west extensional detachments, some with subsequent inversion into west-over-east thrusts (D2). Several aspects of this greenstone belt architecture require special attention with respect to the topic of this section: (1) the low-angle detachment surfaces; (2) the growth-fault arrays: examples of interference between sedimentary rock and structure; (3) the persistence of extensional growth faults during the growth of the Early Archaean stratigraphic column; and (4) the inversion of extension tectonics into thrust tectonics at 3.3 Ga. Low-angle detachment surfaces In the central CGB, 3471 + 5 to 3459 + 2 Ma intermediate to felsic volcanic rocks of the Duffer Formation (Fig. 2.7-2) of the Warrawoona Group reach a maximum thickness of over 5000 m (cf. DiMarco and Lowe, 1989). They are underlain by a detachment surface along which several tens of metres thickness of BIF have been deformed into buckled and recumbent folds in association with listric slide surfaces in a pattern characteristic of eastover-west sliding ([3] in Fig. 2.7-2). Some of the folds envelop 100 m-diameter blocks of felsic volcanic porphyry, detached from the lower part of the Duffer Formation. This megabreccia is interpreted in the context of gravitational collapse of the entire volcanic unit along its base as a detachment surface.

2.7. Early A rchaean Crustal Collapse Structures

Fig. 2.7-1. Reference maps for (a) the East Pilbara craton and (b) the Kaapvaal craton.

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Fig. 2.7-2. Longitudinal cross-section of the Coppin Gap belt (CGB) along the fully drawn part of the traverse in Fig. 2.7-1, through the southern limb of the CGB syncline (for detailed maps see Nijman et al., 1998b; Williams, 1999; Van Kranendonk et al., 2001). The section is restored for post-D2 deformation. BIF and clastic sedimentary rocks, with minor mafic and some felsic volcanic rocks, belonging to the 3.25 Ga Gorge Creek Group unconformably overlie the 3.45 Ga, predominantly volcanic, Warrawoona Group, composed of Salgash Subgroup, Duffer Formation, and Talga Subgroup. The 3.314 Ga Coppin Gap Suite of the Mount Edgar Batholith (Williams and Collins, 1990) intrudes the Warrawoona Group and its shear zones. Structurally, the CGB shows stacking of units along shear zones of east-over-west extensional detachment (D 1), some of them with a subsequent inversion into west-over-east thrusting (D2). [1 ]: extensional normal faults, at the scale of the entire greenstone belt; [2]: idem, as arrays within composing stratigraphical units; [3]: detachment surface at the base of the Duffer Volcanics with megabreccia; [4]: gravitational folds related to D 1 detachment.

Several such low-angle detachments can be distinguished in the CGB, where they tend to descend westwards in the stratigraphic succession. It is here suggested that the recumbent northwest-facing folds described by Collins (1989) in the McPhee Reward section of the adjacent Marble Bar greenstone belt do not represent collapse structures along the west flank of the doming Mount Edgar Granitoid Complex (Collins et al., 1998), but gravitational collapse folds along one of these detachments ([4] in Fig. 2.7-2). The latter appear to be common features in the early history of the architecture of the greenstone belts. They also underlie arrays of normal growth faults, in both the Pilbara and Kaapvaal cratons.

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Growth fault arrays: examples of interference between sedimentary rock and structure Three Early Archaean extensional fault arrays of similar size are compared in Figure 2.7-3: the c. 3.45 Ga South Kittys Gap volcano-sedimentary complex (SKGVSC) of the Panorama Formation of the Salgash Subgroup in the CGB (Fig. 2.7-3a), the c. 3.49 Ga North Pole volcano-sedimentary complex (NP-VSC: Dresser Formation of Van Kranendonk, 2000) of the Talga-Talga Subgroup in the North Pole Dome (Fig. 2.7-3b), both in the Pilbara craton (Fig. 2.7-1a); and the c. 3.44 Ga Buck Ridge volcano-sedimentary complex (BR-VSC: Lowe and Byerly (1999a) use the name Buck "Reef" instead of Buck "Ridge" on grounds of original use by Hall, in 1918, but we prefer the customised use of Buck Ridge) of the Onverwacht Group of the Barberton greenstone belt of the Kaapvaal craton (Fig. 2.7-3c; cf. Fig. 2.7-1 b). North Pole (NP-) Chert and Buck Ridge (BR-) Chert refer to major sedimentary chert units within the respective complexes. The SKG-VSC (Fig. 2.7-3a) and BR-VSC (Fig. 2.7-3c) are both composed of felsic volcanic rocks (rhyolites and dacites), overlain by sedimentary chert and chertified, commonly volcaniclastic sedimentary rocks. The contacts are conformable and, in places, interdigitating. Both volcano-sedimentary assemblages are sandwiched between pillow basalts. In the BR-VSC the basaltic substratum alternates with felsic porphyries and has been intruded by small trondjhemite bodies. The NP-VSC (Fig. 2.7-3b) is devoid of felsic volcanic components and is characterised by an alternation of chert and pillowed basalt. In the three examples, the depositional depth of the cherts is near base level. The approximate depth of the NP-Chert basin floor facies is estimated at about 50-70 m (Nijman et al., 1998a), whereas tidal structures are indicative of an intertidal environment for much of the remainder of the cherty sedimentary rocks. In the BR-Chert, stalactite cements probably even indicate a vadose environment (i.e., conditions of shallow emergence). In the SKGVSC, near-zero depth is also indicated by the channel-and-flat geometry and heterolithic character of the chert and by the immediately overlying massive basalt with well developed flow-foot breccia (Fumes and Sturt, 1976; Cas and Wright, 1987). Sedimentologically, the sequences of felsic volcanic and chert units, between the underlying and the overlying pillow basalts, are regressive to transgressive. Evidence for the growth-fault character of the normal fault arrays (cf. Nijman, 1999) is provided by: (1) Stratigraphic thickness differentials that diminish to zero towards the youngest bed involved. The latter evidently buries the structure. The faults affect all available sedimentary and volcanic rock types, including bedding-parallel porphyry sills. Maximum thickness occurs in the hanging wall, and minimum in the footwall, where hiatuses and local angular nonconformities were also formed (e.g., SKG-VSC, Fig. 2.7-3a; BR-VSC, Fig. 2.7-3c). (2) The distribution of sedimentary facies, which has been controlled by the faults. In the NP-chert, stacked channel-fills cluster in the hanging walls of the fault blocks (Nijman et al., 1998a), whereas in the BR-VSC coarse clastic deposits are concentrated along the downthrown sides of the normal faults. Moreover, sedimentary facies in the BR-VSC gradually change from proximal conglomerate and sandstone along the eastern faults to distal even-laminated dense cherts and Fe-oxide-rich shale in the west. (3) Downwards merging of the normal faults into a bedding-parallel, complex shear zone with sense of shear synthetic with respect to the faults. (4) Fault arrays up

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Fig. 2.7-3. Growth-fault arrays, the relationship between chert sedimentary rocks and tensile growth faulting in Early Archaean greenstone belts. (a) South Kittys Gap felsic volcano-sedimentary complex, Coppin Gap belt, Pilbara (note that because of the subvertical dip of the strata, the map reads as a cross-section; data from Nijman, de Vries, Vos, Louzada; and Williams, field work from 1999). The normal-faulted SKG-VSC is sandwiched between foliated and highly chemically altered zones of shear deformation: the Bamboo Creek Shear Zone truncating the top of the unit, and a complex basal detachment zone into which the listric normal faults merge at about 1 km below the top of the unit. Both shear zones show evidence for inversion from westwards extensional detachment to eastwards thrusting. Above the basal detachment zone, a megabreccia of rotated blocks of felsic volcanic in a matrix of deformed basalt indicates sliding of the SKG-VSC under relatively low-overburden pressure, i.e. under uppermost crustal conditions. The flow-banded quartz-feldspar-porphyritic rhyolites and agglomerates, capped by chert and shale, form wedges that thicken towards the hangingwalls of the normal faults. Some of the fault blocks develop a roll-over anticline. Progressive compressional deformation of previously-formed normal faults in the toe of the array and the occurrence of undeformed normal faults at the head indicate a back-stepping of the normal faults during sedimentation. Densely spaced fractures, often chert-filled, emphasise the extensional character of the deformation. Metres-wide black chert veins arise from within the top tens of metres of the felsic volcanic rock unit and terminate in the capping sedimentary chert.

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Fig. 2.7-3 (continued). (b) North Pole volcano-sedimentary complex, Pilbara (after Nijman et al., 1998a). A cross-section along the NP-VSC, restored for post-Dl tilt and radial faulting due to the formation of the North Pole Dome, shows an array of synsedimentary, listric, tensional growth faults with thickness differential in the infilling sequence of cherts (I-V) and interlayered pillow basalts. Chert units II and IV are used as horizontal reference for the restoration. Swarms of synsedimentary hydrothermal chert-barite veins fan upwards from silica-poor centres that project at or near the growth-fault planes in the underlying basalt at 1500 m below reference. Other chert-vein generations are strike-parallel, or at right angles to bedding without obvious fanning. The faults merge from a basal shear zone with intense hydrothermal chemical alteration of the sheared basalt. The chert-barite unit I formed in a shallow tidal basin (cf. Dunlop et al., 1978; Buick and Dunlop, 1990), in which the distribution of facies, architectural elements, and barite mineralisation is controlled by the faults (Nijman et al., 1998a). to 20 km wide and which have a detachment depth of c. 1-2 km and a fault spacing of c. 1-4 km. Rollover anticlines have also been recognised. Tilt of up to 60 ~ accommodates the thickness differences (BR-VSC, Fig. 2.7-3c). Extension at the head is often accompanied by synthetic, to a lesser extent antithetic, compressional deformation at the toe of the structures (Figs. 2.7-3a, c), a characteristic feature of surficial gravitational collapse deformation (Mandl and Crans, 1981). It is commonly accompanied by chaotisation and megabrecciation of the stratigraphic sequence (BR-VSC: head and toe regions, Fig. 2.7-3c; basal detachment of SKG-VSC, Fig. 2.7-3a). (5) Complex black chert, felsic porphyry and basaltic vein systems radiate from the fault system towards the capping chert (very prominently in NP-VSC, Fig. 2.7-3b). The veins interfere with the sedimentary and early diagenetic structures. There is ample evidence of synsedimentary hydrothermal mineralisation along the faults and around the fault terminations (barite in the NP-Chert, Fig. 2.7-3b). The primary character of the barite was recently confirmed by Runnegar (2001, cf. Lambert et al., 1978; nahcolite; Lowe and Fisher Worrell, 1999, in the BR-Chert, Fig. 2.7-3c).

Persistence of Early Archaean extensional growth faults Extensional structures without conclusive evidence of preceding compression appear to dominate the structural scene during the Early Archaean. In the Pilbara, normal fault ar-

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Fig. 2.7-3 (continued). (c) Buck Ridge volcano-sedimentary complex, Barberton Greenstone Belt, South Africa. Because of the subvertical dip of the strata, the map reads as a cross-section. Based on field data of De Vries, De Wit, King, Houtzager, Louzada, and Nijman; and on De Wit (1983); Dann (2000); all intermediate, mafic and ultramafic intrusions are omitted. The BR-VSC is composed of a chert-capped alternation of predominantly pillow basalts with felsic lava and porphyry, and minor komatiite flows and cherts along the base, intruded by felsic igneous rock: finely crystalline trondhjemite in the core, fsp-qtz porphyritic dacite along the periphery, with offshoots (dykes or pipes) into the overlying BR-Chert. The latter forms a regressive sedimentary sequence (cf. Fig. 2.7-6): from felsic lava and volcaniclastic sediments (unit 1) via conformable banded (heterolithic) sedimentary chert (unit 2) into Fe-oxide-rich sandstone and siltstone with breccia and iron pods (unit 3), overlain by transgressive silicified littoral sandstones (unit 4). To the west, late-stage basalt is intercalated between the lower units of the BR-Chert. The sedimentary facies in the BR-Chert changes from proximal polymict conglomerate and sandstone in the east to distal even-laminated dense cherts and Fe-oxide-rich shale in the west. Geometrically, the BR-VSC is subdivided by 3 to 4 km-spaced listric normal faults (evidence from fault traces, strike discontinuities, and the asymmetrical compartmentalisation of the felsic intrusions). Maximum thickness in the hangingwalls, and minimum thickness with local internal nonconformities in the footwalls result in up to 60 ~ tilt against the controlling faults. Some develop rollover anticlines. At the western toe region of the fault array, several chert layers are--gravitationally?--detached from the normal-faulted base and encased in a matrix of Fe-rich shale. Complex vein systems (black chert, felsic porphyry, basalt) radiate from the fault system towards the capping chert. There is ample evidence of associated hydrothermal mineralisation and alteration (De Vries, in prep.). Two kilometres below the top of the BR-VSC the faults merge as major listric features into a bedding-parallel, complex shear zone (Geluk Fault of Lowe et al., 1985, 1999). It shows east-over-west sinistral shear asymmetry of compressional folds and faults, i.e. synthetic with respect to the normal faults, and separates the deformed BR-VSC from the underlying, regular, mafic rock succession (Dann, 2000). At the rear of the normal-fault array, the stratigraphic succession becomes chaotic, the normal faults are folded, and one of the rollover anticlines is steeply overturned. This is explained by underthrusting (better "undersliding") of normal fault wedges during progressive gravitational collapse along the rear of the fault array (cf., text section on the influence of TTG emplacement on basin development). All features in this example are considered highly diagnostic for the syndepositional character of the normal faults (De Vries, in prep.; Nijman, 1999). Viljoen and Viljoen (1969c) first recognised the normal fault style of the complex, and Lowe and Fisher Worrell (1999), mention minor growth faults within the BR-Chert.

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rays have been found dispersed over the entire craton and at different stratigraphic levels (see also Zegers et al., 1996; Oliver and Cawood, 2001; Kloppenburg et al., 2001; Van Kranendonk et al., 2001). In the CGB, normal fault arrays are vertically stacked, and are incorporated in and affected by similar fault systems of a higher order of magnitude (cf. Gibbs, 1987). Since the tensional deformation is synsedimentary, the growth-fault patterns have apparently migrated upwards with the deposition of the rock succession. Along the cross-section exposed in the CGB, all extensional structures face due west independent of their size and age. In adjacent belts the orientation differs considerably. Given the scale of the extensional structures, the en echelon faults and the upright isoclinal and rootless folds in the 3.51 Ga Coonterunah succession below a basal unconformity of Strelley Pool Chert (Buick et al., 1995c; Green et al., 2000; Van Kranendonk et al., 2001) might also be explained as the result of large scale extensional gravitational gliding and block rotation. Both in the East Pilbara and Barberton areas, compressional tectonics follow the earlier extensional regime at c. 3.3 Ga (Zegers et al., 1998). In the SKG-VSC (Fig. 2.7-3a), for instance, inversion is recorded by D2 refolding of S l--cleavage in the bounding shear zones with development of an $2 axial-planar crenulation cleavage. The phase of compressional tectonics is well dated as it is intruded by 3.314 Ga granitoids in the CGB (Williams and Collins, 1990). Compression interfered with the clastic sedimentation of the Gorge Creek Group in the Pilbara craton (Nijman et al., 1998b) and in the Barberton greenstone belt, with the Moodies Group (Lamb, 1986). The Isua greenstone belt of West Greenland New data have recently become available on the other early Archaean type area, the > 3.7 Ga Isua greenstone belt of West Greenland (Rosing et al., 1996; Appel et al., 2001; Fedo et al., 2001; Myers, 200 l b; section 2.2). Now that the deformation, metamorphism, and metasomatism at Isua have been investigated in more detail, the most common supracrustal rock association is of pillow basalts with chert and BIF, with the contribution of felsic volcanic rock less than was previously thought (see section 2.2). TTG sheets have been recognised to have intruded these early rock sequences (Nutman, 1997), as they do in the Barberton greenstone belt. Conglomerate and sandstone are rare and derived from local source areas. Carbonates, previously considered as primary sediments, are instead now seen as the product of hydrothermal or metasomatic replacement mineralisation. Bathymetric estimates range from an undefined oceanic depth below wave base to emergence (Cas et al., 2001; Fedo et al., 2001). In all these aspects, the supracrustal sequence of the Isua greenstone belt resembles its much better preserved counterparts in the Pilbara and Kaapvaal cratons. Using U-Pb zircon dating, Nutman et al. (2002) recognised at least three pre-3.5 Ga deformational events (c. 3660, 3640, and 3510 Ma) of ultramafic and mafic underplating of the oldest Archaean crust with or without concurrent extension. They rejected the model of intraoceanic convergent plate and collision tectonics proposed by Komiya et al. (1999). On the contrary, the extensional events seem to have preceded orogenic compression and

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concurrent clastic sedimentation, much like they do in the Pilbara and Kaapvaal cratons, albeit with a time difference of c. 200 My.

Crustal Collapse Structures as a Cause of Early Archaean Sedimentary Basin Formation The significance of growth faults in the Early Archaean The recognition of Early Archaean extensional growth-fault systems, like those mentioned above, is crucial for the understanding of the evolution of the earliest sedimentary basins. The observed arrays of listric growth faults may be expected to have been directly related to basin shape and development. Before any reconstruction of the sedimentary basin geometry, however, the extent of later tectonic deformation of the stratigraphic succession within the basins must be assessed. Many shear zones have been truncated by well-dated nonconformities and intrusive contacts, and are folded. An original position subparallel to bedding is often evident. As a consequence, tectonic slicing is considered to have influenced the stratigraphic thickness either in a negative (extensional) or a positive (compressional) sense. This is not only the case in the CGB, but also elsewhere in the Pilbara and the Kaapvaal cratons. Geochronologically, the general upwards-younging of the Early Archaean succession (Fig. 2.7-2) is used as an argument for stratigraphic continuity. However, because it is known from alpine orogens that superposition and shortening may involve successions formed in only a few million years, the density and precision of the available dates is still far from sufficient to exclude tectonic superposition. The style of D2 deformation has allowed transport of tectonic slices over long flats and short ramps (Fig. 2.7-2). This transport process may have involved thin-skin tectonics or gravitational sliding. Those previously formed D I extensional detachment zones properly oriented to accommodate subsequent thrusting, as in the CGB example, show evidence of inversion of motion; others oriented at right angles to the newly generated compression direction, like the NP-VSC, do not. Where inversion was involved, subsequent thrusting may have partly obscured or obliterated the effect of the previous extensional detachment. Therefore, the net tangential effect, shortening or lengthening, may vary according to the place of observation and the angle between the translation directions of the two deformations. Chronological continuity in the stratigraphic succession may therefore have been preserved in places. Sedimentary Basin Geometry Important parameters in defining the sedimentary basin geometry are the shape, determined by outline and depth, and the size. If the inference of extensional growth-faulting as the controlling factor for the Early Archaean volcano-sedimentary sequence is correct, the shape of the extensional basin will be reflected in the original facing (restored dip direction) of the normal faults. Between the Mulgandinnah and Lionel Lineaments (Fig. 2.7-1 ) of the eastern Pilbara, the facing of all Early Archaean extensional normal fault arrays and the vergence of associated compressional folds (Fig. 2.7-4), after correction for later folding

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Fig. 2.7-4. Map of the East Pilbara granite-greenstone terrane (cf. Fig. 2.7-1a) showing the restored fault facing and fold vergence of D I extensional structures. The pattern suggests the existence of at least one ring-shaped collapse structure of considerable size. (New data and partly modified data from Collins, 1989; Zegers, 1996; Nijman et al., 1998a, b; Kloppenburg et al., 2001; Van Kranendonk, 2001). The amount of D2 shortening is difficult to assess (see text) and is not implemented in the figure.

and doming (cf. Blewett, 2002), matches an arched pattern rather than a unidirectional one, as suggested by Zegers (1996) in relation to metamorphic core complex formation, or as expected along a plate boundary. This pattern is discordant to the actual form of the Archaean granitoid domes, which would not be the case in a model of long-lasting and persistent diapiric doming as the main controlling structural mechanism (Collins et al., 1998; see section 2.6). The inferred pre-D2 (pre-3.3 Ga compressional) configuration (Fig. 2.7-5) along the traverse in Figure 2.7-1 shows a partial cross-section of a basin, as yet without knowledge of the western closure. The inferred size is based on assumptions about the amount of eastwest shortening due to the D2-thrust phase (diameter of the basin in Fig. 2.7-5:160 km at about 10% shortening). The growth fault-controlled volcano-sedimentary facies changes from predominantly intermediate and felsic lava and agglomerate at the basin margin (Duf-

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Fig. 2.7-5. Tentative reconstruction of the facies distribution in the Early Archaean Warrawoona Group along the North Pole-Coppin Gap traverse through the East Pilbara (dashed + fully drawn line in Fig. 2.7-1; cf. partial representation of the traverse in Fig. 2.7-2). Marginal facies of intermediate and felsic volcanic rocks are concentrated in the normal-faulted CGB segment of the traverse, the central parts are dominated by basalt and chert as exposed in the North Pole dome. The direction of extension there is at right angles to the cross-section. The felsic volcanic rock (Panorama Formation: cf. Van Kranendonk et al., 2001), in the uppermost part of the North Pole dome succession, is here interpreted as belonging to a basin margin, opposite to that of the CGB, translated eastwards during D2 thrusting (Boulter et al., 1987; Nijman et al.,1998a; and new field data, 2002).

fer Formation volcanic rocks, Fig. 2.7-2; SKG-VSC, Fig. 2.7-3a) to predominantly banded chert and basalt in the centre (NP-VSC, see Fig. 2.7-3b; cf. a similar change from marginal facies to a more shaly basin centre facies in the BR-VSC, Fig. 2.7-3c). The approximate depth of the basin centre during a phase of active felsic volcanism and chert deposition is estimated at about 50 to 70 m on the basis of observations in the NP-Chert (Nijman et al., 1998a). The suggested basin shape and size, as well as the volcanic and sedimentary facies distribution, resemble those of unusually large caldera-like collapse structures (see also, section 4.6), with felsic and intermediary volcanic emanations concentrated along the faulted margin. Basin Dynamics Vertical crustal movement and sedimentation

The crustal collapse model proposed for these Early Archaean basins (Fig. 2.7-6) implies initial uplift from the general subaqueous submarine level of deposition of pillowed flood basalt to above base level with development of extensional faults and fractures. The amount of uplift may have been only in the order of some hundreds of metres. The widespread pillow basalts which encase the felsic volcanic-chert rock suites need not have been deposited at oceanic depth, whatever the definition of the latter may be for the Early Archaean (cf. de Wit, 1998). They lack all evidence of slope and base-of-slope submarine clastic systems, but easily intercalate with inferred shallow-water, tide-influenced cherty sedimentary rocks. A shelf depth is therefore assumed for these flood basalts.

2.7. Early Archaean Crustal Collapse Structures

151

Fig. 2.7-6. Summary scheme of relationships for the formation and infilling of an Early Archaean collapse basin, during a phase of felsic to intermediate volcanism in the overall mafic volcanic environment of subaqueous plateau basalts. Further explanation in text.

Uplift by crustal arching may have caused the local extension responsible for collapse along growth faults. The uplift and onset of collapse coincided with the appearance of TTG magmas and felsic porphyries at shallow crustal levels and with extrusion of dacites and rhyolites. Uplift also formed restricted source areas along the margins of the collapse basin to produce enough volcaniclastic sediments to compensate for the subsidence in the aftermath of volcanic activity. This phase is characterised by hydrothermal circulation in the fractured and faulted collapse structure, the formation of chert veins and hydrothermal vents, causing large-scale silica input in the sedimentation environment and concurrent mineralisation and chemical alterations. The balance of uplift, collapse, and sedimentation and lava input apparently stabilised the sediment-water interface of these basins at about

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sea level. If closed from the open sea, the collapse basin became a lake comparable to caldera lakes or was possibly subaerially exposed, as seems to be the case with the BRVSC; if connected to open sea, tidal influence became perceptible or even pronounced by resonance, as in the NP-VSC and SKG-VSC. The subsequent transgressive regime, recorded by littoral facies at the top of both NP-Chert (Nijman et al., 1998a) and BR-Chert (Unit 4: Fig. 2.7-3c), indicates a final prevalence of subsidence over sediment input, back towards the original plateau depth of subaqueous flood basalts. The combination of uplift, extension and gravitational collapse produced extensional and compressional gravitational structures at the same time. Evidence of early gravitational thrusting or sliding was previously discussed for greenstone belts in the Kaapvaal craton by de Wit (1982, 1991). Emplacement of TTG sheets at high crustal levels may have deformed the host rocks and thus influenced basin development. Stacking of sheets may create sufficient gradient for surficial collapse structures (Hutton, 1996). Lowe and Byerly (1999a) related deformation in the BR-VSC to the emplacement of a much larger dacite body than is actually observed and shown in Figure 2.7-3c. The trondjhemite bodies in the Barberton greenstone belt probably followed the traces of pre-existing or concurrently developing growth faults. Accommodation of felsic intrusions in space created within the extensional regime (Hutton, 1996) is therefore much more probable than lateral thrusting and localised uplift merely by the intrusive force itself.

Bimodal volcanism and deep crustal magmatic differentiation In these Early Archaean basins, chert and associated felsic volcanic rocks occur episodically within a bulk volume of basalt flows. From REE-analysis, a deep crustal eclogitic derivation was suggested for such rocks by Cullers et al. (1993; cf. Glikson, 1976). Vlaar et al. (1994) propose an Early Archaean crustal model of episodic sinking of cool slabs of hydrated basalt through a ductile, relatively thick (c. 45 km) basalt crust. Eclogite so formed tends to delaminate into the underlying upper mantle harzburgite. The buoyant rise of the latter causes remelting of the eclogite/harzburgite with production of both basalt and tonalite. This delamination model, but without the aspect of flake tectonics, has been elaborated by Zegers and Van Keken (2001; see also section 2.5). In particular, when resulting in shallow plume activity, this delamination process may be able to account for the giant caldera-like structures here proposed for the Early Archaean (cf. Nijman, 1999). Crustal collapse versus meteorite impact The extensional collapse structures bear a superficial resemblance to impact structures. Indirect evidence for meteorite impacts during the Early Archaean was presented by Lowe and Byerly (2002). These authors reported the occurrence of four spherule-bearing beds between 3470 and 3243 Ma in the Barberton greenstone belt. The spherule beds are interpreted to have formed by the impact of 20 to 50 km diameter meteorites (Byerly and Lowe, 1994). The older layers (the 3470 Ma layer 1 and 3260 Ma layer 2) consist of clastic sedimentary units with a 5-50% admixture of spherules of different composition. These

2.7. Early Archaean Crustal Collapse Structures

153

authors explain the high-energy character of these spherule-bearing beds in a generally low-energy environment as the result of tsunamis below wave base (see section 1.3). As to this sedimentological argument, the high-energy scour-and-fill and crossbed structures observed in BR-Chert unit 2, fit better with a model of volcanic extrusion and volcaniclastic sedimentation than one invoking impact-related tsunamis: the vertical sequence from felsic lava and coarse volcaniclastic sediments to low-energy sedimentary rocks up to emergence (cf. Fig. 2.7-6), records a well-organised system of decreasing volcanic and transport activity, rather than one of catastrophic impact. Fluctuating energy is also common in chert layers at deeper stratigraphic levels of the Onverwacht Group, which have well-preserved tidal channel-and-flat characteristics (cf. Lanier and Lowe, 1982, for the Middle Marker Chert). The long-lasting and persistent character of the growth-fault systems and the way they control the facies patterns is also inconsistent with models invoking meteorite impacts.

Planetary Analogues The range of island craters and submarine calderas found in the Taupo backarc zone of New Zealand (Stoffers et al., 1998) represents one of the few Phanerozoic and Recent geological settings that can be compared with these Archaean basins. Similar calderas also produce silica and barite (in white smokers, Ishibashi and Urabe, 1995) albeit at considerably greater oceanic depths than inferred for the Archaean basins (cf. Nijman et al., 1998a; Van Kranendonk et al., 2001). The Middle Cretaceous Minarets Caldera in California (Fiske and Tobisch, 1994) is also filled partly with banded chert deposits similar to those in the Archaean. In general, calderas are explosive volcanic structures related to felsic volcanism (see also, section 4.6). The persistent contribution of basalts in the bimodal volcanic fill of the Archaean collapse structures (e.g., the NP-VSC example) does not favour a caldera origin. There are no known ring-shaped collapse structures, unrelated to impact, comparable in size to these Archaean structures. The crust of Venus, however, shows a variety of extensional structures (Solomon et al., 1992; Phillips and Hansen, 1994) of which the coronae are comparable in both size and shape with those of the Archaean terrestrial basins (Nijman and De Vries, 2001). Solomon et al. (1992) describe coronae as volcano-tectonic structures unique to Venus, generally distinguished by a quasi-circular annulus of concentric extensional faults and fractures, 100-2600 km in diameter ( 100-1000 km according to Smrekar and Stofan, 1999). Volcanic flows from these fractures cover large portions of their interior and periphery (Fig. 2.7-7). The interior stands higher than the surrounding mesolands (the intermediate topographic level of Venus, perhaps comparable to base level on Earth). According to Smrekar and Stofan (1997, 1999) the coronae are related to plume-generated rises where partial remelting of delaminated lower lithosphere occurs above small-scale upwellings impinged asynchronously on the bottom of the lithosphere. This is considered an effective means for heat loss on Venus. Phillips and Hansen (1994) and Solomon et al. (1992) also mention an indirect generation by Rayleigh-Taylor instabilities in layers of melt above wider upwards mantle convections. The lithospheric thermal thickness is estimated to be 100 to 150 km,

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Chapter 2: Generation of Continental Crust

Fig. 2.7-7. Magellan radar image of Kuva Corona on Venus: diameter 230 km in bright colours: lava flows extruded from ring-faults (courtesy of NASA). and the thermal gradient to be Earth-like, though episodically (500 My) at a considerably higher general temperature. The correspondingly reduced crustal strength does not allow plate tectonics (Solomon et al., 1992). Venus' stress fields are therefore characterised by a more diffuse strain distribution than exists in the plate-motion-dominated Earth. To a lesser extent some Martian features also deserve notice, in particular the relationships between hydrothermal circulation, Fe-oxide sedimentation and extensional (gravitational) "chaos terranes" along the slopes of volcanic rises and volcanoes mentioned by Farmer (2000). Conclusion

The earliest (> 3.3 Ga) terrestrial sedimentary basins were characteristically filled with felsic and mafic volcanic products and cherty sedimentary rocks, and their development was controlled by normal listric growth faults arranged in non-linear patterns. These faults linked intermittently occurring shallow-level felsic intrusions via porphyry pipes, veins and hydrothermal circulation with the surficial sedimentary basin-fill of cherty sediments and concurrent mineralisation and alteration products. The extension tectonics did not represent a reaction to compression and crustal thickening. It also had no relationship with the present-day distribution of granitoids and greenstone belts. Crustal uplift, collapse and basin formation is best explained by crustal delamination and related plume activity. In the

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155

Pilbara and Kaapvaal cratons this phase of extensional crustal evolution ceased at about 3.3 Ga and was replaced by compression tectonics and associated clastic sedimentation more readily attributable to plate motion. A comparable transition may have taken place 200 My earlier in the Isua greenstone belt of West Greenland.

2.8.

CRUSTAL GROWTH RATES

N.T. ARNDT The ages of many important events in Earth history are known with remarkable precision: the planet formed by accretion 4.6--4.5 Ga ago, the core segregated around 4.56 Ga, and the early atmosphere formed about 20 My later (e.g., Canup and Righter, 2000) (section 1.2 provides a summary of the pre-3.8 Ga Earth). Heavy bombardment of meteorites, the end-game of accretion, ceased around 4.0 Ga, about the same time as the formation of the oldest known rocks. Life may have been present in the oldest preserved sediments (see, however, sections 2.2, 6.1, and 6.2), which were deposited over 3800 Ma ago (Schidlowski, 1988; Mojzsis et al., 1996). The first supercontinent assembled between 2.6 and 2.7 Ga (sections 3.2, 3.4, 3.6, and 5.3), major iron formations were deposited during the period 2.5-2.0 Ga (section 5.4), and the Palaeoproterozoic was punctuated by intense global ice ages (sections 5.6, 5.7, and 5.8). For the Phanerozoic, the three major tools of geochronology, palaeontology and palaeomagnetism, allow accurate dating of Wilson cycles, evolution of animals and plants, marine transgressions and regressions, glaciations, and major volcanic episodes. Oceanic crust forms and subducts continuously, and the oldest oceanic crust beneath present oceans is about 190 Ma. We do not know the age of formation of the Earth's continental crust. There are two schools of thought. Armstrong (1981, 1991) proposed that the continental crust grew rapidly in the Hadaean (see also, section 3.6) and had reached its present volume 4.0 Ga ago. Thereafter growth of crust was balanced by its destruction, mainly through subduction. Many other scientists, and most geochemists, reject this interpretation and believe instead that the continental crust started to grow at c. 3.9 Ga and has continued to grow progressively ever since (e.g., DePaolo, 1983; Jacobsen, 1988; Albar~de, 1998; Coltice et al., 2OOO). Arguments in support of the two theories are summarised in Table 2.8-1. Strong evidence for the Armstrong model comes from the discovery of zircons with ages up to 4.4 Ga (see section 1.2). These zircons form a small but significant proportion of the detrital minerals in a quartzite deposited c. 3.1 Ga ago. Their significance is two-fold: (1) these zircons resemble those in modern granites or felsic gneisses and probably came from such rocks; they provide evidence for the very early existence of granite, which is the essential constituent of continental crust; (2) the zircons have survived for c. 1300 My at the surface of the hot, tectonically unstable, meteorite-bombarded Early Earth and must have been protected by a stable, buoyant platform, one that resisted subduction back into the mantle. The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Table 2.8-1. Arguments relating to the two models of continental crustal growth Arguments for rapid early crustal growth (Armstrong's (1981) model)

Arguments for the continuous crustal growth model

4.4 Ga zircons from Mt Narryer, Australia: evidence of granitoids at the start of Earth history

Rarity of continental rocks older than 3.5 Ga; absence of rocks older than 4.0 Ga

4.0 Ga gneiss from Acasta, Canada; evidence of extant continental crust early in the Archaean

Continent growth continues to the present day in subduction zones

Isotopic depletion in the Archaean mantle: early extraction of an enriched reservoir

Lack of isotopic evidence for older crust in Archaean granitoids and sediments Isotopic data limiting the amount of continent recycled back into the mantle

This platform can only have been continental crust (see section 3.6 for a relevant model for early continental crust). Other arguments for the Armstrong crustal growth model include the discovery of the pre-4 Ga Acasta granitic gneiss (Bowring et al., 1989), the oldest known terrestrial rocks, and the pattern of early isotopic depletion of the Archaean mantle (DePaolo, 1980; Stein and Hofmann, 1994). This depletion records the extraction of enriched material, which, in the opinion of many authors, was continental crust. Advocates of the continuous growth model (e.g., Taylor and McLennan, 1985; O'Nions, 1992) were impressed by the relationship between the volume and age of continental crust: most material in the continents is younger than c. 2 Ga and very little is older than 3.5 Ga (Fig. 2.8-1). However, it is incorrect to argue that such distribution indicates necessarily that crust first started to form only 4000 Ma ago. Continental crust is continuously being destroyed, by recycling at subduction zones or by other processes. A small but significant proportion is periodically reincorporated into younger crust leaving no geochemical record of its existence. Programs like the Canadian Lithoprobe (Calvert and Ludden, 1999), which combines mapping, geophysics and geochronology, demonstrate that large volumes of old crust are hidden by surface veneers of younger crust--the proportion of pre-3.5 Ga crust may be higher than previously thought. Other arguments for continuous crustal growth are based on geochemical or isotopic tracers that provide upper bounds on the proportion of continental material that has returned to the mantle. Pb and Nd isotopic systems define limits to the amount of subducted continental sediments in the upper mantle (Kramers et al., 1998), but say less about material recycled from other parts of the continents. Ar isotopes provide further constraints (Coltice et al., 2000). Final resolution of the debate awaits better definition of the compositions of major reservoirs in the Earth (upper and lower continental crust, depleted mantle, lower mantle), the development of better geochemical tools, and, above all more comprehensive, multidisciplinary studies of all major regions of old continental crust. Even more interesting, perhaps, is the rate at which continental crust grew, particularly through the early part of the Earth's history. Compilations of ages in crustal rocks and of detrital zircons in sediments from major rivers produce spectra with several pronounced peaks separated by deep troughs (e.g., Moorbath and Taylor, 1981; Taylor and McLennan,

2.8. Crustal Growth Rates

157

20

JUVENILE CONTINENTAL CRUST A

c

15

IO" L_

I.l=

-o9 10 _e-

$ E

5

0.2

0.6

1.0

1.4

1.8

2.2

2.6

3.0

3.4

3.8

AGE (Ga) Fig. 2.8-1. Compilation of ages of crustal rocks indicating episodic crustal growth (from Condie, 1997).

1985; Condie, 1994a; Goldstein et al., 1997). A major peak at 2.7 Ga corresponds to a global surge of crustal growth (Fig. 2.8-1). This event is recorded on every continent, either by the formation of voluminous juvenile crust or by thermal overprinting of older crust. The other peaks are more regional: 2.5 Ga is prevalent in China and India; 2.1 Ga in West Africa and South America; and 1.8-1.9 Ga in North America and Australia (Goldstein et al., 1997). Before 2.7 Ga and after 1.8 Ga, the crustal growth pattern is more continuous. Between the peaks very little seems to have happened. Of the few zircons with betweenpeak ages, many correspond to anorogenic intrusions or granulite cooling ages. Convergent margin sequences seem to be rare or absent. If this pattern is confirmed it has major implications because it implies that between the peaks, there was little to no subduction (see Lindsay and Brasier, section 5.3, for similar periods of global plate tectonic stasis and activity). Each growth peak opened with massive eruption of basalt and komatiite, mainly in ocean basins but in part on flooded continental platforms (e.g., Arndt, 1999). Thirty million years after the eruption peak, the volcanic plateaus aggregated, together with oceanic arc sequences, to form the nucleii of continents. Voluminous granites then intruded the volcanic successions. From their thermal and geochemical signatures, many of the volcanic rocks appear to have formed by partial melting in mantle plumes (Abouchami et al., 1990; Arndt et al., 1997; sections 3.2 and 3.3). Each peak of crustal growth therefore started with a surge of plume activity and terminated with massive subduction. Then followed a long period of inactivity.

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158

What caused the surges of crustal growth? Perhaps the Precambrian was a transitional period between layered mantle convection of the Hadaean and present-day whole-mantle convection (Stein, 1994; see also, sections 3.2, 3.4, and 3.6). In a hotter Hadaean mantle, subducting slabs may not have been able to penetrate the transition zone and may have been restricted to the upper mantle (Condie, 1997a). Only in the late Archaean, around 2.7 Ga, was subducted oceanic lithosphere able to reach the lower mantle, in the form of massive avalanches that provoked return flow in the form of massive upwelling plumes (see also section 3.2). The plumes heated the upper mantle, accelerating the normal cycles of formation and subduction of oceanic crust. The combination of massive plume-related magmatism and accelerated oceanic crust formation set off the cycle of enhanced crustal growth. What caused the quiet inter-peak intervals? During partial melting, heat diffuses from residual unmelted solids towards the molten region, to provide the latent heat necessary for fusion. The extraction of the melt leaves a cooler-than-normal residue. Massive plume magmatism efficiently extracts heat and leaves behind a relatively cold residue. Following each peak of crustal growth, the Earth went through a longer period of sluggish activity, then both its interior and exterior were cooler than normal. During most of the Precambrian our planet may have oscillated from hot to cold, and only towards the end of the Proterozoic, as the Earth gradually cooled down, did our planet escape from the cycles of periodic crustal growth.

2.9.

COMMENTARY

D.R. NELSON AND W.U. MUELLER A persistent theme in the contributions of this chapter is the lack of consensus concerning the field interpretation of the preserved Archaean rock record. Field observations of the geological record provide the foundation for all interpretations of the processes responsible for the formation of the Earth's continental crust. Field-based investigation of the Isua greenstone belt by Myers (section 2.2) shows the critical importance of careful and detailed field observations for the correct interpretation of the Archaean geological record. This is particularly important in high-grade terranes, where intense deformation, the formation of gneissic compositional layering, and metasomatism, can generate geological features that may superficially resemble sedimentary layering. Similarly, careful field studies are also central to our understanding the geological evolution of low-grade Archaean terranes, such as the granite-greenstone terranes of the Pilbara and Kaapvaal cratons. This is demonstrated by the differing interpretations of the same field evidence that are central to the controversy concerning the mechanisms of emplacement of the ovoid granitic complex "domes" of early Archaean granite-greenstone terranes, discussed by Zegers (section 2.5) and Hickman and Van Kranendonk (section 2.6). In part, different interpretations arise because many geological features of Archaean continental crust--even the granitegreenstone terranes themselves--are unique to the Archaean era, and must have formed by processes for which there are no exact modern analogues. The Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

2.9. Commentary

159

Polat et al. (section 2.3) demonstrated that it is possible to use immobile-element geochemical attributes to draw comparisons between the igneous rocks of the Isua greenstone belt and those from better-understood, modern-day tectonic settings. These authors argued that, provided the geochemical characteristics of the Isua greenstones have the same geodynamic significance as their modern counterparts, the lsua greenstones probably originated in an intra-oceanic subduction zone setting. This implies that plate tectonic processes were operating as early as c. 3.8 Ga. Daigneault et al. (section 2.4) also argued that most volcano-sedimentary sequences of the Late Archaean Abitibi greenstone belt display the salient features of arc evolution, arc-arc collision and arc fragmentation, but noted the presence of plume-related volcanism during several phases of Abitibi subduction-related magmatism. Nijman and de Vries (section 2.7) presented new field evidence, indicating that the development of some of the earliest (> 3.3 Ga) and best-preserved terrestrial sedimentary basins was controlled by normal listric growth-fault systems arranged in non-linear patterns. These authors argued that development of these early sedimentary basins was unrelated to crustal thickening or the present-day distribution of granitoids and greenstone belts, but is best explained by crustal delamination and/or mantle plume activity. Emerging from the present lack of concensus about the processes by which Archaean continental crust was formed is the suggestion that subtle differences may have existed between the operation of those processes involved in the formation of Early Archaean granite-greenstone terranes, and of those responsible for formation of Late Archaean examples. Given the vast time span covered by the Archaean era, the existence of any such differences should not be surprising. The contributions within this chapter provide an overview of the nature of the evidence, and the diversity of views about the formation of the Earth's continental crust arising from that evidence.

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The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) 9 2004 Elsevier B.V. All rights reserved

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Chapter 3

TECTONISM AND MANTLE PLUMES T H R O U G H TIME

3.1.

INTRODUCTION

EG. ERIKSSON AND O. CATUNEANU The previous chapter has dealt with the generation of continental crust and crustal growth rates, and more particularly with Archaean greenstone belts and associated granite emplacement to form the granite-greenstone terrains characteristic of cratonic nuclei. The operation of plate tectonics and mantle plumes in the generation of cratons is implicit, at least in the view of many scientists (see section 3.6). This chapter examines the temporal distribution of mantle plumes and superplumes (cf. superplume events and large igneous provinces; sections 3.2 and 3.3), and the possibility of global magmatic events marking a transition between an Earth dominated by thermal/mantle processes and one where heat loss was predominantly through Phanerozoic-style plate tectonics (section 3.4). How far back the latter processes, as a crust-generating mechanism, can be justified is examined in section 3.6, where Trendall's (2002) "plughole" model of Hadaean geological evolution is discussed. The latter again provides for a possible transition from thermal to plate tectonic regimes. More generally speaking, an intimate association between "plume tectonics" and plate tectonics may be considered central to Precambrian geological evolution (e.g., Eriksson et al., 2001 a, b).

Basic Principles Acceptance of the plate tectonic paradigm as central to Archaean continental crustal growth (e.g., de Wit et al., 1992; Sleep, 1992; Krapez, 1993; Windley, 1995; Kusky and Vearncombe, 1997; Brandl and de Wit, 1997; de Wit, 1998; Mueller and Corcoran, 2001), when applied to well-founded evidence for Archaean heat flow of c. 2-3 times modern-Phanerozoic values, is responsible for a long-lived assumption of small, fast plates due to greatly enhanced mid-ocean ridge length (Hargraves, 1986). Similarly, structuralstratigraphic basin models based on the same paradigm are applied widely in interpretation of the cratonic geological record (e.g., Miall, 2000). Assuming Archaean plate tectonics to be viable, low angle subduction is inferred by many to have been common (section 3.5). Other workers argue that magmatic, plate-independent models may be more applicable to the Archaean (e.g., Campbell and Hill, 1988; R.I. Hill et al., 1992; Goodwin, 1996; Hamilton, 1998). Mints and Konilov (section 3.9) question the suture interpretation (i.e., Himalayan-style collisions) applied to many early Precambrian granulite-gneiss orogenic

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Chapter 3: Tectonism and Mantle Plumes Through Time

belts, and argue for a within-plate and plume-related origin. Almost all researchers agree, however, that by c. 2 Ga, plate tectonism was firmly established on Earth. Rates of pre-2 Ga plate movement inferred for several examples (c. 2.6-2.1 Ga Transvaal basin, Kaapvaal craton; c. 2.4-2.1 Ga attenuation of the postulated Neoarchaean Kenorland supercontinent; c. 3.1-2.8 Ga Witwatersrand basin, Kaapvaal) suggest slower rather than enhanced rates (Aspler and Chiarenzelli, 1998; Catuneanu, 2001; Eriksson et al., 2001c). More logically, the interaction of "plumes and plates" led to variable rather than universally higher or lower rates of plate migration during the later Archaean and earlier Proterozoic, prior to c. 2 Ga (Eriksson et al., 2001 a) (section 3.6). As an illustration of these ideas, large mantle plumes impinging on cratonic plates would logically have stopped subduction of cold oceanic crust at their margins, while simultaneously shifting island arc systems further offshore; there, arc imbrication and the formation of large obductive arc complexes (cf. the "intra-oceanic" tectonic model of de Wit, e.g., 1998) may have resulted from continued mid-ocean ridge-push. Plume abatement would then have been followed by enhanced rates of arc complex accretion onto the continents, raising crustal growth rates (section 2.8). Of course, during a global superplume event, many more plume heads would have hit beneath ocean crust, and most likely would have raised rates of growth of juvenile (oceanic) crust at constructive ocean plate margins. The data base for determining plate movement rates is much better constrained for the Neoproterozoic. Strong evidence for anomalously rapid plate motion then is related to deep-seated mantle plumes and concomitant thermal instability beneath the lithospheric plates (Meert and Tamrat, section 3.11). Evidence in favour of a major global superplume event (Condie, section 3.2) or large igneous province (LIP) cluster (Ernst, section 3.3) at c. 2.7 Ga is complemented by the ideas of Nelson (section 3.4), that catastrophic mantle overturn events encompassing wholemantle convection rather than a layered mantle circulation system, may reflect a transition from an earlier thermally dominated Earth to one characterised by plate tectonics. Trendall (2002) provides an essentially thermally-driven Hadaean model for formation of the earliest cratonic nuclei, the growth of cratonic keels or roots, as well as the evolution of greenstone belt successions through predominant volcanism, subordinate largely chemical sedimentation, and the operation of both extensional and compressional tectonic regimes (section 3.6). An intimate relationship is inferred between mantle superplume events and the supercontinent cycle (section 3.2). Archaean greenstone belts can be considered, to a degree at least, as "LIPs" (Eriksson et al., 2001 b). Many greenstone belts (see also detailed discussions of greenstone evolution in chapters 2 and 4) probably formed, also, as plumegenerated oceanic plateaus (e.g., Abbott, 1996; Polat et al., 1998; Puchtel et al., 1998a), later accreted tectonically onto continental nuclei. By the Mesoproterozoic, when operation of Phanerozoic-style plate tectonics is accepted almost universally, a strong inter-relationship between such processes and those of mantle plumes remains pertinent, being well illustrated by the Midcontinent Rift System, U.S.A. (Ojakangas et al., 2001c). For the Mesoproterozoic supercontinent, Rodinia and Palaeopangea configurations apart, multiple plate tectonic genetic events can be determined in relatively great detail; however, the influence of plumes is also pertinent

3.2. P r e c a m b r i a n S u p e r p l u m e E v e n t s

163

(Frimmel, section 3.10). The earliest Himalayan-style orogenic belts, the approximately coeval, c. 2.7 Ga Limpopo and Hoggar belts, strongly support plate tectonism in the Neoarchaean (section 3.8). Although considered by some as controversial, identified ophiolite complexes cluster at times in the Archaean and Palaeoproterozoic when supercontinental assembly is inferred (Moores, 2002). A causative link with the supercontinent cycle and assembly above geoidal lows (cf. mantle downwarps, Condie, section 3.2) is a logical interpretation (Chiarenzelli and Moores, section 3.7). Based on the papers in this chapter, the importance of the interaction of mantle plumes and plate tectonics as first-order controls on crustal evolution during the Precambrian is emphasised strongly. It is conceivable that early, Hadaean proto-cratons evolved from fully thermal processes within a whole mantle convection scenario (Trendall, 2002) to an Earth, where heat loss was largely achieved through mid-ocean ridges. This change was likely related to catastrophic mantle overturn events and/or a major global superplume event, both possibly at c. 2.7 Ga, and the onset of recognisable plate tectonics followed. It is notable that evidence for an early supercontinent ("Kenorland"), for the first Himalayanstyle plate collisions and for significant chemical changes in the ocean atmosphere system (chapter 5) occur at about this same Neoarchaean time period.

3.2.

PRECAMBRIAN SUPERPLUME EVENTS

K.C. CONDIE Introduction

Because the term "superplume" has been used in different ways in the literature, it is necessary to constrain the term as it will be used in this section (see also section 3.3). A superplume is a large mantle plume with a well-defined head and tail, presumably coming from the D" layer above the core-mantle interface (Condie, 2001a). At the base of the lithosphere, the plume head flattens, spreading to a diameter of 1500 to 3000 km. Single superplumes typically give rise to large erupted volumes of mafic magma (> 0.5 • 106 km 3) in periods of time < 3 My. In contrast, mantle upwellings are large volumes of mantle that move upwards as part of the return flow of mantle convection ( ~> 10,000 km across). These upwellings are not plumes in that they do not rise from thermal boundary layers as distinct blobs that divide into head and tail components. Condie (1998) and Isley and Abbott (1999) have presented arguments that superplume events have been important throughout Earth history. A superplume event is a short-lived mantle event (~< 100 My) during which many superplumes and smaller plumes rise through the mantle and bombard the base of the lithosphere (Condie, 2001a). During a superplume event, plume activity may be concentrated in one or more mantle upwellings, as during the Mid-Cretaceous superplume event some 100 Ma ago, when activity was focused mainly in the Pacific mantle upwelling (Larson, 1991 a). The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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The strongest evidence for superplume events in the geologic past is the episodic age distribution of plume-related igneous rocks (see also section 3.3), such as komatiites, picrites, flood basalts, and giant dyke swarms (Isley and Abbott, 1999; Ernst and Buchan, 2002a). Also, the distribution of U-Pb zircon ages coupled with Nd isotopic data suggest two major peaks in juvenile continental crust production rate, one at 2.7 Ga and another at 1.9 Ga, both of which may be associated with superplume events (Condie, 1998, 2000) (Fig. 3.2-1).

The Cretaceous Superplume Event Larson (199 la, b) suggested that one or more superplumes beneath the Pacific basin could explain numerous Cretaceous geological features. The major evidence presented to support a Mid-Cretaceous superplume event is an enhanced rate in production of juvenile crust (oceanic arcs and oceanic plateaus), which peaks at 120-110 Ma. Many of the largest oceanic plateaus in the modern ocean basins were formed at 120-80 Ma (e.g., Ontong Java and Caribbean) (Kerr, 1998). Also, the Mid-Cretaceous pulse in production of juvenile crust correlates closely with a superchron (long period of normal magnetic polarity), suggesting that the heat source for the crustal pulse is located near the core-mantle boundary, thus supporting the superplume idea.

3.2. Precambrian Superplume Events

165

Near-surface reservoirs on the Earth, and the carbon cycle are also sensitive to superplume events (Kerr, 1998; Condie et al., 2001 a). For instance, the palaeotemperature curve as deduced from oxygen isotopes shows a broad increase from 150 to 100 Ma, which appears to require excess CO2 in the atmosphere to produce global warming (Larson, 1991 b; Barron et al., 1995). Such an increase in atmospheric CO2 may have been caused by increased submarine volcanism related to superplumes in the Pacific basin. In addition, an approximately 125 m increase in eustatic sea level, reaching a maximum at about 100 Ma, also can be related to increased ocean-ridge activity, displacement of seawater by oceanic plateaus, and uplift of the oceanic lithosphere over superplumes (Larson, 1991a, b; Kerr, 1998). However, some of this rise in sea level was probably related to the continuing breakup of Pangea (Hardebeck and Anderson, 1996), and the effect of the superplume event is superimposed on the supercontinent breakup effect. Extensive deposition of black shale from 130 to 85 Ma also may reflect increased CO2 related to a Mid-Cretaceous superplume event (Jenkyns, 1980). Black shale deposition requires anoxia resulting from increased organic productivity and poor water circulation in basins on continental platforms, both of which can result from a superplume event: directly, by hydrothermal spring input of CH4 and CO2 into the oceans; and indirectly, by increasing sea level and the frequency of partially closed basins on continental shelves (Kerr, 1998; Condie et al., 2001 a).

Causes of Superplume Events Slab avalanches Models. Stein and Hofmann (1994) were among the first to advocate that episodic instability at the 660-km seismic discontinuity controls the episodic growth of continental crust. They suggested that convection patterns changed in the mantle from layered convection (the normal case), when the growth rates of continental crust were relatively low, to whole-mantle convection (see also sections 3.4 and 3.6) when the growth rates were high. Whole mantle convection occurs in short-lived episodes during which subducted slabs that have accumulated at the 660-km discontinuity catastrophically sink into the lower mantle, in a manner similar to that proposed by Tackley et al. (1997). Davies (1995a) proposed catastrophic global magmatic and tectonic events at 1 to 2 Gy spacings. The favoured models show layered convection, which becomes unstable and breaks down episodically to whole-mantle convection as in the Stein-Hofmann model. During the catastrophic mantle overturns, hot lower mantle material is transferred to the upper mantle and may be responsible for rapid episodic growth of juvenile crust. Peltier et al. (1997) extended thermal constraints to more thoroughly evaluate the catastrophic mantle models. They quantified the physical processes that control the Rayleigh number at the 660-km discontinuity, which in turn controls the frequency of slab avalanches through this discontinuity. They also suggest a correlation between avalanche events and the supercontinent cycle. Their results imply that slab avalanches occur at a spacing of 400-600 My, and that they are brought about by the growth of an instability in the thermal boundary layer at the 660-km discontinuity. During and after slab avalanches a

Chapter 3: Tectonism and Mantle Plumes Through Time

166

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large mantle downwelling is produced directly above the avalanches, and this downwelling attracts fragments of continental lithosphere, thus leading to the formation of a supercontinent. Condie (1998, 2000) also proposed a model to explain the episodic growth of juvenile crust based on superplume events in the mantle (Fig. 3.2-2). In this model, the supercontinent cycle operates independently of slab avalanches at the 660-km seismic discontinuity, except for the first supercontinent at 2.7 Ga, which may have formed in response to the first slab avalanche. The production rate of continental crust increases during slab avalanches in response to increased production rate of oceanic plateaus and of subduction-related crust. Crustal recycling rate may drop significantly below crustal production rate during slab avalanches due to the formation of supercontinents, which trap juvenile crust. The difference in timing between production of oceanic juvenile crust and continental juvenile crust may be as short as 20 My or as long as 100 My. What triggers a slab avalanche? Although seismic tomography clearly suggests that many descending slabs sink into the lower mantle today (Grand et al., 1997), this may not have been the case in the geologic past when the Earth was hotter. Numerical simulations by Yuen et al. (1993) show that at the higher temperatures that existed in the Archaean, the mantle would convect more chaotically. Their results show that during this time with a higher Raleigh number, which is also temperature dependent, phase transitions such as the perovskite transition, at 660-km, become stronger barriers and result in layered convection. Models of Christensen and Yuen (1985) and Zhao et al. (1992) yield similar results. This implies that during the Early Archaean, subducted slabs may have accumulated at the

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660-km boundary. It may have been in the Late Archaean, when the 660-km discontinuity became less "robust", that slabs catastrophically fell through to the lower mantle (Peltier et al., 1997; Condie, 1998). It is this catastrophic collapse that may have triggered the first superplume event in the D" layer above the core. What may have triggered later superplume events? Possibly the same process. As suggested by Stein and Hofmann (1994), the Earth may have reverted to layered convection after the Late Archaean event and slabs again collected at the 660-km discontinuity, which again failed some 800 My later to produce the 1.9-Ga event. Alternatively, the breakup of the Late Archaean supercontinent at 2.2-2.1 Ga may have triggered a slab avalanche resulting in the 1.9-Ga event.

Core rotational dynamics Another possible cause of a superplume event is resonance of tidal waves in the fluid outer core (Greff-Lefftz and Legros, 1999a). When the core rotational frequency and solar tidal waves are in resonance, frictional power may be converted into heat, destabilising the D" layer above the core, leading to the generation of many mantle plumes. Numerical models predict two major resonances in the past, one at about 3 Ga and another at about 1.8 Ga. These times correspond closely with the observed peaks in juvenile crust production at 2.7 and 1.9 Ga. During the core resonance periods, the temperature near the inner core boundary should increase, an effect that could stop inner core growth and produce a new momentum equilibrium for the geodynamo. This, in turn, could lead to a decrease in magnetic reversal frequency, thus accounting for the superchrons associated with Phanerozoic superplume events.

Superplume Events and Supercontinents If both supercontinents and superplume events exist, a perplexing question is that of how they are related in space and time. A related question is: can the supercontinent cycle operate independently of superplume events? The timing of various geologic events resulting from the supercontinent cycle and from superplume events is constrained chiefly by data from two sources: U-Pb zircon ages of juvenile continental crust; and results of computer simulations of mantle processes (Tackley et al., 1994; Condie, 1998, 2000). Results clearly suggest that superplume events occur near the beginning of supercontinent formation (Fig. 3.2-2). Regardless of the trigger, from the time a slab avalanche begins in the mantle to the time juvenile crust is produced is probably quite short, < 100 My. This is because slabs can sink to the bottom of the mantle in 100 My or less (Larson and Kincaid, 1996), and in a mantle in which viscosity increases with depth, mantle plumes can rise to the base of the lithosphere in a few million years (Larsen and Yuen, 1997). Peltier et al. (1997) suggested that the supercontinent cycle also results from slab avalanche events in the mantle. In their model, the avalanches produce mantle downwellings directly over the avalanches, which act as "catchment basins" for an aggregating supercontinent. However, if supercontinents accumulate over mantle downwellings and

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break up over mantle upwellings (Anderson, 1982; Gurnis, 1988; Lowman and Jarvis, 1996), both of which are a consequence of the supercontinent cycle, slab avalanches in the mantle may not be a necessary part of supercontinent formation. As previously mentioned, supercontinent breakup may actually trigger slab avalanching that leads to another superplume event. Numerical models clearly show that supercontinents affect the thermal state of the mantle, with the mantle beneath supercontinents becoming hotter than normal, expanding and producing mantle upwellings (Anderson, 1982; Gurnis, 1988). Actual breakup of supercontinents may be caused by increased mantle plume activity within these upwellings. Does this constitute a superplume event? Perhaps there are two types of superplume events. One type of event results from thermal blanketing caused by a supercontinent, and another type results from a slab avalanche in the mantle, perhaps triggered by the breakup of a supercontinent. Only the second type of superplume event would appear to increase significantly the production rate of juvenile continental crust.

The First Supercontinent One of the intriguing yet puzzling questions of any episodic model for production of continental crust is that of just how and why the first supercontinent formed. For a supercontinent to form requires a significant volume of continental crustal fragments that survive recycling into the mantle. Prior to the Late Archaean, the high mantle temperatures and inferred large mantle convection rates in response to large Rayleigh numbers probably resuited in rapid recycling of continental crust, presumably before continental pieces had time to collide to make a supercontinent (Bowring and Housh, 1995; Condie, 2002b). So what happened in the Late Archaean that led to formation of the first supercontinent? One possibility is that the first slab avalanche in the mantle at 2.7 Ga, which liberated mantle plumes from the D 'f layer, led to the production of large volumes of continental crust in a relatively short period of time (~< 100 My). However, if there were no earlier supercontinents to fragment, what triggered the slab avalanches? Although the triggering mechanism is unknown, possibly the total mass of slabs accumulated at the 660-km discontinuity reached a critical value and collapsed through the boundary. Mantle plumes resulting from a slab avalanche can produce juvenile crust in two ways: directly, by the production of oceanic plateaus, and indirectly by heating the upper mantle and increasing the production rate of ocean crust due to increased convection rates or/and increasing the total length of the ocean ridge system (Larson, 1991 a). The first supercontinent may have formed by collision of oceanic plateaus, surviving fragments of continental crust older than 2.7 Ga, and oceanic arc systems. Also contributing to growth of a Late Archaean supercontinent is the thick Archaean subcontinental mantle lithosphere, which is relatively buoyant (Griffin et al., 1998), thus resisting subduction during plate collisions. The relative abundance of Late Archaean greenstones with oceanic plateau geochemical affinities supports the idea that oceanic plateaus were a major contributor to a Late Archaean supercontinent (Condie, 1994b; Tomlinson and Condie, 2001).

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Supercontinents, Superplume Events and Near-Surface Reservoirs Because new ocean ridges form during supercontinent breakup, the supercontinent cycle may have important consequences for near-surface Earth systems (Worsley et al., 1986; Veevers, 1990). Supercontinent breakup creates new, narrow ocean basins having restricted circulation and hydrothermally active spreading centres. These features promote anoxia in the deep ocean as do new rift basins accompanying supercontinent breakup. The amount of shallow marine carbonate deposition critically depends on redox stratification of the oceans, as reducing environments are not conducive to carbonate precipitation. Should anoxic deep-ocean water invade continental shelves, it would facilitate organic carbon burial on the shelves, including the deposition of black shale and the accumulation of gas hydrates. The increase in length of the ocean-ridge network that accompanies supercontinent fragmentation promotes increased degassing of the mantle, and increasing atmospheric CO2 levels, and rising sea level should lead to warmer climates resulting in increased weathering rates (Berner and Berner, 1997). Increasing carbonate in the oceans together with a growing ocean-ridge system should also enhance rates of removal of seawater carbonate by deep-sea alteration. To the extent that these developments enhance the fraction of carbon buried as organic matter, they would also lead to an increase in the 613C of seawater because 12C is preferentially incorporated into organic carbon (Des Marais et al., 1992) (see also section 5.3). During a superplume event, ascending plumes warm the upper mantle and lithosphere, and thereby elevate the seafloor by thermal expansion and create oceanic plateaus by the eruption of large volumes of submarine basalt. The extensive volcanism associated with a superplume event should also pump significant amounts of methane into the oceanatmosphere system, where it is rapidly oxidised to CO2 (Caldeira and Rampino, 1991; Kerr, 1998; Condie et al., 2000). In addition, uplift of oceanic lithosphere and eruption of oceanic plateau basalts should release large amounts of methane from gas hydrates on the sea floor (Jahren, 2002). The increased greenhouse gases, in turn, warm the climate and enhance weathering rates. Oceanic plateaus can locally restrict ocean currents (Kerr, 1998), thus promoting local stratification of the marine water column leading to anoxia. Superplume events also should result in rising of sea level due to isostatic uplift and thermal erosion of the oceanic lithosphere above plume heads (Kerr, 1994; Eriksson, 1999). Also, oceanic plateaus contribute to a rise in sea level, and during superplume events when many oceanic plateaus form, the effect could be significant. In addition, a superplume event can account for several features of BIF (banded iron-formation) (section 5.4) deposition. First, the enhanced submarine volcanism and hydrothermal venting associated both with ocean-ridge and oceanic plateau volcanism may be the source of the iron and silica in BIE Furthermore, the elevated sea level caused by a superplume event provides extensive shallow marine basins along stable continental platforms necessary to preserve B IF against later subduction. Biological productivity during superplume events is enhanced by increased concentrations of CO2, increased nutrient fluxes from both hydrothermal activity and enhanced weathering, and elevated temperatures due to CO2-driven greenhouse warming. Carbonate

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precipitation is enhanced by increased chemical weathering and by marine transgressions. Increased hydrothermal activity on the sea floor should also increase the rate of deep-sea alteration, which in turn should increase the removal rate of carbonate from seawater.

Precambrian Superplume Events Komatiites, flood basalts, mafic dyke swarms, and layered mafic intrusions have been used as proxies for superplume events in the Precambrian (Isley and Abbott, 1999; Ernst and Buchan, 2002a; Abbott and Isley, 2002b) (section 3.3). A time series analysis of the data shows major superplume events at 2.75-2.70, 2.45, and 2.0-1.9 Ga and several minor or possible events between 2.5 and 1.75 Ga. The 2.75-2.70 and 2.0-1.9-Ga events correspond well with the peaks in juvenile crust formation at these times (Fig. 3.2-3). The 2.45-Ga peak corresponds to a juvenile crust formation event recorded in India and the North China craton, and an inferred superplume event at about 2.1-2.15 Ga correlates with the 2.15-Ga crustal formation event in the Guiana shield and in West Africa (Condie, 1992a, 1998). One or two peaks at 1.75-1.70 Ga correlate with widespread continental growth in Southern Laurentia and Southern Baltica at this time.

The 1.9-Ga event The widespread occurrence of giant mafic dyke swarms and flood basalts and a peak in abundance of juvenile continental crust suggest a major superplume event at 1.9 Ga (Condie et al., 2000; Ernst and Buchan, 2002a) (Fig. 3.2-3). In addition, widespread remnants of shallow marine sediments with depositional ages of 1.9-1.8 Ga suggest that sea level was relatively high at this time, a feature consistent with a superplume event. A peak in abundance of shallow marine sediments at 1.9 Ga suggests that a 1.9-Ga superplume event may have overpowered supercontinent formation at this time, resulting in a significant rise in eustatic sea level. This may reflect the relative timing of the two events: supercontinent formation with many craton and arc collisions at 1.85-1.70 Ga occurred on the tail end of the 1.9-Ga superplume event, and may have contributed to the lowering of sea level following the superplume event. Also supporting high sea level at about 1.9 Ga is the widespread occurrence of submarine flood basalts on continental platforms. Examples of such basalts are in the Ungava orogen in Quebec, the Birrimian in West Africa, and in the Baltic shield in Scandinavia (Arndt, 1999). Also consistent with a 1.9-Ga superplume event is a peak in black shale abundance and in the black shale to total shale ratio at this time (Condie et al., 2000, 200 lb) (Fig. 3.2-3).

Fig. 3.2-3. Time series distribution of mantle plume proxies (based on plume-related igneous rocks), banded iron-formation (BIF), shallow marine sediments, black shale/total shale ratio, and CIA for shales. After Isley and Abbott (1999) and Condie et al. (2000). CIA = [A1203/(A1203 +CaO + Na20 + K20) x 100] molecular ratio, with CaO representing the silicate fraction only.

3.2. Precambrian Superplume Events

Fig. 3.2-3.

171

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Chapter 3: Tectonism and Mantle Plumes Through 7~me

The Chemical Index of Alteration (CIA) (section 5.10) can be used to track the degree of chemical weathering in shale source areas, and thus, may be useful as a proxy for palaeoclimates (Condie et al., 2001b). Although CIA data from shales show scatter, due perhaps to later remobilisation of Ca, Na, or K, there is a major peak in CIA at about 1.9 Ga and another at 1.7 Ga. These peaks suggest that palaeoclimates were unusually warm at these times supporting increased input of greenhouse gases (CH4 and CO2) into the atmosphere, a feature expected during superplume events. The last major period of BIF deposition was at about 1.9 Ga when the large BIFs of the Labrador Trough in northern Quebec, the Animikie basin in Minnesota, and the Nabberu basin in Western Australia were deposited (Klein and Beukes, 1992) (section 5.4). As shown by Isley and Abbott (1999), this last peak in BIF deposition correlates well with a 1.9-Ga superplume event (Fig. 3.2-3). Stromatolites, layered structures thought to be deposited by microbial mat communities (section 6.5), are widespread in the Proterozoic with a prominent peak (or peaks) in distribution at about 1.9-1.8 Ga. Maxima at this time are found in the number of stromatolite occurrences, the diversity of stromatolites, and in the number of occurrences of microdigitate stromatolites (Grotzinger and Kasting, 1993; Hofmann, 1998). The peaks in abundance and diversity of stromatolites at about 1.9 Ga may reflect a combination of global warming, high sea level stands, and enhanced input of greenhouse gases into the sedimentary cycle, all of which may be related to a superplume event at 1.9 Ga. The 2.7-Ga event

In addition to widespread juvenile continental crust, a possible 2.7-Ga superplume event is recorded in the Kaapvaal craton in southern Africa by eruption of the Ventersdorp flood basalts (2.72 Ga). The fact that Ventersdorp lavas are almost entirely subaerial indicates that sea level did not rise on the Kaapvaal craton as it should during a superplume event. Supporting this interpretation is the deposition of rift-related, subaerial detrital sediments in the medial part of the Ventersdorp Supergroup and the absence or near absence of black shale and BIF (Eriksson et al., 2002b). The probability that the Kaapvaal craton rode high during a 2.7-Ga superplume event may be due to a direct hit of a superplume elevating the craton. An increase in sea level on the Kaapvaal craton following the superplume event, as reflected by widespread shallow marine sediments (lower Transvaal Supergroup, 2.6-2.4 Ga), may be due to gradual collapse of the large mantle plume head beneath the craton. Decreasing amounts of greenhouse gases pumped into the atmosphere during waning of the superplume event, negative feedback of continental weathering, and increasing albedo caused by a newly formed Late Archaean supercontinent(s) may have cooled worldwide climates and led to widespread glaciation at 2.4-2.2 Ga in Laurentia, Baltica and South Africa (Young, 199 la) (sections 5.6-5.8). There is also a good correlation between the alleged 2.7-Ga superplume event and peaks in the abundance of BIF (Fig. 3.2-3). As at 1.9 Ga, the number of occurrences of marine stromatolites shows a peak at 2.7 Ga (Hofmann, 1998), again perhaps recording enhanced input of C02 into seawater.

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Epilogue Whether or not superplume events have occurred in the geologic past remains debatable and speculative. However, the episodic distribution of continental crustal growth in Earth history would seem to necessitate some kind of episodic event in the mantle. What does the geologic record tell us about the possibility of superplume events in Earth history? With the accuracy of isotopic ages, sediment and fossil distributions in the stratigraphic record allow superplume events in the geologic past and strongly support major events at 2.7 and 1.9 Ga, as well as several minor events, the youngest of which is at 110 Ma. The effects on the atmosphere-ocean-biosphere system predicted to result from superplume events at 2.7, 1.9 and 0.11 Ga are impressive. As more precise ages become available, it should be possible to distinguish the effects of a superplume event lasting the order of 50 My from a supercontinent event that lasts for 100 My or more.

3.3.

LARGE IGNEOUS PROVINCE RECORD THROUGH TIME

R.E. ERNST, K.L. BUCHAN AND A. PROKOPH

Introduction Large igneous provinces (LIPs) constitute one of the most significant modes of magmatism throughout Earth history. These were originally defined as "massive crustal emplacements of predominantly mafic (Mg- and Fe-rich) extrusive and intrusive rock which originate via processes other than 'normal' seafloor spreading" (Coffin and Eldholm, 1994). However, the definition has evolved to focus on "transient" events that cover an area of > 100,000 km 2, are emplaced in a short time interval, and can be linked with the arrival of a mantle plume (section 3.2) originating in the deep mantle (Eldholm and Coffin, 2000; Coffin and Eldholm, 2001). Others (e.g., Condie, 2001a) (section 3.2) have used the term superplume for plumes rising from the deep mantle, and reserved plume for those originating at shallower levels in the mantle. Given the present preliminary level of understanding of both the size and depth-of-origin of plumes, and the size of their LIP products (e.g., Ernst and Buchan, 2001b, 2002b), we prefer to retain the original usage, applying the label plume to buoyant material rising through the mantle regardless of depth of origin (e.g., Campbell and Griffiths, 1992; Coffin and Eldholm, 1994). Current usage includes continental flood basalts, volcanic passive (continental) margins, oceanic flood basalts (oceanic plateaus and oceanic-basin flood basalts), but excludes submarine ridges and hotspot chains, which were part of the original definition. The surface exposure of LIPs varies with age. In the Cenozoic and Mesozoic, the LIP record consists mainly of continental and oceanic flood basalts. In contrast, in the Palaeozoic and Proterozoic, flood basalts are commonly deeply eroded thereby exposing their plumbing system of dykes, sills, and layered intrusions. The Precambrian l-arth: Temposand Events Edited by P.G. Eriksson, W. Alterrnann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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The extrapolation of the LIP record into the Archaean is more speculative. Erosional remnants of originally more extensive, relatively flat-lying, "normal" flood basalts include the Fortescue succession of the Pilbara craton in Australia and the Ventersdorp succession of the Kaapvaal craton in southern Africa (Eriksson et al., 2002b). However, greenstone belts (sections 2.4, 4.3 and 4.4) represent the predominant volcanic mode in the Archaean. These comprise variably deformed and metamorphosed sequences that are typically fragmented into fault-bounded packages. Even though short-duration events of significant (multi-kilometre) stratigraphic thickness have been recognised in many greenstone belts, correlation of these events across large areas has been difficult (see also sections 2.5-2.7). Nevertheless, two classes of non-arc-related greenstone belts, the so-called mafic plains and platform assemblages (Thurston and Chivers, 1990), may be analogous to modern LIPs (e.g., Condie, 2001a). Based on the presence of minor komatiites, these classes of greenstone belt likely have a plume origin. However, only in a few cases so far, notably in the Kam Group of the Slave Province, Canada and perhaps in the Bababudan Group of the Dharwar craton, India, and the Abitibi Belt (section 2.4) of the Superior Province, Canada is there sufficient data to correlate between greenstone belts and to demonstrate scales of basaltic magmatism similar to modern LIPs (Bleeker, 2003).

Large Igneous Province Record Ernst and Buchan (2001b) developed a data base of plume-head events, in part based on earlier compilations of young LIPs (Coffin and Eldholm, 1994, 2001), Archaean greenstone belts containing komatiites (Tomlinson and Condie, 2001), and units older than 1.5 Ga (Isley and Abbott, 1999). The data base of Ernst and Buchan (2001b) was updated in Ernst and Buchan (2002b) and Prokoph et al. (2003). Here we utilise this plume-head data base (after Prokoph et al., 2003) as a proxy for the LIP record (Figs. 3.3-1 and 3.3-2). This is an excellent approximation in the Phanerozoic and Proterozoic because in this period nearly all of the events are linked to a plume-head using LIP criteria (large size and short duration). Less certain are the Archaean events in our data base, for which the main criterion is the presence of komatiites. This is a plume criterion (e.g., Campbell and Griffiths, 1992) (see also section 3.2) but not necessarily a LIP criterion because the original, pre-deformation size of events is usually unknown. Events were divided into two classes, "A" and "B", on the basis of the likelihood of a link to a mantle plume-head, following the criteria of Ernst and Buchan (2002b). Thirty five events rated "A" are confidently associated with the arrival of a mantle plume-head, based on the following criteria: emplacement of a large amount of basaltic magma (areal coverage of volcanic and intrusive rocks ~> 100,000 km 2 in a short time of a few million years); the presence of a giant radiating mafic dyke swarm; or a link with a present-day hotspot. One hundred and thirty two events are rated "B", on the following criteria: "plume" geochemistry; the presence of high-Mg rocks (picrites and komatiites); event size and duration, > 100,000 km 2 within uncertain age range, or > 20,000 km 2 (or > 20,000 km 3 for layered intrusions) within a few million years; or the presence of giant dykes (> 300 km long). All the above criteria are also LIP criteria, except for the "link with a hotspot",

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175

"plume geochemistry" and the presence of "high-Mg rocks". The original compilation of Ernst and Buchan (2001 b, 2002b) contains an additional "C" class of events, the majority of which are rift-related. They require more work to assess their plume-head and LIP links and are not considered further here.

Interpretation of LIP Record The LIP record through time is fairly continuous (Fig. 3.3-2a, b). Only a few significant gaps are observed (e.g., at 350-500, 615-720, 2220-2400 and 3000-3300 Ma), but it is not clear if these are real or merely artifacts of an incomplete data base. The apparent greater plume frequency from 150 Ma to the Present is due to numerous oceanic LIPs, which augment the continental LIP record during this period (Ernst and Buchan, 2002b). When only the continental data are used for the period between 150 Ma and Present the plume rate matches the pre-150 Ma rate (Fig. 3.3-2a). However, in the older record the oceanic LIPs have mostly been lost to subduction, with the exception of thick oceanic plateaus and plume-related volcanics preserved above a buoyant arc (e.g., Cloos, 1993). A late Archaean increase in plume frequency from 2800 to 2700 Ma (see also sections 3.2 and 3.4) reflects an apparent increase in plume generation and alternatively may be explained by increased preservation of LIPs. The decrease in plume flux prior to 2800 Ma may be real, or could be an artifact of our incomplete understanding of this older fragmentary record. Estimates of average frequency based on the LIP record indicate about one continental LIP every 20 My since the end of the Archaean (Ernst and Buchan, 2002b). Analysis of the young record which still preserves the oceanic LIPs, suggests that the combined continental and oceanic flux exceeds one LIP every 10 My (Coffin and Eldholm, 2001).

LIP Clustering and "Superplume" Events Clusters of roughly coeval LIPs can be identified throughout the record (Figs. 3.3-1, 3.3-2c; Ernst and Buchan, 2002b). These may represent the ancient analogue of events such as the mid-Cretaceous Pacific "superplume" (section 3.2), a multiple plume event that includes the largest known LIP, Ontong Java, with a volume of 40 x 106 km 3 (Larson, 1991 a; Coffin and Eldholm, 1994). Some such clusters of plumes, sometimes termed superplume events (Condie, 200 l a) have been correlated with supercontinent breakup (e.g., Gondwana: Storey, 1995; and Rodinia: Li et al., 2003) (sections 3.2, 3.10 and 3.11). However, a proposed link between major plume clusters (" superplume events") and periods of enhanced juvenile crust production (Condie, 2001) (section 3.2) is more difficult to assess because it requires a burst in the production of oceanic LIPs, which are poorly preserved during ocean closure (Cloos, 1993).

Time Series Analysis Wavelet, spectral and cross-spectral analyses have been applied to the LIP record to assess cyclicity in the distribution of LIPs through the past 3500 My (Prokoph et al., 2003). In

Chapter 3: Tectonism and Mantle Plumes Through Time

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3.3. Large Igneous Province Record

177

Opposite: Fig. 3.3-1. P l u m e - h e a d e v e n t s t h r o u g h t i m e u s e d as a p r o x y for l a r g e i g n e o u s p r o v i n c e ( L I P ) r e c o r d . C o n t i n e n t a l a r e a s are c o d e d as f o l l o w s : N o r t h A m e r i c a i n c l u d i n g G r e e n l a n d ( N . A M . ) , E u r o p e ( E U R . ) , S o u t h A m e r i c a ( S . A M . ) , A s i a , i n c l u d i n g I n d i a a n d the S e y c h e l l e s ( A S I A ) , A f r i c a i n c l u d i n g the A r a b i a n P e n i n s u l a ( A F R . ) , A u s t r a l i a ( A U S T . ) , A n t a r c t i c a ( A N T . ) a n d o c e a n i c a r e a s ( O C N . ) . E v e n t l i s t i n g a n d f i g u r e is a f t e r P r o k o p h et al. ( 2 0 0 3 ) , i t s e l f m o d i f i e d f r o m E r n s t a n d B u c h a n ( 2 0 0 1 b , 2 0 0 2 b ) . L o n g - b a r e v e n t s are r a t e d "A", s h o r t - b a r e v e n t s are r a t e d " B " . R a t i n g s ("A" a n d " B " ) are f r o m E r n s t a n d B u c h a n (2001 b) a n d the c r i t e r i a are d i s c u s s e d in the text. S o l i d lines h a v e 2 o a g e u n c e r t a i n t i e s less t h a n 20 M y ( a n d u s u a l l y less t h a n 10 M y ) a n d t h o s e w i t h d a s h e d lines h a v e u n c e r tainties o f 2 0 - 5 0 My. In g e n e r a l , i n c r e a s i n g e v e n t n u m b e r c o r r e s p o n d s to i n c r e a s i n g age. E x c e p t i o n s i n c l u d e e v e n t 137 a n d 211 w h e r e a b e t t e r a g e e s t i m a t e ( E r n s t a n d B u c h a n 2 0 0 2 b ) w a s o b t a i n e d after E r n s t a n d B u c h a n ( 2 0 0 1 b ) w e n t to press. Plume head events (label, event name and rating). Numbered labels correspond to listing in Ernst and Buchan (2(X)lb). 1, Columbia River (A); 2, Afar (A); 5, NAVP (North Atlantic Volcanic Province) (A); 6, Deccan (A); 7, Maud Rise (B); 8, Sierra Leone Rise (B); 9, Carmacks (B); 1___00,Madagascar (A); 11, CCCIP (Caribbean-Colombian Cretaceous Igneous Province) (A); 12, Alpha Ridge-Queen Elizabeth Islands (A); 1_33,Wallaby Plateau (B); 1__44,Hess Rise (B); 1__55,Naturalistc Plateau (B); 1__6_6Hikurangi , Plateau (B); 18, Kerguclen-Rajmahal (A); 1___99Nauru , Basin (B); 2___00,Ontong Java (A); 2__!_1,Manihiki Plateau (A); 22, Pifi6n (B); 24, Paranfi-Etcndeka and Equatorial Circum-Atlantic (A) Itwo nearby coeval plume centres (Ernst and Buchan 2002b)]; 2__55,Gascoyne Margin (B); 2___66,Magellan Rise (B); 2"7, Shatsky Rise (B); 2___88,Sorachi (A); 29, Argo Basin Margin (B); 31, Karoo-Ferrar-Chon Aike (A) Imultiple plume centres (Ernst and Buchan 2002b)1; 32, CAMP (Central Atlantic Magmatic Province) (A); 3_33,Angayucham (B); 34, Wrangellia (A); 36, Siberian Traps (A); 37, Emeishan (A); 38, Cache Creek (I3); 3__99Himalaya Ncotethys (B); 41, Jutland (A): 44, Yakutsk (A); 45, East European Craton (A); 48, Crimson Creek (B); 4_9, Antrim lin Australia] (A); 50, Wichita Mountains (B); 52, Late Central-lapctus (B); 53, Middle Ccntral-lapetus (A); 5__44,Early Ccntral-lapctus (B); 55, Baltoscandian (B); 5.__6_6,Volyn (B); 5___88,Franklin (A); 6__!_1,Mundine Well (B); 66, WillouranGairdncr (A); 67, South China (B); 6___99,Bukoban (B); 7__22,Southern-PRT (B); 7__33,Arabian-Nubian Shield (B); 7_55,BlekingeDalarna (B); 8___55Laanila-Kautokcino , (B); 89, Southwestern USA diabase province (B); 9_9_0.Kewecnawan (A); 91, Umkondo (A); 9___22Camucuo , (B); 9__4_4,Abitibi Idike swarml (B); 9__55,Late Gardar (B); 9___88,Protogine Zone-2 (B); 10____33Boyagin , (B); 10___44,Sudbury [dike swarml (B); 10___55,Seal Lake-Mealy (B); 10____66Central , Scandinavian Doleritc Group (CSDG) (B); 107, Mackenzie (A); 108, Harp (B); 10____99Middle , Gardar (B); 116, Derim (B); 11____99,Bukoba-Kavumwe (B); 12(___)),Bangcmall-I (B); 12___!.1Axamo , (B); 122, Michacl-Shabogamo (B); 123, Trond-G6ta (B); 12___44,Moyic (B); 126, Kuonamka (B); 128, .~,land-,~,boland and V~innland (B); 13____44,HamE (B); 13"7, Uruguayan (B); 141, Taihang-Hengshan, (B); 14__22,Eastern Creek (B); 14.3, Hart (B); 144, Avanavero (Roraima) (B); 14___6_6,Sparrow (B); 147, East Kimberley (B); 154, Flin Flon belt (B); 15___66,SoutpansbcrgWatcrberg-Olifantshock (B); 16___0,Northern Baltica-2 (A); 162, Minto-Eskimo (B); 166, Kenncdy (B); 167, Lac de Gras (B); 169, part of Povungnituk Group (B); 17___Q0,Kangfimuit (B); 173, Bushvcld (A); 174, Fort Franccs (B); 17___88,Karelian (B); 179, Griffin (B); 18___00,Marathon--reversed magnetic l~)larity (B); 181, Marathon--normal magnetic polarity (B); 18____22Labrador , Coast (B); 18____44,Biscotasing (B); 186, Tulcmalu (B); 187, Birimian-Bandamian (B) (2 distinct events); 19___11,Ungava (A); 193, BN-I (B); 19___55,Koli (B); 196, Ongcluk-Hckpoort (B); 19"7, Malley (B); 20___22,Widgicmooltha (B); 20____33,Kaminak (B); 205, Woongana (B); 206, Matachewan (A); 207, Northern Baltica- 1 (A); 2 I___LRampur-Garhwal (B); 21___33,Great Dyke of Zimbabwe (B); 217, Maddina (B); 218, Upper Bulawayan (B); 21___99,Pnicl (B); 220, Prince Albcrt-Woodburn Lake (B) (in part dated at 2.73 Ga after Sandcman and Skulski, unpub, data in Skulski et al. 2002); 221, Eastern Goldfields (B); 22___22,Stillwater (B); 223, Platbcrg-Klipriviersberg (B); 224, Abitibi belt (B); 225, Kylcna (B); 227. Kam Group (B); 230, Wawa (B); 233, Mount Roe (B); 234, Dcrdcpoort (B); 235, Rio das Velhas (B); 236, Vizien (B); 237, Kuhmo (B); 241, Kostomuksha (B); 242, BarlccYcllowdine (B); 244, Pickle Crow (B); 250, Forrestania-Lakc Johnston (B); 251, West Pilbara (B); 252, Steep Rock (B); 253, Pongola (B); 256, Balmer (B); 263, Olondo (B); 267, Verkhovtscvo (B); 270, Nondweni (B); 271, Lower Wanawoona-Upper Onvcrwacht (B); 273, Lower Onvcrwacht (B); 276, Coonterunah (B). Additional events in Prokoph et al. (2003) that update the listing in Ernst and Buchan (2(X)lb): A, Mctchosin (Coast Range Basalt Province) (B) [#339 in Buchan and Ernst (2003)]; B, Port Nollah-Gannakouriep (B) [combined events #57 and #59 in Ernst and Buchan (2001b)]; C, Wcstcrn North America (Gunbarrcl) and Windcrnlere (A) [combined events #63 and #77 in Ernst and Buchan (2001b)]" D, Giles-Bangemall (B) [combined events #84 and #88 in Ernst and Buchan (2001b)]" E, Frascr-Gnowangerup (B) Icombined events #101 and #102 in Ernst and Buchan (2001b)]; F, McRae Lake-Hadley Bay (B) [combined events #105 and #106 in Buchan and Ernst (2003)1; G, Molson-New Quebec Cycle-2 events (B) [combined events #150 and 151 in Ernst and Buchan (2001b)]; H, Avayalik (B) [#66 in Buchan and Ernst (2003)]; I, Kikkertavak (B) [event #47 in Buchan and Ernst (2003)]. Numerous events fiom Table 1 of Ernst and Buchan (2001b) are not listed here: "C" events, because of their uncertain plumehead origin, and miscellaneous events including no's. 278-304 that have age uncertainties greater than • My. Events which are underlined in the listing above can bc linked to a plume-head based on a LIP criteria (involving size of duration of the magmatic event; see text). The remaining events arc linked to a plume on the basis of chemistry or composition; most notably this includes grccnstone belts containing komatiites.

178

Fig. 3.3-2.

Chapter 3: Tectonism and Mantle Plumes Through Time

3.3. Large Igneous PIvvince Record

179

general, wavelet analysis (example in Fig. 3.3-2d) shows only weak cyclicity over limited time intervals. Possible exceptions are a c. 170 My cycle from 1500 Ma (possibly from Neoarchaean) to Present, a c. 330 My cycle from c. 2700 to 1500 Ma, and a cycle that progressively changes in length from 730 to 600 My over the interval from 2600 Ma to Present. Other weak cycles occurring over shorter intervals include a c. 105 My cycle in the early Proterozoic, a c. 230 My cycle in the Phanerozoic and a c. 250 My cycle in the late Archaean. The high frequency part of Fig. 3.3-2d exhibits a more complicated pattern because of the presence of non-persistent cycles ranging in length from 60 to 15 My (Prokoph et al., 2003). The most notable of these are a c. 16 My cycle present during much of the past 300 My and a 27 My cycle occurring from 2250 to 2000 Ma, and also in the oceanic portion of the record of the past 300 My. However, our data do not show a single dominant cyclicity. Cycles of c. 26, c. 35, c. 273, and c. 800 My were reported in a previous time series analysis of high-Mg units in LIPs by Isley and Abbott (2002). With the exception of the 26 My cycle, the correlations with our results (Fig. 3.3-2d; Prokoph et al., 2003) are not strong. As illustrated by the four curves in Fig. 3.3-2a, varying the criteria that are used for inclusion of LIPs in the data base can have a dramatic effect on the observed pattern of LIPs. Therefore, we infer that the difference between our result and that of Isley and Abbott (2002) is primarily a result of our different (expanded) LIP data base. As the LIP data base becomes more complete and better dated, time series analyses should converge on more consistent results. Summary

Large Igneous Provinces (LIPs) represent dramatic magmatic events of large volume and short duration. They punctuate Earth's history on average every 20 My (continental LIPs) or probably every 10 My (combined continental and oceanic LIPs). In general, the surface exposure of Cenozoic and Mesozoic LIPs is dominated by flood basalts, whereas in the Palaeozoic and Proterozoic, more widespread erosion has exposed the plumbing system

Fig. 3.3-2. Distribution of plume events in time: (a) Cumulative frequency curves (after Fig. 4 in Buchan and Ernst, 2002b) for subsets of the database based on assessment of reliability of the links with a plume, "A", "B", and the 2o age uncertainty -t-50 My, -t-20 My (2o-). The steeper the curve the more numerous the events. No events occur where the curve is horizontal. The dotted curve between 150 Ma and the Present is based only on data from continental LIPs. (b) Bar diagram showing spectrum of "A" and "A + B" subsets of the data (modified after Ernst and Buchan, 2001b, 2002b). Events with age uncertainty ~< 20 My are indicated with solid lines whereas those with 20-50 My uncertainty are located with a dashed line. (c) Potential plume clusters (modified after Ernst and Buchan, 2001b, 2002b) to include only events rated "A" and "B". (d) Wavelet analysis using Morlet wavelet with scaling factor l = 10 (Prokoph and Barthelmes, 1996) of LIP "A + B" subset with a <~ 10 My 2o" uncertainty (120 events) (after Prokoph et al., 2003). The most significant cycles are labelled with their approximate peak frequencies.

Chapter 3: Tectonism and Mantle Plumes Through Time

180

consisting of dykes, sills and layered intrusions. In the Archaean, LIPs may be represented by certain classes of plume-related greenstone belts, particularly those containing komatiites (see also sections 4.3 and 4.4). LIPs are linked to plumes rising from the deep mantle, and generally have a semicontinuous distribution in time and space. However, some may cluster in "superplume events", some of which are associated with continental breakup (section 3.2). Time series analysis indicates only weak cyclicity, most notably of wavelengths c. 170, c. 330, and 730-600 My, that is usually restricted to limited portions of the record.

3.4.

EPISODIC CRUSTAL GROWTH DURING CATASTROPHIC GLOBAL-SCALE MANTLE OVERTURN EVENTS

D.R. NELSON The granite-greenstone association is characterised by linear or arcuate belts of predominantly mafic volcanic rocks (or greenstones) that are commonly in fault contact with voluminous tonalitic, trondhjemitic, granodioritic and/or (TTG) rocks (cf. Condie, 1981) (sections 2.5-2.7, 4.3 and 4.4). As formation of granite-greenstone crust was largely confined to the Archaean era, the processes responsible for its creation may have differed substantially from those operating today. The dominant lithologies represented within greenstone belts (sections 2.3 and 2.4) are ultramafic, high-Mg and tholeiitic basaltic lavas (section 4.3). These are commonly interbedded with or conformably overlain by felsic volcanic (section 4.4) and/or chemical and clastic sedimentary rocks. Komatiites are unusual high MgO (~> 18 wt.%), low viscosity ultramafic lavas that were erupted at high temperatures (1400-1700~ and that were derived by high degrees of partial melting of hot, upwelling mantle peridotite, at depths of ~> 120 km (Nisbet et al., 1987; Herzberg et al., 1988; Bickle, 1993) (see also section 4.3). These lavas, commonly found in greenstone belts in close association with thick high-Mg and tholeiitic basalt sequences, are largely (with a few exceptions) unique to the Archaean era. Large volumes of granite-greenstone crust, including that of both the Yilgarn craton of Western Australia and of the Superior craton of North America (section 2.4), were formed between 2760 and 2620 Ma. Comparison of the available U-Pb zircon dates for volcanic and plutonic rocks of the Eastern Goldfields Province of the Yilgarn craton, and of the Abitibi and Wawa Subprovinces of the Superior craton (Fig. 3.4-1), indicates that greenstone volcanism in these regions of both cratons occurred between 2740 and 2680 Ma. The time of crystallisation of the granitic rocks in the eastern Yilgarn also overlaps with those in the southern part of the Superior craton (Fig. 3.4-1). Remarkably, komatiitic lavas within the greenstones of both the Yilgarn and Superior cratons have been dated precisely at 2705 Ma. One possible explanation for these geological similarities is that the Yilgarn and Superior cratons represent dispersed fragments of a once-united craton. If this were the case, details of the geological evolution in both cratons, including the timing of magmatism, The Precambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

3.4. Episodic Crustal Growth

181

deformation and metamorphism, should be identical. In detail however, the major episode of plutonism in the southern Superior craton predated that in the eastern Yilgarn craton by more than 15 My. Widespread trondhjemite, tonalite and granodiorite intrusions within the southern part of the Superior craton mostly date from between 2695 and 2680 Ma, whereas crystallisation ages determined for regionally extensive monzogranites of the Eastern Goldfields Province of the Yilgarn craton are generally between 2665 and 2650 Ma, with few granitoid rocks older than 2675 Ma. Episodes of regional compression, strike-slip deformation, the emplacement of late tectonic syenites, metamorphism, and deposition of clastic metasedimentary belts all commenced and continued later in the eastern part of the Yilgarn craton compared to the southern part of the Superior craton (Nelson, 1998a, and references therein). Compression within the southern part of the Superior craton had ceased by 2670 Ma, whereas in the eastern part of the Yilgarn, closure of a series of continental margin rift basins had only just commenced at this time, and deformation resulting from the closure of these rifts was not completed until ~< 2645 Ma. Furthermore, plutonism in the southern part of the Superior craton is mostly tonalitic to granodioritic in composition, whereas throughout the eastern Yilgarn it is dominantly monzogranitic, with tonalitic to granodioritic compositions comprising only a volumetrically minor component. These differences preclude the possibility that parts of the Yilgarn and Superior cratons could have been formed along the same tectonically active continental margin. If the Yilgarn and Superior cratons were not united during the Late Archaean, then the most plausible alternative explanation for the synchroneity of volcanism, and of the eruption of komatiites at 2705 Ma, within the greenstones of these two cratons is that this was due to a global-scale magmatic process, possibly related to catastrophic convective overturn of the Earth's mantle. About three times more heat was generated by radioactive decay within the Earth during the Archaean than at present (O'Nions et al., 1978). This was likely to have resulted in a more chaotic mantle convection r6gime during the Archaean compared with the present-day (see also section 3.6). Chaotic convective mantle overturn is likely to be manifest at the Earth's surface by the eruption of magmas derived from hot upwelling mantle, such as komatiitic lavas. Large-scale mantle upwelling episodes (see also section 3.2) may have resulted in synchronous ultramafic and mafic volcanism on a number of dispersed cratons. The compressive stresses induced in the Earth's tectonic plates following such convective overturn episodes will be eased by subduction and plate collision. Because each craton will have occupied its own unique tectonic environment, plutonic, deformation and metamorphic events in response to such mantle overturn episodes will not necessarily occur synchronously on affected cratons. Global-scale catastrophic magmatic events may have played an important role in crustformation throughout the Archaean. However, a comparison of available U-Pb zircon data from the 3.5 to 3.1 Ga granite-greenstone basements of the Pilbara and Kaapvaal cratons indicated a high degree of disparity in the timing of major granite-greenstone crustforming events on both cratons (Nelson et al., 1999). There was no clear evidence for synchronous magmatic episodes common to both cratons during this interval. This implies that magmatism was comparatively localised during the Early Archaean (see also section 3.6) and was not associated with large-scale mantle overturn episodes, as was inferred

182

Fig. 3.4-1.

Chapter 3: Tectonism and Mantle Plumes Through Time

3.5. Unusual Palaeoproterozoic Magmatic Event

183

for formation of 2.7 Ga granite-greenstone crust. It is possible that the Late Archaean era represented a time of transition, during which catastrophic convective overturn episodes may have played a role in crust-formation as modern-day plate tectonic processes became increasingly dominant (see also discussion in section 3.6).

3.5.

AN UNUSUAL PALAEOPROTEROZOIC MAGMATIC EVENT, THE ULTRAPOTASSIC CHRISTOPHER ISLAND FORMATION, BAKER LAKE GROUP, NUNAVUT, CANADA: ARCHAEAN MANTLE METASOMATISM AND PALAEOPROTEROZOIC MANTLE REACTIVATION

B.L. COUSENS, J.R. CHIARENZELLI AND L.B. ASPLER

Introduction At c. 1.83 Ga, the western Churchill Province of northern Canada witnessed one of the largest ultrapotassic magmatic events in Earth history. Minette dykes injected across c. 240,000 km 2, and voluminous minette flows and pyroclastic deposits accumulated in numerous continental sub-basins, collectively comprising the Christopher Island Formation (CIF) of the Baker Lake basin (Fig. 3.5-1) (see also section 7.6). Characterised by high volatile contents, incompatible element enrichment, and distinctive isotopic compositions, ultrapotassic rocks are widely accepted as melts of metasomatised lithospheric mantle (Mitchell and Bergman, 1991; Foley, 1992). Thus they are valuable probes to as-

Fig. 3.4-1. (a) Comparison of U-Pb zircon geochronology data for the granite-greenstone components of the Yilgarn craton, Western Australia, and the Superior craton of Canada. Volcanism in both regions occurred synchronously, whereas plutonism in the Eastern Goldfields region of the Yilgarn craton culminated at c. 2665 Ma, more than 15 My after that in the Abitibi-Wawa Subprovinces of the southern part of the Superior craton, which occurred mainly between 2700 and 2680 Ma. (b) Schematic diagram showing the possible effect of global catastrophic convective overturns of the Earth's mantle during the Archaean on cratons located in different tectonic settings. Nelson (1998) postulated that global-scale catastrophic mantle overturn events between 2780 and 2670 Ma caused the eruption of voluminous flood basaltic lavas of the Fortescue and Ventersdorp sequences onto the stabilised Pilbara and Kaapvaal cratonic platforms, the rifting of the eastern margin of the proto-Yilgarn craton, and eruption of a bimodal sequence of felsic and ultramafic lavas (now preserved as the greenstones of the Eastern Goldfields Province), and eruption of ultramafic, mafic and felsic lavas of the Abitibi and Wawa Subprovince greenstones onto young, thin crust adjacent to the southern margin of a proto-Superior craton. It was argued that the plate tectonic settings occupied by these different Archaean cratons responded differently following the disruption caused by the mantle overturn events. Such global-scale mantle overturn events may have played an important role in crust-formation on the Earth during the Priscoan and Archaean eras, on the Moon, and on Venus until about 700 Ma ago (see section 1.2). Diagrams adapted from Nelson (1998). The Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

184

Chapter 3: Tectonism and Mantle Plumes Through Time

Fig. 3.5-1. Simplified geological map, western Canadian Shield. Inset left: simplified stratigraphic column of the Baker Lake Group, showing interstratification of sedimentary (patterns) and volcanic (CIF, grey) rocks.

sess the evolution of the lithospheric mantle beneath cratons and cratonic tectonic events (e.g., Lambert et al., 1995; Canning et al., 1998; Feldstein and Lange, 1999; Peccerillo, 1999; Wannamaker et al., 2000; Cousens et al., 2001). Predominantly volcanic rocks of the CIF are part of the Baker Lake Group, the lowest unit of the Dubawnt Supergroup (Fig. 3.5-1, inset) (Gall et al., 1992). The Dubawnt Supergroup (see also section 7.6) consists of continental siliciclastic rocks and intercalated volcanic rocks deposited between 1.84 and 1.72 Ga (Rainbird et al., submitted). The CIF has proven difficult to date precisely, but recent efforts show that volcanism extended from c. 1.84 to 1.79 Ga (Rainbird et al., submitted), coincident with collisional and

3.5. Unusual Palaeoproterozoic Magmatic Event

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post-collisional processes in the Trans-Hudson orogen, on the southern flank of the western Churchill Province (Fig. 3.5-1).

Regional Distribution of the CIF Samples of CIF lavas and dykes were collected from the Angikuni Lake-Rack Lake, MacQuoid-Gibson Lake, Kaminak Lake, Baker Lake, Kasba Lake-Lake Athabasca (Martin Formation), and Whitehills Lake areas. Data from this suite are compared to results from the Dubawnt Lake area (Peterson and Rainbird, 1990; Peterson and LeCheminant, 1993; Peterson, 1994; T.D. Peterson et al., 1994; Peterson and LeCheminant, 1996) (Fig. 3.5-1). CIF lavas and dykes from all areas range from phlogopite-bearing mafic minettes to aphyric or feldspar-bearing felsites. In the Dubawnt, Kamilukuak and Baker Lake areas, a felsic-mafic-felsic minette sequence is interstratified with siliciclastic rocks (Rainbird and Peterson, 1990; Peterson, 1994; Rainbird et al., submitted). Near Baker Lake, Archaean basement is usually overlain by orange to red-weathering, potassium feldspar-bearing felsite with only minor phlogopite. Sanidine porphyries, including dusty potassium feldspar crystals up to 1 cm long, also occur (Smith, 2001). Overlying mafic minettes are rich in phlogopite and clinopyroxene. Upper felsites are grey to orange, slightly potassium feldspar-phyric flows and domes. In the Angikuni Lake area, Baker Lake Group rocks outcrop in two northeast-trending sub-basins (Fig. 3.5-1) that extend from northern Angikuni Lake (Aspler et al., 1998, 1999). Mafic units commonly include up to 30% phlogopite and clinopyroxene phenocrysts in a potassium-feldspar-rich matrix. Felsic rocks contain variable proportions of feldspar phenocrysts, commonly including minor corroded phlogopite crystals. We sampled core from three holes drilled by WMC International Ltd. near the centre of the easternmost sub-basin at Rack Lake (informal name, Fig. 3.5-1) (Cousens, 1999). Hole 94-2 consists of over 170 m of poorly-phyric, phlogopite-clinopyroxene, variably carbonate-rich mafic minette flows of uniform compositions, and Hole 94-1 includes over 140 m of mafic pyroclastic minette. Hole 95-1 intersected a 500 m-thick section of siliciclastic rocks (Angikuni Formation) that intervenes between Archaean basement and the CIF, yet contains CIF-like detritus (Aspler et al., 2002a). In the MacQuoid-Gibson and Kaminak Lake areas (Beaudoin, 1998; Sandeman et al., 2000b), the CIF is represented exclusively by dykes. CIF dykes at MacQuoid-Gibson can be split into three types: (1) hornblende-phlogopite-plagioclase spessartites, (2) typical CIF phlogopite minettes, and (3) rare black, poikilitic, phlogopite-potassium feldspar dykes exemplified by the diamondiferous Akluilak dyke (MacRae et al., 1995; Armitage, 1998). A CIF dyke was sampled from the Snowbird tectonic zone at Kasba Lake (Hanmer et al., 1995). Farther southwest, along the north shore of Lake Athabasca (see also section 8.3), potassic to sodic dykes, sills, and flows of the Martin Formation are similar geochemically to the CIF (Ashton et al., 1999; Hartlaub, 1999; Morelli et al., 2001). Martin Formation volcanic rocks include plagioclase and sanidine phenocrysts, accompanied by variably altered clinopyroxene, but lack phlogopite.

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Chapter 3: Tectonism and Mantle Plumes Through Time

Mineral Chemistry

We selected ten samples, which represent the different textural and compositional characteristics of the suite, for microprobe analysis. Our data from phlogopite and feldspar phenocrysts mirror previously published data for the CIF (LeCheminant et al., 1987). Clinopyroxene is a ubiquitous macrocryst in CIF minettes, and crystals are commonly zoned from Cr-rich diopside cores to Fe-Na salite rims (Fig. 3.5-2). Three analyses of groundmass clinopyroxene grains are all Cr-diopsides. We interpret the zoned diopside-salite grains to be mantle xenocrysts, perhaps disaggregated wall rocks from within or above the

Fig. 3.5-2. Variations in clinopyroxene chemistry in CIF lavas and dykes (this study, plus "CIF" field from LeCheminant et al., 1987). Note Na- and Fe-enrichment of rims.

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zone of melting. A clinopyroxene separate from Christopher Island dyke PHA-97-T138A from the Kaminak Lake area yielded initial 87Sr/86Sr of 0.70309 and eNd1830 of --0.8 (Beaudoin, 1998), yet whole-rock isotope ratios for this sample are 87Sr/86Sr = 0.70288 and eNd1830 = --8.1 (Cousens, 1999). The isotopic difference between bulk rock and clinopyroxene separate support the hypothesis that the clinopyroxene is xenocrystic. Common groundmass phases include rutile, apatite, Ti-poor magnetite, titanomagnetite, alkali-poor amphibole, pyrite, quartz, and calcite. Calcite occurs as large clots in many samples, but is particularly common in lavas from Rack Lake drill Hole 94-2. The variety of minor phases suggests that normally incompatible trace elements (elements not partitioned into crystallising mineral phases in basaltic systems) would not behave incompatibly during CIF magma evolution.

Geochemistry CIF lavas span a huge range in major and trace element characteristics (Table 3.5-1; Figs. 3.5-3 and 3.5-4) (T.D. Peterson et al., 1994; Cousens et al., 2001). Major element compositions range from mafic minettes to feldspar-rich felsites (SIO2 = 41-72 wt.%). K20/NazO is variable in CIF lavas, ranging from 0.01 to 102:16% of CIF lavas are sodic, 20% are potassic, and 64% are ultrapotassic (Fig. 3.5-3, inset). Whole-rock chemical compositions cluster about the dividing line between lamproites (Group I) and lamprophyres (Group III) in plots of CaO versus A1203 (see Foley et al., 1987), CaO versus SiO2 and a ternary plot of A1203-KaO-MgO (Fig. 3.5-4a) (Bergman, 1987). However, TiO2 contents in the CIF place them in a low-Ti subgroup that is usually associated with ultrapotassic volcanism in recent to modern subduction environments (e.g., Roman Province, Spanish lamproites). In a plot of SiO2 versus peralkalinity (Fig. 3.5-4b), CIF lavas and dykes fall within the Roman Province field, which also includes minettes (Mitchell and Bergman, 1991). Primary carbonate is an important phase in CIF volcanic rocks: CO2 contents are as high as 12.5% (by weight) in mafic minettes, whereas felsites generally have < 4% CO2. Cr and Ni abundances in the most MgO-rich CIF rocks are typical of mantle-derived magmas, ranging from 1000 to 400 ppm Cr and 400 to 200 ppm Ni. Sc and V abundances are lower than expected for "primary" mantle-derived magmas, with maxima of 180 ppm V and 25 ppm Sc, but are typical of minettes (Mitchell and Bergman, 1991). Ni, Cr, Sc and V all behave as compatible trace elements in CIF rocks, presumably partitioning into phases such as olivine, Cr-spinel, clinopyroxene and Fe-Ti oxides, respectively. TiO2 abundances also decrease with increasing fractionation, as measured by the Mg# or SiO2 (Fig. 3.5-5). In addition, P205 abundances drop dramatically as the rocks become more fractionated, accompanied by scattered, but generally decreasing Th and U abundances, all suggesting that apatite is also a fractionating phase (Fig. 3.5-5). Rare earth element abundances remain constant or decrease as the Mg# decreases, which also may be a function of apatite fractionation. Thus, with the fractionating assemblage of clinopyroxene, potassium feldspar, titanomagnetite, phlogopite, and apatite, only A1203, Na20 and SiO2 increase in abundance as the magmas become more evolved. Crustal contamination during magma ascent conceiv-

-

Table 3.5-1. Representative chemical and Nd isotopic analyses of CIF lavas and dykes Sample Area UTM Zone Northing Easting

97-AN-7 Angikuni L.k. 13 6895700 459750

98-RL-2 Rack Lake 13 6922939 503436

SiO2 (wt.%)

45.65 11.34 0.08 9.28 8.37 1.27 4.08 0.56 0.42 8.20

64.16 13.72 0.05 1.87 2.28 4.52 4.66 0.33 0.25 3.1 1

50.82 11.80 0.14 10.33 8.57 1.49 4.42 0.84 1.29 8.12

52.79 11.53 0.1 1 9.23 4.69 1.48 5.91 0.97 0.98 7.53

55.90 13.00 0.09 7.32 5.11 3.00 7.00 0.58 0.65 6.30

46.95 12.20 0.12 8.05 7.42 3.65 2.43 0.66 0.5 1 7.81

9.43

3.92

0.80

3.20

1.50

5.60

98.68 0.01 5.50 3.92 5 140 20 480 74 I900 900 37 80

98.87 0.06 2.81 I .05 8 147 11 395 129 3118 38 11 17

98.62 0.01

98.42

100.45

95.40

-

0.49 2.71 18 279 26 650 25 1 6753 289 29 60.07

A1203

MnO MgO CaO Na20 K20

Ti02 p205

Fe2Ogt LO1 Total S COZ H?O Nb (PPW Zr Y Sr Rb Ba Cr co cu

98-TXH-461A MacQuoid 15 7038905 470250

00 30

0.80 12 409 31 1868 186 -

39 8

RH98-275 Athabasca 13 6596000 3 19900

-

ZB99-398 Whitehills 14 7 170673 641596

-

0.30 1.20 7 181 13.2 1327 377 2800 31 1 26 37

00-RAT-TL 17 Baker Lake 14 7065700 603300

-

3.51 2.09 7 148 20 1531 113 9319 104 26 136

2

B z

?

3 3.

2

29.

3i ;a b

$2

fr

39

aY

$

Table 3.5- 1 (continued). Sample Ni Sc

v

Zn La Ce Pr Nd Sm Eu Gd Tb DY Ho Er Tm Yb Lu Th U Hf Ta Pb EN^ 1830

97-AN-7 200 22 150 110 40.26 80.95 9.50 37.22 6.93 1.89 5.31 0.68 3.82 0.66 1.97 0.24 1.57 0.21 9.68 2.04 3.29 0.31 -

-8.6

98-RL-2 30 6 48 46 44.00 84.15 9.79 36.73 6.15 0.26 3.87 0.42 2.33 0.32 0.99 0.11 0.74 0.11 12.41 4.84 4.02 -

98-TXH-461A 206 21 146 110 130.78 242.93 33.77 129.88 21.36 5.31 14.20 1.57 8.03 1.17 3.10 0.33 1.98 0.28 16.14 2.96 8.85 0.58

-

-

-7.3

-7.0

RH98-275 179 16 135 82 66.62 130.09 16.49 61.80 1 1.98 2.93 8.22 1.09 4.97 0.89 2.11 0.28 1.77 0.26 35.15 10.76 7.51 1.20 52 -8.8

ZB99-398 142 14 98 85 53.0 110.0 13.00 50.0 9.00 2.10 5.90 0.66 3.00 0.5 1 1.30 0.19 1.20 0.20 18.00 7.80 4.20 0.61 42 -7.6

00-RAT-TL 17 80 21 131 91 52.21 114.21 14.15 55.56 11.48 2.88 7.18 0.92 4.17 0.71 1.80 0.26 1.70 0.22 12.96 19.37 3.25 0.33 2 -8.0

Notes: Major elements and Nb through Zn by fused disk XW La through Pb by dissolution ICP-MS. See Cousens (1999) for details. Major element precision < 5% except for S (15%). XRF trace elements ilo%, and ICP-MS i5%. The complete data set is available from the senior author and will soon be published as a Geological Survey of Canada Open File.

sF

-s2

5 a2 +,a

%

s.

3

Oc

3

2. 2 5

-

\O w

190

Chapter 3: Tectonism and Mantle Plumes Through Time

Fig. 3.5-3. K20 versus SiO2 (weight %) for CIF lavas and dykes of this study and Dubawnt Lake (LeCheminant et al., 1987; Peterson and LeCheminant, 1993; Peterson et al., 1994; T. Peterson, 1999, pers. comm.). Inset: histogram of K20/Na 20 ratios in CIF rocks. Y-axis is fraction (fr) of total number of samples (221). ably could have contributed to these systematics, but the estimated rapid ascent rates of hours to days (LeCheminant et al., 1987; Peterson and LeCheminant, 1993) imply inadequate time for significant interaction with wall rock. It is clear from plots of P205, K20, and TiO2 versus SiO2 that a range of primary magma compositions exists in the CIF (Fig. 3.5-5). At SiO2 < 50%, CIF dykes from the MacQuoid-Gibson Lake and Kaminak Lake areas exhibit a large range of P205 and K20 contents. Dykes and lavas of the Martin Group (see also section 8.3) are enriched in TiO2 and Sc compared to all CIF, but otherwise fall in the CIF range. Primitive-mantle-normalised (Sun and McDonough, 1989) incompatible element patterns for CIF lavas and dykes from the study area are shown in Fig. 3.5-6. Incompatible element patterns in CIF samples are subparallel (in particular, rare earth element (REE) patterns, Fig. 3.5-6 inset) and exhibit pronounced depletions in Nb, Ta, and Ti, and large enrichments in Rb, Ba, Th, U, K, and the light REE relative to primitive mantle. CIF rocks

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Fig. 3.5-4. Ultrapotassic rock classification diagrams. (a) A1203-K20-MgO (weight %) ternary plot distinguishing lamproites, lamprophyres, and kimberlites (Bergman, 1987). (b) Peralkalinity index (molar) versus SiO2 (weight %) distinguishing Roman Province (RPT), Leucite Hills (LHT), and Toro-Ankole (TAT) type potassic lavas (Mitchell and Bergman, 1991). Most minettes, including young subduction-related minettes from the Mexican Volcanic Belt (Luhr, 1997), plot in the RPT field. Sources of Dubawnt Lake data same as in Fig. 3.5-3. from the Angikuni Lake, Rack Lake and Baker Lake areas consistently have low incompatible element abundances relative to samples from the Kaminak Lake and MacQuoidGibson Lake Akluilak-type dykes. Other areas exhibit a range of incompatible element enrichment. The Martin Formation rocks have REE abundances and patterns similar to other CIF rocks, but are not as depleted in Ti or as enriched in the large ion lithophile elements

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Chapter 3: Tectonism and Mantle Plumes Through Time

Fig. 3.5-5. Plots ofTiO2, P2 05 (weight %), La, Th, Cr and Ba (weight parts per million) versus SiO2 (weight %) demonstrating magmatic evolution of CIF magmas. Additional data from Dubawnt Lake (sources from Fig. 3.5-3) and Lake Athabasca area (Morelli et al., 2001). (Rb, Ba, U and K). These incompatible element patterns are typical of subduction-zone magmas (e.g., Luhr, 1997), suggesting that the source of Christopher Island magmas was mantle-modified by subduction-related fluids.

Radiogenic Isotopes Whole-rock analyses display a remarkable uniformity in Nd isotope composition (Fig. 3.5-7a); 75% of the eyd 1.830 values are between - 9 and - 7 . The lack of systematic variation in end 1830 with SiO2 suggests that crustal contamination was not a major process affecting these magmas (Cousens et al., 2001). Calculated depleted mantle model ages (TDM) for Christopher Island volcanic rocks, based on the 147Sm/144Ndratio measured in each sample, range from 2671 to 3150 Ma (average, 2861 Ma). Similar to examples described from elsewhere (e.g., Van Kooten, 1981; Nelson, 1992; Canning et al., 1998), ultrapotassic rocks of the CIF have isotopic compositions consistent with mantle metasomatism and subsequent long-term storage in the lithospheric mantle. In contrast, calculated initial 87Sr/86Sr in whole-rock powders are highly variable, ranging from 0.64088 to 0.76076 (Figs. 3.5-7a, b). Some of the samples have initial 87Sr/86Sr lower than that of the Earth's initial 87Sr]86Sr at 4.5 Ga (Tilton, 1988). Because this is

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Fig. 3.5-6. Primitive mantle (Sun and McDonough, 1989)mnormalised incompatible element plot for selected CIF lavas and dykes (black lines) compared to all available CIF analyses (grey field). Symbols as in previous figures. Additional data from Dubawnt Lake (sources from Fig. 3.5-3) and Lake Athabasca area (Morelli et al., 2001). Inset: primitive mantle-normalised rare earth element patterns for all available CIF analyses. Note uniformity of rare earth element patterns, prominent negative Nb, Ta, and Ti anomalies, depletions in P, and enrichment in Rb, Ba, and Th relative to the light rare earth elements (subduction signature). highly unlikely, these samples could not have remained closed systems since crystallisation, and some loss of Sr or gain of Rb must have occurred in the last 1.83 Ga to raise the Rb/Sr ratio. Samples with very high initial 87Sr/86Sr may have lost Rb or gained Sr subsequent to crystallisation, thereby lowering the Rb/Sr ratio. Consequently, calculated initial 87Sr/86Sr are in some cases over-corrected or under-corrected for radiogenic ingrowth (e.g., Cousens et al., 1993). Rb (and K) mobility is not unexpected in volcanic rocks that have remained near the surface of the Earth for nearly 2 billion years. Separation of the carbonate from silicate phases in whole-rock powders shows that as much as 75% of the whole rock Nd is concentrated in the carbonate phase (Cousens, 1999). Fig. 3.5-8 plots the abundances of incompatible elements in the carbonate versus silicate phases in a CIF flow from Rack Lake. The light to middle REE and Sr are concentrated in the carbonate phase, whereas the large ion lithophile elements (Rb, Ba and K) and high field strength elements (Th, Nb, Ta, Zr and Hf) are concentrated in the silicate phases. In particular, Rb is concentrated in phlogopite and Sr in the carbonate phase. Whereas initial end in the silicate and carbonate phases overlap within analytical error, calculated initial 87Sr/86Sr are dramatically different. The silicate and carbonate phases are not in iso-

194

Fig. 3.5-7.

Chapter 3: Tectonism and Mantle Plumes Through Time

3.5. UnusualPalaeoproterozoic Magmatic Event

195

topic equilibrium, and the initial ratio calculated for the silicate phase is below that of the initial ratio of the Earth (i.e., unlikely to be a true magmatic value). The carbonate has extremely low Rb/Sr so that corrections for 87Sr ingrowth are minimal, and Sr concentrations are high enough that meteoric water exchange has likely been minimal. Analysis of the carbonate phase alone in eight samples reveals a much smaller range of initial 87Sr/86Sr ' between 0.7045 and 0.7083 with a mean of 0.7067 (Fig. 3.5-7b). This coincides with the average 87Sr/86Sr in the whole-rock samples, and hence the carbonate analyses appear to provide the best estimates of magmatic values. We interpret the abundance of carbonate in mafic to intermediate minettes and the geochemistry of the carbonate phase to indicate that the carbonate is a primary magmatic phase (Cousens, 1999). The indistinguishable Nd isotopic compositions in the carbonate and silicate phases and the low 87Sr/86Sr of the carbonate are strong evidence that the carbonate is largely magmatic. Our experience analysing Archaean rocks that have suffered pervasive, post-crystallisation CO2 alteration shows that silicate and carbonate phases rarely have the same calculated end at the age of crystallisation. Pb isotopic data for CIF dykes and flows are unradiogenic compared to most other ultrapotassic suites (initial 2~176 -- 11.2-14.6) (T.D. Peterson et al., 1994; Cousens, 1999). Pb isotope ratios and Ce/Pb variations in CIF magmas are consistent with mixing between two components, depleted mantle and ocean floor sediments (or fluids derived from a sedimentary source). Stratigraphic Variations in Geochemical Composition

Samples from stratigraphic sections through the Baker Lake basin at Thirty Mile Lake and the Aniguq River (Rainbird and Hadlari, 2000; Rainbird et al., submitted) demonstrate that although SiO2 contents vary from 46 to 68%, rare earth patterns are extremely similar and eNd1830 values show no correlation with La/Sm, SiO2 or stratigraphic height. The rocks analysed include typical CIF sills, dykes, plugs and flows, as well as a less common plagioclase-hornblende monzonite dyke. Thus where stratigraphic control exists, there is no evidence for a preferred liquid line of descent; individual flows are the product of small melt batches from an isotopically homogeneous source that evolved independently from earlier and later melt batches. Although magma mixing between some of these melt batches may have occurred at shallow levels, evidence of a long-lived magma chamber is lacking.

Fig. 3.5-7. (a) Initial ENd versus initial 87Sr/86Sr for CIF lavas and dykes of this study compared to Dubawnt Lake and various CIF (Keewatin) dyke analyses (Peterson et al., 1994). 2o uncertainty in 87Sr/86Sr is smaller than symbols. Inset: expanded 87Sr/86Sr scale to show all isotopic analyses of CIF rocks. Note that initial 87Sr/86Srof Earth (4.5 Ga) is c. 0.6990 and that two Baker Lake analyses have 87Sr/86Srbelow that value. (b) Histogram of 87Sr/86Sr in CIF whole rock (black) and carbonate fraction only (light grey) analyses from this study.

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Chapter 3: Tectonism and Mantle Plumes Through Time

Fig. 3.5-8. Normal mid-ocean ridge basalt (N-MORB, Sun and McDonough, 1989)mnormalised incompatible element abundances, Sr and Nd isotope ratios in carbonate and silicate fractions of Rack Lake sample 94-2-577 A. Note that eNd values overlap within analytical error (+0.8).

Models for the Origin of the CIF The remarkable uniformity of incompatible element patterns and eNd 1830 values in rocks of the CIF throughout the western Churchill Province indicate that they originated from a common mantle source. Several constraints can be applied to the age and origin of the CIF source. High Ni, Cr, and MgO contents in mafic minettes require that they are mantle-derived melts. CO2 is an important component in magma genesis, as reflected in its high abundance in mafic to intermediate minettes. Incompatible element patterns imply that this mantle source was metasomatised by fluids (silicate or H20/CO2) carrying a subduction-related trace element signature. Nd isotopic compositions are remarkably constant throughout this large area, requiring an isotopically uniform source. Nd model ages for Christopher Island magmas cluster closely at c. 2.8 Ga, similar to the average model age of Neoarchaean volcanic rocks from the western Churchill Province (T.D. Peterson et al., 1994; Theriault and Tella, 1997).

Palaeoproterozoic metasomatism of the lithospheric mantle Emplacement of the CIF was coincident with collisional events in the Trans-Hudson orogen, and it is possible that the CIF source formed during subduction of oceanic crust and sediments beneath the western Churchill Province prior to 1840 Ma. If so, a mechanism is required to lower uniformly the eNd of the mantle source (lithosphere or asthenosphere) to c. - 9 from c. + 5 beneath a large volume of the mantle inboard of the subduction zone.

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Conceivably, subduction of sediment derived from western Churchill upper crust at 1.85 Ga with an average eNa of --8 (Theriault and Tella, 1997) could have produced metasomatic fluids with highly negative eNa values, but only if this sediment was the sole source of Nd incorporated into the fluids. This is rarely the case in modern oceanic island arcs, where the contribution to the mantle wedge from the subducting slab (sediment plus crust) is < 15% for most elements (e.g., Hawkesworth et al., 1991). Only positive eya values have been reported for arc-related volcanic and plutonic rocks of the Trans-Hudson orogen (e.g., Stern et al., 1995a; Whalen et al., 1999; MacHattie, 2002), and hence sediments with highly negative eNa do not appear to have contributed significantly to the Nd budget of the Trans-Hudsonian mantle wedge but were instead subducted deep into the mantle. Furthermore, subduction angles were steep enough to produce large batholiths along the flanks of the western Churchill (e.g., 1.99-1.92 Ga Taltson magmatic zone, 1.86-1.84 Ga Wathaman batholith; Corrigan et al., 2000; McNicoll et al., 2000; Hollings and Ansdell, 2002; MacHattie, 2002) but too steep to promote extensive metasomatism in the western Churchill interior. Palaeoproterozoic crust~mantle imbrication and melting

An innovative model for the origin of the CIF suggests that the Palaeoproterozoic Hudson granitoid plutons in the Baker Lake area, dated at 1.84-1.81 Ga, are directly related to the formation of the CIF source in the lithospheric mantle (Peterson et al., 2000, 2003). In the MacQuoid-Gibson Lake area, spessartite-type CIF dykes are commingled with Hudson granitoids (Sandeman et al., 2000b). Hudson granitoids share similar ages, incompatible element patterns and initial eya 1830 ( - 7 to - 1 3 ) with CIF minettes, even though the former are interpreted as lower crustal melts and the latter are clearly mantle-derived. The Peterson et al. model proposes that the CIF mantle source formed during the Trans-Hudson collision as a result of imbrication of the lower crust and lithospheric mantle, at which time K- and incompatible element-rich fluids migrated from the lower crust into the lithospheric mantle during lower crustal melting. This fertilisation of the lithospheric mantle induced mantle melting, producing the CIF minettes whose trace element and isotopic budget was dominated by the lower crustal contribution to the lithospheric mantle. Although the common geochronological and geochemical traits of these two dramatically different rock types are undeniable, we are sceptical that this mechanism can explain important aspects of CIF geochemistry and geology. First, carbonate minerals are extremely rare in rocks from the lower crust (e.g., Rudnick, 1992; Rutter and Brodie, 1992), yet the CIF mantle source must include significant CO2. Second, the lower melting point of the lower crust relative to mantle rocks would require that most of the lower crust be partially molten before the mantle would even begin to melt, and incompatible elements should strongly partition into the granitoid melts rather than infiltrate the lithospheric mantle. Third, mingling of Hudson granitoids occurs with early spessartite-type CIF dykes but not with the "regular" or Akluilak-type CIF dykes in the MacQuoid-Gibson Lakes area (Sandeman et al., 2000b; H. Sandeman, 2002, pers. comm.), so the field relationship between the Hudson granitoids and the CIF is limited to the earliest phase of CIF volcanism.

198

Chapter 3: Tectonism and Mantle Plumes Through Time

Archaean metasomatism of the lithospheric mantle We postulate that metasomatism of the lithospheric mantle occurred during the Archaean and that a large region of enriched lithospheric mantle evolved as an isotopically closed system until much later (c. 1.83 Ga), when shortening related to the Trans-Hudson orogen led to regional uplift and internal stretching of the western Churchill Province. This stretching is inferred to have produced thinning of the lithosphere, generating minette melts by adiabatic decompression, and creating mechanical pathways for rapid magma release. Thinning of the lithosphere also brings the asthenosphere closer to the crust/mantle boundary. Slab breakoff and foundering at the end of Trans-Hudson subduction may have amplified heat transfer to the lithosphere as asthenosphere rose rapidly through the resulting slab window (Hollings and Ansdell, 2002). The combined thermal effects allowed heat to be transferred into the lower crust at this time and potentially caused melting to produce the Hudson granitoids. Nd model ages for Christopher Island magmas cluster at c. 2.8 Ga, which is similar to the average model age of Archaean volcanic rocks from the western Churchill Province (T.D. Peterson et al., 1994; Theriault and Tella, 1997; Sandeman et al., 1998; Sandeman et al., 1999). Archaean mafic volcanic rocks commonly include a small subduction component, and Archaean felsic volcanic and plutonic rocks of the region are interpreted to be partial melts of Archaean lower crust or underplated oceanic crust (i.e., adakites; Drummond and Defant, 1990), indicating that very low-angle subduction was an important Archaean process in the western Churchill Province (H. Sandeman, 2002, pers. comm.). Thus, metasomatism of the lithospheric mantle, perhaps including sediment underplating, may have occurred during Archaean crust/lithospheric mantle formation. If the lower crust also inherited a subduction component, and hence evolved isotopically in a way similar to the lithospheric mantle, Hudson granitoids would carry the same isotopic signature as mantlederived CIF melts. One prediction of this model is that pre-CIE Palaeoproterozoic mafic rocks in the western Churchill Province might include an enriched mantle source component. Data from three such events, as well as a late Archaean amphibole lamprophyre (Cousens, 1999), confirm this prediction. The eNd values of c. 2.45 Ga Kaminak dykes, c. 2.19 Ga Tulemalu dykes, and c. 2.2 Ga basaltic flows in the Hurwitz Group (Happotiyik Member, Ameto Formation; Sandeman et al., 2000a) suggest that they include variable proportions of Christopher Island-type and more depleted mantle sources, implying that enriched lithospheric mantle sources existed well before 1.85 Ga (Cousens et al., 2001). Further, a recent geochemical study of the 1.86 Ga Wathaman Batholith indicates that granitoids in the northern edge of the batholith include an enriched mantle component similar to the source of the CIF (MacHattie, 2002). One other igneous suite in the western Churchill Province, the 2.1 Ga Griffin gabbro sills, does not include a contribution from a CIF-type enriched mantle source. However, field evidence for lateral transport of magmas at mid-crustal levels from a distant source is found in the Griffin gabbros, so these magmas did not interact with sub-western Churchill Province lithospheric mantle (Aspler et al., 2002b).

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Mantle Source Formation The extraordinary uniformity in eNd 1830 values and the parallelism of incompatible element patterns displayed by the CIF require that a large area of the western Churchill Province was underlain by an extremely uniform mantle source. Furthermore, melting only of enriched, rather than depleted, mantle is indicated. These observations are difficult to reconcile with petrogenetic models that derive alkalic magmas from enriched mantle domains distributed as discrete veins or clots within less enriched mantle (Menzies and Murthy, 1980; Francis and Ludden, 1990; Foley, 1992; Harry and Leeman, 1995; Esperanqa et al., 1997). Such models predict a spectrum of magma types, demonstrated in some ultrapotassic provinces (e.g., Van Kooten, 1981), ranging from highly alkalic (melts of enriched domains only) to less alkalic (melts of enriched and depleted mantle), depending on the degree of mantle melting and melt and wall-rock interaction. We propose that the extensive, uniform source of the CIF was a volatile-rich sheet trapped within (or at the base of) the subcontinental lithosphere. Speculatively, such a sheet might have formed as a result of a process analogous to fiat subduction, resulting from elevated Archaean geotherms and lower-density subducting slabs (sections 3.6 and 3.7). Fluids (silicate, H20-CO2, or both) driven from a relatively young, hot, subducting slab (sediments plus crust) possibly infiltrated the overlying mantle across a large area and these fluids then froze at the base of thickening lithosphere to form a sheet of enriched mantle. The existence of a melt and peridotite reaction zone at the base of the lithospheric mantle may have enhanced the formation of volatile-rich, low-degree melts that could infiltrate large volumes of overlying peridotite and thereby produce a widespread, pervasively metasomatised mantle source (Verni~res et al., 1997). Isolated from the convecting upper mantle, this sheet could then evolve isotopically until 1.83 Ga to yield the highly negative eNd values displayed by CIF magmas. Carbonate and phlogopite can co-exist in the mantle at pressures greater than 2.2 GPa (70 km) (Olafsson and Eggler, 1983). Between 1.5 and 2.2 GPa, amphibole and carbonate are stable phases in the lithospheric mantle. The relatively old (Sandeman et al., 2000b) hornblende, phlogopite-bearing spessartite-type dykes in the MacQuoid-Gibson area and the plagioclase-sanidine-clinopyroxene dykes and flows of the Martin Formation may be partial melts of metasomatised mantle at shallow levels in the amphibole-carbonate stability field, reflecting melting as the lithosphere began to thin at 1830 Ma. Continued thinning of the lithospheric mantle would bring mantle rocks with stable phlogopite and carbonate to pressures < 1.5 GPa (45 km) at > 1050~ where they would melt to produce phlogopite, sanidine and carbonate-rich magma of the more common CIF type.

Ultrapotassic Magmatic Events and Formation of Enriched Lithospheric Mantle During the Archaean Ultrapotassic volcanic rocks and dykes with Neoarchaean Nd model ages occur in the Wyoming craton, Greenland, the Baltic Shield, and the western Sao Francisco craton of

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Fig. 3.5-9. Initial end versus initial 87Sr/S6Srfor various ultrapotassic volcanic suites (Mitchell and Bergman, 1991; Luhr, 1997; Miller et al., 1999; Rosa et al., 1999; Rosa et al., 2000; Andronikov and Foley, 2001; Chung et al., 2001). Field patterns indicate age of basement rocks. K-1 and K-2 (diagonal ruled fields) are Type 1 and Type 2 kimberlites, respectively. DM indicates composition of modern depleted upper mantle. Dashed fields enclose CIF whole-rock (CIFs) and carbonate only (CIFc) analyses from this study.

Brazil. These rocks have incompatible element patterns and isotopic signatures similar to those of the CIF, and are all interpreted to have been derived from enriched lithospheric mantle that formed during the Archaean (Mitchell and Bergman, 1991; T.D. Peterson et al., 1994; Nikitina et al., 1999; Rosa et al., 1999; Buhlmann et al., 2000). Ultrapotassic rocks erupted through Archaean terranes typically have very negative end and unevolved 87sr/g6sr at the time of eruption (Fig. 3.5-9), unlike ultrapotassic rocks erupted through exclusively Proterozoic or younger basement rocks that have more radiogenic Sr (e.g., Spanish and West Australian lamproites). Thus Archaean mantle metasomatism produced sources with low Sm/Nd and Rb/Sr along with a distinctive subduction-related

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trace element signature. Younger metasomatised mantle sources apparently have higher Rb/Sr that may reflect the greater abundance of Rb-rich, high 87Sr/86Sr ' continentallyderived sediments undergoing subduction during the Proterozoic and Phanerozoic eras. Perhaps parts of most Archaean cratons are underlain by "low Rb-Sr"-type metasomatised lithospheric mantle similar to the western Churchill Province, and laterally widespread mantle enrichment due to fluid release during shallow subduction was a common Archaean event (see section 3.7). Although younger mantle enrichment events have occurred beneath post-Archaean continental crust, generally positive end values indicate that melting of the enriched source generally occurs soon after enrichment, and storage times are short (Miller et al., 1999; Righter, 2000; Andronikov and Foley, 2001; Chung et al., 2001). Thus lithospheric mantle enrichment during the Archaean was an important process that has produced isotopically distinct alkaline and potassic melts with a source history of long-term incompatible element storage.

3.6.

A COMMENTARY ON PRECAMBRIAN PLATE TECTONICS

P.G. ERIKSSON AND O. CATUNEANU Introduction

A strong polarisation between uniformitarian (e.g., Tarney et al., 1976; Dewey and Windley, 1981; Windley, 1993, 1995; Burke, 1997; de Wit, 1998) and transformist (e.g., Davies, 1992b; Goodwin, 1996; Hamilton, 1998) schools of thought has generally characterised the long-lived debate on Precambrian, and particularly Archaean plate tectonics (de Wit and Ashwal, 1997a). Within the voluminous literature on this debate, the uniformitarians have the advantage of numbers (Kusky, 1997), a consensus rightly viewed with a jaundiced eye (Hamilton, 1998) when used to justify application of an essentially modern plate tectonic paradigm to the Precambrian. Collections of interpretations, viewpoints and models in favour of Archaean plate tectonics (e.g., de Wit, 1998) must be judged objectively; they serve as collective wisdom rather than absolute truth. Geochemistry, reflecting magma reservoirs and melting conditions, is of little value in the debate (Burke, 1997; see, however, Polat et al., section 2.3, and Cousens et al., section 3.5). The application of Lyellian uniformitarianism has been inconsistent in many of these divergent publications. The null or Lyellian hypothesis (see Burke, 1997), that plate tectonics being responsible for most heat loss from the mantle today also applied in the Precambrian, is often invoked (e.g., Windley, 1993). Apart from uncertainties related to the K/U ratio in the mantle, almost all workers support radiogenic heat having been 3-6 to 2-6 times greater in the 4.0-3.0 Ga interval (Pollack, 1997; de Wit, 1998). As a result, the potential temperatures of komatiitic rocks are thought to have been c. 170 -+- 20~ higher in the Mid-Archaean (e.g., Galer and Metzger, 1998). In contrast, de Wit (1998) discusses experThe Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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imental and field observations that certain high-Mg melts were hydrous and that temperature regimes similar to those beneath modern mid-ocean ridges would have been pertinent. This argument rests essentially on a relatively hydrous mantle, but in the same paper, it is noted that early Archaean subduction zones were probably shallow as a result of low angles, and that consequently recycling of large volumes of water through the mantle wedge was not to be expected. The concept of low angle subduction in the Archaean (see also section 3.5) is related to the thicker oceanic crust inferred from higher heat flow (de Wit, 1998). Similarity of komatiitic melt temperatures to those beneath mid-ocean ridges is only circumstantial evidence in favour of Archaean plate tectonics, and such arguments are not governed by the principle of uniformitarianism. Low angle subduction is also used to explain the bimodal dacitic-basaltic volcanic rocks typical of the Archaean, in contrast to the essentially intermediate compositions found above high angle Phanerozoic subduction zones (e.g., Harris et al., 1993). The interpretation within a plate tectonic paradigm is obvious in this example. The application of Lyellian uniformitarianism thus appears to be, at least to a certain degree, opportunistic. The Lyell thesis defends the idea that similar geological conditions arising from invariate chemical, physical and biological laws will produce the same geological products throughout time; it does not in any way support the notion that because certain rock types develop in a specific tectonic setting today, they should also be expected, largely, from the same setting in the Archaean. Donaldson et al. (2002) review actualism, an amplification of uniformitarianism, and they and others stress that it is the rates and intensities at which processes occur which changes over geological time, rather than the processes (e.g., Eriksson et al., 2001a). Adoption of the plate tectonic paradigm for the Precambrian, combined with strong evidence for higher mantle heat flux in the Archaean, has led to arguments for much longer constructive plate margins (mid-ocean ridges), up to 27 times present values, resulting in the concomitant idea of small, possibly fast moving plates (Hargraves, 1986). However, there is no direct evidence for systematically higher Archaean plate velocities (Kr6ner, 1991). The classical scientific method makes no provision for the application of paradigms. Pollack (1997) notes that the beginning of the Archaean was highly exotic, particularly in its thermal behaviour, relative to the modern Earth, that a tectonically and thermally familiar world was active from the end of the Archaean, and that after the lapse of some future time period of geological scale, Earth will cool enough to replace the present layered mantle convective system with one where heat is lost through a conductive thermal regime. When plate tectonics is viewed over this enormous time scale, a case for simplistic uniformitarian application of this paradigm becomes essentially irrelevant.

Mantle Superplumes Abbott and Isley (2002b) stress the importance of mantle superplumes (and superplume eras, characterised by an anomalous abundance thereof) as influencing tectonic change on

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Earth, by building continents (through formation of thick and extensive oceanic plateaus that are not amenable to subduction) and by fragmenting them (see also Condie, 1998, 2001 a; sections 3.2 and 3.3). Condie (2001a) defines a superplume as having a plume head diameter of 1500-3000 km and an eruptive volume ~> 5 x 106 km 3, and a superplume event as comprising regionally clustered occurrences over <~ 100 My periods. However, inherent incomplete preservation of Precambrian plume products necessitates using proxies to estimate plume volumes over time and hence to propose superplume eras (Abbott and Isley, 2002b); komatiites, flood basalts, massive mafic dyke swarms and layered igneous intrusions are used (Isley and Abbott, 1999; Ernst and Buchan, 2001 a, 2002a). Additional proxies are outlined by Condie in section 3.2. The most complete record of inferred large global magmatic events is provided by Ernst and Buchan (2001b) (section 3.3). Komatiites are the most robust of the proxies proposed by Abbott and Isley (2002b). Estimation of original flood basalt surface area is possible, based on preserved surface area, age, a decay constant and an erosion correction (Condie et al., 2000; Abbott and Isley, 2002b). An almost linear relationship between the square of dyke width and total magma volume passed through the feeder (Fialko and Rubin, 1999) is best applied to the widest dyke in a swarm, to enable estimation of areal extent of flood basalt provinces; superplumes have dyke widths > 70 m, but thermal erosion in dykes wider than 300-400 m becomes problematic in calculations (Abbott and Isley, 2002b). The fourth proxy for superplumes is layered intrusions (with no evidence for arc affinity) containing Cr and/or PGE enrichment. Applying these methods, Abbott and Isley (2002b) demonstrate that the extent of Precambrian flood basalts and the intensity of superplume eras dwarfed equivalents in the Phanerozoic. As an example of this, the largest Precambrian superplume events formed flood basalts 10-20 times the surface area of the biggest Phanerozoic event, and these Precambrian events would have covered 14-18% of Earth's surface. Superplume activity shows a progressive decline from 2.9 Ga, the three largest superplume eras are in the Archaean and the strongest superplumes were those at c. 2.76 Ga and 1.65 Ga (Abbott and Isley, 2002b); Condie (section 3.2) proposes two major superplume events at c. 2.7 Ga and 1.9 Ga. Applying similar methods and proxies, Ernst and Buchan (2002a) define similar superplume eras, and, like Condie (2001a) and Abbott and Isley (2002b), find the c. 2.8-2.7 Ga event to have been the greatest (section 3.3). Prior to 2.8-2.9 Ga, preservation problems and concomitant lack of data for application of the proxies makes estimation of superplume eras impossible (Ernst and Buchan, 2002a; Abbott and Isley, 2002b). Arguments about hydrous mantle-derived komatiites (Grove et al., 1994; Parman et al., 1997) aside, evidence discussed briefly here strongly supports significantly higher mantle heat flux in the Archaean, which decreased over Precambrian time (Pollack, 1997), and analogously higher production of Archaean komatiites (e.g., de Wit, 1998) again subject to decreasing importance over time. An enhanced role for magmatic/thermal processes relative to plate tectonic processes with increasing age in the Precambrian thus appears to be pertinent (e.g., Pollack, 1997). Sylvester et al. (1997) point out that modern volcanic successions commonly are ascribed to plate tectonics and mantle plumes, and that this combination is therefore pertinent also for Archaean successions.

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Assumptions, Facts and Interpretations Considered Critical for Formulating a Model for Early Precambrian Crustal Evolution by a Combination of Magmatic and Early Plate Tectonic Processes In this contribution, we wish to avoid the pitfalls of adopting either plate tectonic or "plume tectonic" paradigms. In addition, we support some basic premises: (1) at some stage in its earliest post-planetary accretion history and following differentiation into core and mantle, Earth was subject to whole mantle convection, and an entirely thermally driven regime preceded the onset of plate tectonics (Pollack, 1997); (2) a transition from this to a combined magmatic/thermal and plate tectonic system must have occurred (Trendall, 2002); (3) the relative relationship between plate tectonic and thermal/magmatic (plume) systems has changed over geological time (Pollack, 1997; Abbott and Isley, 2002b); and (4) that a combination of plate tectonic and magmatic processes (plumes) has controlled Precambrian crustal evolution (Condie, 2001 a; Eriksson et al., 200 lb). In attempting to discuss a possible model for early Precambrian crustal evolution, a number of basic facts reported from the rock record, and unresolved paradoxes must be addressed, and these are detailed in the following four paragraphs. It must first be stressed that Early Archaean (4.0-3.0 Ga) rocks account for only 0.5% of present continental surface area, and those from the Late Archaean (3.0-2.5 Ga) for c. 7% (de Wit, 1998). The oldest rocks on Earth, the Acasta gneisses (Slave Province, Canada) which are just over 4.0 Ga, appear to have been continental and were likely derived from hydrated, chemically depleted mafic protoliths (Bowring and Housh, 1995). They may be considered part of the so-called "grey gneisses" (Martin, 1994), small remnants in many cratons of Early Archaean crustal evolution; some are found associated with highly altered volcanic rocks, others with equally deformed and metamorphosed shelf-like quartzites, pelites and carbonate rocks (de Wit, 1998). Continental crustal growth appears to have been rapid early in the Archaean (de Wit, 1998), and deep craton roots or keels are also typical of those from the Archaean in contrast to younger cratons (e.g., Polet and Anderson, 1995). Granite-greenstone assemblages are the dominant crustal rocks on Archaean cratons, and the granitoids form two major groups, the tonalite-trondhjemite-granodiorite (TTG) and granodiorite-granite-monzogranite (GGM) assemblages (de Wit, 1998). Modelling indicates that the thermal lithosphere thickness at c. 4.0 Ga was about 0.3 of modern values, and that it rose to about 0.6 by the end of the Archaean; in contrast, thicker lithospheric keels are inferred (Pollack, 1997). Much work has been done on greenstone belts (sections 2.3, 2.4, 4.3 and 4.4), most of which are highly tectonised, with strong field and geophysical evidence for structural imbrication (Kusky and Vearncombe, 1997). Horizontal shortening and extension both appear to have occurred, and metamorphic grades vary from below greenschist up to granulite facies (de Wit and Ashwal, 1997a). These authors find that komatiite volumes have often been over-estimated in these belts, and that there is no unequivocal decrease of these rocks through time, although a definite peak in preservation occurs in the Neoarchaean. In the two oldest cratons, Kaapvaal and Pilbara, early extensional faults are noted, and upwelling mantle at zones of rifted, rigid continental crust is strongly supported (de Wit, 1998).

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Strains estimated from greenstones are compatible with those associated with modern-type plate tectonic regimes (Hudleston and Schwerdtner, 1997). The largest Archaean craton, Superior is also the best studied, especially in terms of geochronology; most of these rocks formed in < 100 My, and strong field evidence exists for accretionary tectonics, terrane juxtaposition over an even shorter period of c. 30 My, and thrust-stacking of crustal and supracrustal architectural elements of varying ages (Stott, 1997; de Wit, 1998). Similar data are recorded from many of the well studied greenstone belts (Kusky and Vearncombe, 1997; Stettler et al., 1997), such as that at Barberton (e.g., Brandl and de Wit, 1997). However, Barberton had a polyphase tectonic history of close to 500 My (de Wit and Ashwal, 1997a). It should be borne in mind that interlayered mafic and felsic volcanic successions, interpreted by many to result from shortening, may also have a purely magmatic origin, independent of tectonism (Sylvester et al., 1997). The same authors stress that while evidence for Archaean plate tectonics is based largely on identifying large scale horizontal deformation and inferred ophiolite remnants (e.g., de Wit et al., 1987b), no unequivocal ophiolites of this age have been documented (section 3.7). Sedimentary rocks in greenstone belts are mostly epiclastic lithologies associated with volcanic successions, and siliciclastic rocks; this points to exposed continental crust and weathering thereof, yet the absence of extrabasinal detritus combined with common shallow water facies in greenstone belts leads to inherent difficulties in models of greenstone volcanics having extruded onto continental crust (Eriksson et al., 1997). Arndt et al. (1997) identify two major suites of volcanic rocks in greenstone belts: more voluminous lower basaltic-komatiitic rocks (which they relate to plumes and compare with modem oceanic plateaus), and upper mafic-felsic lithologies of calc-alkaline affinity (which they find comparable to modem arcs). In a more recent publication, Arndt (1999) emphasises that eruption of crust-contaminated flood basalts onto submerged continental platforms is a tectonic setting common in Archaean-Proterozoic greenstone belts, but one which is rare within the Phanerozoic. This has implications for continental crustal and oceanic volumes (Arndt, 1999); the former may have been voluminous already early in Archaean history (cf. model of early growth, Armstrong, 1981; see also section 2.8), and the latter may either have been greater (due to degassing of hotter Archaean mantle; cf. Fyfe, 1978) or smaller (due to a more hydrated Archaean mantle; cf. de Wit, 1998). Thicker oceanic crust, inferred by other workers too (e.g., Sleep and Windley, 1982; Pollack, 1997; de Wit, 1998) may also have reduced ocean volumes resulting in flooding of cratons (Arndt, 1999). In an alternative model, episodic continental crustal growth (e.g., Eriksson, 1995) with eruption of oceanic plateaus (cf. Abouchami et al., 1990) could have led to submergence of continental platforms largely during volcanism itself (Arndt, 1999). Arguments have long continued on the role of granitoid plutonism (and concomitant vertical tectonics) in the evolution of apparently correlatable autochthonous greenstone stratigraphies developed on some well studied cratons, such as Pilbara, Dharwar and Zimbabwe (e.g., Hickman, 1983; Collins, 1989; Treloar and Blenkinsop, 1995; Chardon et al., 1996) (sections 2.5-2.7). For the Zimbabwe craton, episodic crustal growth has been related to plumes causing rifting of older continental crust, with concomitant partial melting of mafic material and crust to form bimodal supracrustal volcanics within a greenstone stratigra-

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phy widely developed across this craton (e.g., Bickle et al., 1994; J.E Wilson et al., 1995; Hunter et al., 1998). More recently, detailed structural and geochronological data have enabled the alternative plate tectonic model, where lateral accretion and thrust stacking are important, to be strongly supported (Kusky and Kidd, 1992; Kusky, 1998; Dirks and Jelsma, 1998, 2002; Hofmann et al., 2001). For the Pilbara craton, contrasting viewpoints remain more entrenched (e.g., Bickle et al., 1980; Boulter et al., 1987; Trendall, 1995; Zegers et al., 1996, 1998a; Pike 2001; Pike and Cas, 2002) (sections 2.5-2.7). Basically, the role of granite diapirism versus tectonic shortening (and extension) in Archaean crustal evolution remains unresolved (de Wit and Ashwal, 1997a).

Trendall's "Plughole" Model for Hadaean and Early Archaean Crustal Evolution Trendall (2002) presents a speculative hypothesis on the origin and early growth of continental crust, based on forwards modelling from concepts of early Earth accretion (e.g., Stevenson, 1987; Allbgre et al., 1995; McCulloch and Bennett, 1998) (see also section 1.2) and inverse modelling from the known rock record, which goes back as far as c. 4.0 Ga (e.g., Bowring and Williams, 1999). This model accounts for many of the basic facts and many of the unresolved paradoxes inherent in any discussion of early Archaean crustal evolution, and we here speculate further on the basis provided by this hypothesis. Trendall's (2002) modelling begins arbitrarily at 4.3 Ga, with a molten Earth having a differentiated core, and whole mantle convection, following reasoning outlined by Davies (1992b) and Pollack (1997). Convection is postulated to have occurred in large polygonal cells, with upwelling at polygon margins and with centres of convective descent (CCDs) where downwards flow took place (Fig. 3.6-1). A strength of this postulate is that it provides for later evolution of the polygon sides into primitive mid-ocean ridges, thus making possible a transition from an early Earth dominated by thermal-magmatic processes to one where rigid plates and subduction became paramount. Within these polygonal systems, Trendall (2002) infers rapid cooling and formation of a first crust of devolatilised mantle composition, thinnest at the peripheral "ridges" and increasing in thickness towards the CCDs. Much higher convective velocity due to significantly enhanced radioactive heat generation at this stage in Earth's evolution (McKenzie and Weiss, 1975) would have carried this crust downwards at the CCDs, where it became recycled into the mantle. As cooling proceeded and convective velocity decreased, partial melting of this descending ultramafic, transient crust would have generated basaltic magmas, with potential diapiric buoyancy; however while descent rates exceeded diapiric rise potential, these differentiates were also recycled into the mantle (Trendall, 2002). With time and increase in diapiric rise forces, large plugs of differentiates could have accumulated above the CCDs (Fig. 3.6-1), where the downwards pull of convecting mantle would have been less. The Trendall (2002) model differs from previous viewpoints in generating early sialic material over centres of descent rather than above rising convective cells, and thus follows the ideas of Davies (1992b, 1993). The postulated differentiate plugs would also have been subjected to tectonism and remelting, and the > 4.0 Ga zircons from the Yilgarn craton, Australia (Nelson et al., 2000) are related tentatively to these recycling processes (Trendall, 2002). During this period, when differ-

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entiate plugs began to form, it is surmised that surface water started to condense, forming small "oceans" centred on the CCDs and separated by the exposed upwelling ridges of the polygons (Fig. 3.6-1); this suggestion is compatible with the proposal of de Wit and Hynes (1995) that Hadaean mid-oceanic ridges (or something analogous) were emergent above sea level, thus hindering the onset of plate tectonics s e n s u stricto. By applying Stoke's Law (to control the rate of diapiric rise of the differentiate plugs) to continuing cooling and exponential slowing of mantle convection systems, the (Trendall, 2002) model predicts a cross-over point at which the lightest, granitoid differentiates beneath the CCDs were in balance with convective descent potential (Fig. 3.6-1). Trendall hypothesises that this may have occurred at c. 4.0 Ga, and that when this cross-over took place, sialic plugs ascended isostatically to form protocontinents, preserved as the "ancient gneiss complexes" of the Archaean cratons known today. These protocontinents were very small, relatively hot and ductile and thus subject to flow rather than fracture when placed under stress; they likely had high freeboard, and their formation raised water levels of the incipient oceans lying above the CCDs (Fig. 3.6-1). Aggressive weathering (e.g., Corcoran et al., 1998) (sections 5.10 and 5.11) would rapidly have peneplained the protocontinents, and quartzose siliciclastic deposits could have formed locally, possibly with > 4.0 Ga zircons (Trendall, 2002). Continuance of mantle convection beneath the protocontinents (cf. Davies, 1993) established a process which may be viewed as an early version of subduction. Partial melting of mantle material (i.e., protooceanic crust) and differentiates below the protocontinents also continued, and diapiric and isostatic ascent of magmas ranging in composition from komatiitic through to silicic would have become increasingly common, resembling processes from the modern Earth (Clemens, 1998). In this way, the protocontinents would have grown gradually, both in extent and thickness, and by the addition of siliceous intrusions, underplating and overplating (flood basalts) of mafic-ultramafic magmas, the protocratons extended outwards from the early ("ancient gneiss complex") cores. This could explain why, in many Archaean cratons, the addition of both TTG and greenstone material, apparently accreted successively and in various geometries, around an early core shield or protocraton (e.g., de Wit and Ashwal, 1997b). Discussion

As the protocratons greyS, they would have been subject to local extensional stresses due to the impingement of ascending magmas from below, leading to localised rifting of the developing continental crust and to the deposition of komatiitic, basaltic and felsic lavas and volcaniclastics. This would be compatible with Arndt's (1999) model of eruption of crustcontaminated flood basalts onto submerged continental platforms discussed earlier. As the protocratons grew and slowly developed into true cratons, they would have displaced water in the ponded micro-oceans within the polygonal convection systems postulated by Trendall (2002), thereby drowning the thinner and thus isostatically lower marginal cratonic areas. These envisioned conditions closely resemble those inferred for many early Archaean greenstone belts (sections 2.3, 2.4, 4.3 and 4.4), with the development of predominant volcanic rocks of diverse composition, many of them subaqueously; volcaniclas-

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tic deposits, deeper and shallower water sediments and even small proportions of emergent sediments would have been possible within such essentially linear rift systems within the growing protocratonic margins. The requirement that greenstones were subjected to significant horizontal shortening (in addition to extensional stresses) and also to thrusting, discussed previously (e.g., de Wit, 1998), could easily have resulted from the continuing convective descent of mantle material at the CCDs, now lying beneath developing cratons; as already argued, these convective systems would have begun to resemble subduction systems. Obviously, large scale intrusion of silicic differentiates (sections 2.5-2.7) would have been necessary to thicken and strengthen the growing protocratons and to provide the necessary brittle strength of continental crust. A major advantage of the Trendall (2002) model outlined briefly here is that it provides for a transition from an entirely thermal magmatic system (whole mantle convection) to one that begins to resemble "normal" plate tectonics, with protocratons and "subduction" of primitive mantle crust beneath them. The transition from this hypothetical model (Fig. 3.6-1) to a recognisable plate tectonic regime would presumably have depended upon a number of changes in the model parameters: (1) change of the "ridges" at the outer boundaries of the convectional polygons into the prototypes of mid-oceanic ridges; (2) implicit in this would have been the development of basaltic ocean crust in place of the mantle compositions in the Trendall (2002) model, probably due to adiabatic partial melting of upwelling mantle in these zones; (3) also implicit in this transition would have been extension of the separated, ponded proto-oceans within each polygon, until they amalgamated into a global ocean, most likely due to a combination of continued devolatilisation of the mantle, and to growth of protocratons which raised water levels until the "ridges" became submerged (cf. de Wit and Hynes, 1995); (4) a change from whole mantle convection to a more familiar two-layered mantle convection regime (see also sections 3.2 and 3.4). It is possible that lateral growth of the emerging cratons above the CCDs displaced areas of downwelling mantle further and further away from the growing continental keels and thus closer to the upwelling "ridges"; eventually an unstable convective system may have resulted, leading to collapse of the whole mantle circulation and its replacement with a two-layered mantle convection regime. It is uncertain when the requisite changes

Opposite: Fig. 3.6-1. Trendall's (2002) "plughole" model for Hadaean-Archaean crustal evolution. (a) Polygonal systems of marginal mantle upwelling and centrally located centres of convective descent (CCDs). (b) A transient ultramafic crust at the CCDs is overlain by the first ponded and small oceans, separated by marginal polygonal "ridges". Differentiation of downwelling mantle (i.e. transient crust) beneath CCDs produces a suite of different magma compositions, which are not able to ascend. (c) Schematic representation of the relationship between rate of magma descent at CCDs and the diapiric rise potential of silicic differentiates, with possible equivalence being achieved at c. 4.0 Ga. (d) Subsequent evolution of proto-continents above CCDs, bounded by marginal seas which amalgamated into the global ocean. As proto-continents expand laterally due to a combination of under- and overplating and granitic to ultramafic intrusions, rifting results from continuing mantle descent (the start of "subduction"?) beneath proto-cratons. Greenstone belts develop within these rifts.

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(1-4 above) would have occurred, to enable the transition from the "plug-hole" and protocontinent regime of the Trendall (2002) model to a recognisable plate tectonic system, except to note that this likely occurred within the 4.0-2.5 Ga period and that these various changes may not all have occurred simultaneously, either globally, or for any individual poygonal cell. Another necessary factor for the onset of plate tectonics s e n s u s t r i c t o would probably have been the necessity to approach approximate balance in rates of marginal upwelling at polygon outer "ridges" and in rates of convective descent at the CCDs. Obviously, these two rates would only have approached each other as the protocratons grew outwards enough to make the circular zones of "subduction" around them much larger than the initial, restricted, centrally-lying CCDs. It is also postulated here that collapse of the whole mantle convective system may have led to a period when heat loss was achieved largely by an abundance of mantle superplumes (sections 3.2 and 3.3), before plate tectonics in recognisable form became predominant with heat loss mainly from mid-ocean ridges. Nelson (1998a) recognises a global catastrophic convective overturn of Earth's mantle at c. 2.7 Ga (section 3.4), along with evidence for the operation of plate tectonic processes at the same time. As outlined above already, the greatest global superplume era is thought to have occurred at c. 2.8-2.7 Ga (Abbott and Isley, 2002b; Ernst and Buchan, 2002a) (sections 3.2 and 3.3). There is thus evidence that the 2.8-2.7 Ga period may have been one characterised by both large scale mantle instability and the onset of plate tectonics in recognisable form. There is very strong evidence for plate tectonics having been paramount in Late Archaean continental crustal growth, particularly for the well preserved Superior and Slave Provinces of Canada (e.g., Stott, 1997; de Wit, 1998; Mueller and Corcoran, 1998, 2001) (sections 2.4 and 3.9). For the largest and best preserved Archaean basin in the > 2.8 Ga period, the Witwatersrand (c. 3.1-2.8 Ga) of Kaapvaal, Catuneanu (2001) discusses evidence for a lack of dynamic loading in this flexural foreland basin system, which he ascribes to lower rates of subduction. Older basins are well studied from the Pilbara craton, Australia; the ~> 3009 + 4 to c. 2945 Ma Whim Creek basin offers no direct evidence for plate subduction or an arc setting (Pike and Cas, 2002), and the ~< 2990 to > 2955 Ma Mallina basin similarly offers no evidence for an active plate margin or subduction, and it is interpreted as an intracratonic rift depository (Smithies et al., 2001). The enhanced continental crustal growth rates (section 2.8) close to the Archaean-Proterozoic boundary (e.g., Eriksson, 1995) may be ascribed to the onset of a fully plate tectonic regime at about 2.7-2.5 Ga, with accretion of island arcs, oceanic plateaus and small cratons (e.g., Limpopo belt, southern Africa; section 3.8). Additionally, there is reasonable evidence, especially geochronological (Aspler and Chiarenzelli, 1998) for the first supercontinent, Kenorland in the Neoarchaean (sections 3.2 and 3.9). It is thus possible that the c. 3.0-2.7 Ga period saw a transition, of diachronous character on the global scale, from a system possibly analogous to the Trendall (2002) model to a fully plate tectonic regime, and that the Witwatersrand basin was deposited during this transition, when Phanerozoic type subduction processes had begun, but were not yet efficient. The extended attenuation and final breakup postulated for Kenorland, occurred from c. 2.5-2.1 Ga, indicating that although plate tectonics was fully operational, it was still relatively slow by Phanero-

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zoic standards. There is a paradox in these suggestions, that despite higher Neoarchaean mantle heat flow, emerging plate tectonics may have been characterised by slower rather than higher plate movements. It is possible that heat was trapped beneath growing continents, and that the latter were "suspended" above upwelling mantle, or even superplumes (see also section 3.2), thereby accounting for slower subduction rates, and for the apparent paradox. It is also possible that the "plughole"mplate tectonic transition occurred prior to Witwatersrand deposition. The concept discussed here, of a transition from some postulated Hadaean-Palaeoarchaean crustal evolutionary regime to recognisable plate tectonics, is seen by many to be unnecessary, if recourse is had to shallow, low angle subduction zones (cf. Martin, 1994) (section 3.5). This idea provides for a geological environment in which relatively modem magmas of TTG affinity can be produced (e.g., Karsten et al., 1996). Phanerozoic adakites are suggested as analogues for Archaean TTGs, and low angle subduction (section 3.5) of significantly thicker oceanic crust under Archaean geothermal gradients would have yielded TTGs with Mg numbers analogous to those experimentally derived from partial melting of basalts (D.H. Abbott, 2002, pers. comm.). Modelling of Archaean heat loss supports the possibility of low angle subduction (e.g., Pollack, 1997) as does the typical bimodal volcanism of Archaean greenstones (Abbott and Hoffmann, 1984). Once Archaean mid-ocean ridges had become subaqueous, inferred to have occurred at c. 4.0 Ga (de Wit and Hynes, 1995) (note the Trendall (2002) model suggests that diapiric rise of protocratons and expansion of ponded Hadaean oceans occurred at this same approximate age), hydration and serpentinisation at the ridges would have formed more buoyant upper oceanic lithosphere (de Wit, 1998). Delamination of this lighter portion from underlying mantle material is an implicit part of the intra-oceanic obduction complex, or oceanic crust duplex model for Early Archaean cratonic evolution (de Wit and Hart, 1993; de Wit, 1998), which must be seen as an alternative to the Trendall (2002) model discussed here. This delamination could also explain the lack of high-Fe lower crust and the rare underplating inferred in Archaean terranes (D.H. Abbott, 2002, pers. comm.). Obviously, due to a total lack of a preserved Hadaean rock record, and the paucity of rocks older than 3.0 Ga, the formation of early Precambrian continental crust will always remain, at least partially, enigmatic. The Trendall (2002) model discussed here at some length, however, does resolve many paradoxes inherent in studying these ancient terranes. Apart from providing for greenstone lithologies to develop above sialic crust (cf. Arndt, 1999), the Trendall model is also compatible with the concept of a small number of ancient cratonic nuclei (e.g., Horstwood et al., 1999; B6hm et al., 2000; Dirks et al., 2002; Trendall, 2002), explains the abrupt beginning of the known geological record at c. 4.0 Ga, and can account for rapid early crustal growth (cf. Armstrong's (1981) model) (section 2.8). The Trendall (2002) model provides for quasi-simultaneous formation of early cratonic nuclei of similar character at diverse locations on Earth's surface, and for their lateral growth by underplating, overplating, granitoid (and other compositions) intrusions as well as greenstone belt formation; the decreasing age of this new continental crust outwards from the early cores is analogous with the accretion of arcs, terranes and crustal fragments by recognisable plate tectonic processes (e.g., de Wit et al., 1992; Windley, 1993, 1995). Metamor-

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phic grades from Archaean terranes, especially greenstones, do not indicate significantly hotter regimes than young orogens (Pollack, 1997), and Archaean cratons are characterised by anomalously low heat flow (Nyblade and Pollack, 1993). These observations are difficult to reconcile with crustal evolution within the plate tectonic paradigm, but are easily explained by Trendall's (2002) model where deep and relatively cool roots are inferred to have formed early beneath protocratons. The absence of such roots beneath younger continental crust provides additional support for this aspect of the Trendall model (Polet and Anderson, 1995). Trendall (2002) also points out that his "plughole" model has no significant relationship to meteorite bombardment at the Hadaean-Archaean boundary and in the earliest Archaean, as such bolide impacts would have had little effect on a whole mantle convecting magma ocean or its quenched komatiitic crust. Later, smaller such events (see sections 1.3 and 1.4 for detailed discussions) may have been relatively common, and are preserved in the cratonic record from the Archaean onwards. Conclusions

Our tentative conclusions are thus that, instead of there being two mutually exclusive models for early Precambrian crustal evolution, namely that plate tectonics either applied or did not (e.g., de Wit, 1998; Hamilton, 1998, respectively), a gradual transition may be more pertinent (see also sections 1.2 and 3.4). This transition may have been from a fully magmatic Hadaean Earth, with whole mantle convection, to one where heat loss was predominantly from mid-ocean ridges in a fully plate tectonic regime, with episodic influences from mantle plumes, superplumes and superplume eras (sections 3.2 and 3.3). We see the Trendall (2002) "plughole" model as a reasonable hypothesis for formation of Early Archaean continental crustal nuclei, by differentiation of transient komatiitic crust and eventual diapiric rise at his "convective centres of descent" (CCDs), with a relatively sudden, global onset of the geological record close to 4.0 Ga as a result. Continued growth around these protocratons could have proceeded by further ascending magmas of diverse composition, including eruption of bimodal greenstone magmas into early crustal rift depositories. Continued descent of komatiitic and later basaltic crust from marginal "ridges" defining polygonal CCD systems would have provided a primitive form of ridge and subduction zone tectonics (Trendall, 2002). Growth of ponded oceans above CCDs would finally have overstepped the polygonal "ridges", forming the global ocean. Outwards growth of protocratons and concomitant trapping of heat beneath them, as well as division of downwelling convective cells adjacent to growing cratonic keels may have led to eventual collapse of whole mantle convective systems as proposed in Trendall's model, and to the onset of a layered mantle convection system. It is possible that the major global superplume era, at c. 2.8-2.7 Ga represents a time of transition in these mantle convective regimes. Early intracratonic basins from Pilbara (Smithies et al., 2002; Pike and Cas, 2002) suggest rifting as a major control, and the c. 3.0-2.8 Ga Witwatersrand basin indicates the possibility of slow subduction within a foreland basin model (Catuneanu, 2001). Evidence in favour of enhanced continental crustal growth rates close to the Archaean-Proterozoic boundary (section 2.8) may reflect the onset of a fully plate tectonic system globally, with accretion of

3.7. P r e c a m b r i a n Ophiolites

2 13

arcs and smaller crustal fragments about the earlier-formed "plughole" cratonic nuclei. In this, c. 2.7-2.0 Ga period, plate tectonic spreading and subduction rates may still have been lower than their Phanerozoic-Modern equivalents, as evidenced by the apparently slow attenuation and breakup of the Kenorland supercontinent from c. 2.45-2.1 Ga (Aspler and Chiarenzelli, 1998) and the equally slow amalgamation of the Eburnean supercontinent from c. 2.15-1.8 Ga (Eriksson et al., 1999). Once plate tectonics, punctuated by mantle superplumes and superplume events (sections 3.2 and 3.3), became well established on Earth, their combination would have had a major control on the global carbon cycle, with isotopic excursions directly related to the supercontinent cycle and to major periods of mantle overturn and superplume genesis (Lindsay and Brasier, 2002) (section 5.3). Planets lacking the crust-mantle interactions implicit in plate tectonics probably never developed life beyond primitive single-cell organisms and the rich organic evolutionary history of Earth rests, ultimately, on this plate tectonic basis (Lindsay and Brasier, 2002), active since at least the Neoarchaean, if not earlier in a number of possible guises.

3.7.

PRECAMBRIAN OPHIOLITES

J.R. CHIARENZELLI AND E.M. MOORES

Introduction Ophiolites, as defined in the next paragraph, have been the subject of considerable interest and intensive research since their recognition in the Alps nearly a century ago (Steimann, 1906, 1927) and especially since the recognition, in the late 1960s, of their formation at oceanic spreading centres at mid-ocean, back-arc, and forearc settings. Because ophiolites represent the only surviving vestiges of oceanic crust formed before 180 Ma, they are of crucial importance in constraining tectonic, thermal, and magmatic processes operative during the Precambrian. Well-preserved Phanerozoic ophiolites are emplaced (obducted) by aborted subduction of a continental margin or island arc. Some 35 probable and possible ophiolites predating the assembly of Rodinia (pre-1.0 Ga) (sections 3.10 and 3.11) suggest ocean floor spreading as early as 3.5 Ga. Temporal trends in ophiolites may reflect the planetary evolution of Earth (see also sections 1.2 and 3.6). Archaean ophiolites may represent the equivalent of thickened oceanic plateaus where higher heat flows and greater plume activity (sections 3.2 and 3.3) enhanced partial melting and oceanic spreading. Oceanic thinning during the Proterozoic may have led to conventional plate tectonics and temporal atmospheric changes, eventually leading to the Phanerozoic.

Recognition of Precambrian Ophiolites The classic pseudostratigraphic ophiolite sequence includes, in ascending order, tectonised suboceanic (peridotitic) lithosphere, mafic-ultramafic plutonic rocks, sheeted-dyke (with dyke-within-dyke relations) intrusive complexes, extrusive basaltic volcanics (pillowed or 777ePrecambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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massive), and deep-sea sedimentary or volcaniclastic rocks (Anonymous, 1972). Ophiolite emplacement (obduction) demonstrably is a product of collision of a subduction zone with a continental margin or island arc. Fragmentary preservation of partial sequences also occurs in subduction complexes. Ophiolites are relatively abundant in Phanerozoic orogenic belts, but even so, they comprise only about 0.0012% of ocean crust developed during that time (Coleman, 1977, p. 16). Collisional and/or accretionary tectonic processes are such that only rarely are complete or "ideal" ophiolites found. Inherent spatial variability in the spreading centre environment(s) where oceanic crust originates provides an additional complication for the comparison of ophiolitic sections. Distinct ophiolite sequences may reflect differences in spreading rates, observed lithologies, thicknesses of individual units, and structures, among other factors (Hess, 1962; Dilek et al., 1991; Moores, 2002). In modern oceans, slow-spreading crust contains evidence of intermittent magmatism, abundant normal faulting, and abundant serpentinite, whereas fast-spreading crust contains thicker magmatic rocks, less evidence of faulting, and a more complex ophiolitic sequence. These differences are present within individual ophiolite complexes as well. A number of secondary features can influence the ultimate preservation of ophiolites. As implied above, the emplacement of ophiolites along plate boundaries results in a relatively high tectonic position of ophiolite nappes early in the convergence/collision orogenic process. Even if preserved relatively intact, extensive changes resulting from tectonic dismemberment, structural interleaving, and/or metamorphism during subsequent orogenesis are common. Because ophiolites are emplaced by subduction-related thrust faults rooted in the mantle, their basal contacts represent fossil subduction zones. Emplacement in a tectonically high position with subsequent uplift and erosion in orogenic belts may lead to extensive destruction of the thickest and best-preserved portions of the complex, leaving only the attenuated root. Major differences in ophiolite stratigraphy corresponding with geologic age also occur (Moores, 2002). Because of their great age, Precambrian ophiolites are more likely to be altered, laterally or vertically discontinuous, and/or less well preserved than Phanerozoic ophiolites. In some deeply eroded Precambrian orogenic belts, former sutures may contain little more than highly metamorphosed and strained, lensoidal tectonic inclusions of a once more extensive ophiolite suite (e.g., Limpopo orogenic belt, section 3.8), interleaved with rocks of continental, arc-related, or accretionary affinities. Nonetheless, if verified as remnants of oceanic crust, their presence implies the prior operation of oceanic spreading processes and subsequent emplacement (obduction) of the oceanic crust during continental convergence and collision. In much the same way, the erosion of extensive flood basalt provinces on Earth and other planets has been inferred from the numerous well-preserved radiating dyke swarms that presumably fed them (Ernst and Buchan, 1997) (section 3.3). Vast areas of the world's shields are highly metamorphosed and deeply eroded, rendering the recognition of Precambrian ophiolites difficult. Classic features supportive of an ophiolitic origin are of limited value in high-grade or tectonically dismembered terranes (Table 3.7-1). In addition, numerous other possibilities (Table 3.7-1) must be ruled out before an ophiolitic affinity can be verified (Moores, 2002). Because lithologies associated with ophiolites can occur in a variety of other geologic environments, the geologic

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Table 3.7-1. Evidence for ophiolite sequences and possible alternative geological settings (after Moores, 2002) Possible alternative geologiPossible evidence for ophiolite sequences cal settings Ophiolite Part or all of intact complex Fossil island arc Disrupted fragments of a complex Sheeted dykes, particularly when with other parts of the se- Oceanic plateau quence Serpentinite overlain depositionally by marine sediments Stratiform ultramafic to mafic and or extrusive volcanics, or intruded by hypabyssal rocks intrusion Fault or shear zone with lensoid fragments of ophiolite Early or aborted rift-related magmatism affinity Metamorphosed mafic or ultramafic rocks perhaps derived M61ange zones (convergent or transform) from an ophiolite complex Geophysical anomaly in continental crust (high density Subcontinental mantle slice and/or magnetism) Overlying sedimentary rocks indicating deep-water envi- Impact related magmatism ronments context is an essential ingredient for sound interpretation. Unfortunately, the interpretation of ophiolite chemistry is complex and generally cannot be used by itself as a tectonic indicator. Conversely, the chemistry, and magmatism and extensional processes preserved during initial stages in the development of the Izu-Bonin-Mariana forearc in the Western Pacific closely resemble those of some supra-subduction zone ophiolites (Bloomer et al., 1995). The chemical affinities of individual ophiolites can be ascribed to several tectonic environments, and long-lived mantle heterogeneities (Moores et al., 2000; Cousens et al., 2001) play a significant role in the chemistry of mantle-derived magmatism, regardless of emplacement location or tectonic setting (see also section 3.5). Thus the geologic context, including associated lithologies and structures, is essential in evaluating the significance of a suspected ophiolite. Temporal and Spatial Occurrence of Ophiolites Moores (2002) notes that the age of identified and inferred ophiolite complexes cluster between 1.0-1.5 Ga, 1.8-2.3 Ga, 2.5-2.7 Ga and at c. 3.4 Ga. These times fit remarkably well with periods of global assembly of supercontinents and may suggest a causative link (sections 3.2, 3.8-3.11 and 5.3). One possible explanation is that the assembly of a supercontinent rimmed by subduction zones over a geoidal low (section 3.2) provides greater opportunities for obduction and ophiolite preservation in its interior than other phases in the supercontinent cycle (Hoffman, 1992). Alternatively, the convergence of previously separate continents that culminates in a supercontinent, involves closure of ocean basins containing intraoceanic island arcs and fracture zones, all formed at oceanic spreading centres. Thus continent-continent collision will be preceded by continent-subduction zone (either incipient or mature or island arc) collision, that is, ophiolite emplacement. In

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other words, ophiolite emplacement precedes continent-continent collision, as seen in the present oceans (e.g., Taiwan, New Guinea, Cyprus) and older intact Precambrian orogens (e.g., Trans-Hudson orogen). The location of ophiolites within ancient subduction zones indicates that they mark intracontinental sutures (remnants of disappeared oceans; e.g., Moores, 1981) or terrane margins. Reconstructions based on this assumption have been useful in interpreting the palaeogeography of Phanerozoic orogens. A possible Palaeoproterozoic example from the Trans-Hudson orogen includes several recognised ophiolites (Purtuniq, Flin-Flon, New Quebec, Fox River-Thompson) marking the intra-orogenic suture(s) between the Churchill and Superior Provinces along which the c. 2.1-1.9 Ga Maniwekan ocean opened and closed (Scoates, 1981; Baragar and Scoates, 1981; Cummings et al., 1982; Stauffer, 1984; Scott et al., 1991, 1992; Rohon et al., 1993; Skulski et al., 1993; Stern and Lucas, 1994; Lucas et al., 1996; Aspler et al., 2002b). However, it should be noted that the full width of the Trans-Hudson orogen is preserved in northern Canada; correlation of ophiolite obduction events of Precambrian age will rely ultimately on the validity of post-ophiolite continental reconstructions and geochronological verification.

Evolutionary Trends A number of observations suggest both similarities and differences between Precambrian and Phanerozoic ophiolites. Aside from the greater occurrence of komatiites in ophiolites > 1.8 Ga in age (cf. oceanic plateaus; Kerr et al., 1997; Moores, 2002), most lithologies present in Precambrian examples also occur in modern oceanic environments, implying formation by broadly similar processes. With decreasing age, a greater proportion of the known ophiolite complexes display supra-subduction zone chemical affinities, perhaps suggesting both temporal changes in geothermal gradients (fewer komatiites) and increasing mantle heterogeneity with time (Moores, 2002). Greater volumes of basalt with oceanic island affinity have been noted in pre-1.6 Ga ophiolitic sequences; this may be due to the greater role of hotspots, preservation potential of thickened oceanic crust, or greater magmatic activity due to higher heat flow favouring plateau development. Pre-1.0 Ga ophiolite complexes generally lack mantle tectonites, as do Phanerozoic oceanic plateaus preserved on land. Abrupt thinning of oceanic crust at 1.0 Ga may have occurred. Data from pre-1.0 Ga ophiolites suggests the thickness of oceanic crust has decreased with time, with major breaks at 2.5 and 1.0 Ga (Moores, 2002). Two physical constraints that may control ophiolite emplacement processes include the maximum thickness of ophiolite thrusts (c. 10 km; Moores, 1986, 1993) and the maximum thickness of subductable oceanic crust (c. 17 km; Cloos, 1993). If pre-1.0 Ga oceanic crust was > 10 km in thickness, there would be a tendency to truncate the ophiolite section, preferentially removing the lowermost tectonised mantle unit. In addition, if the thickness of Archaean oceanic crust was > 17 km, detachments accommodating displacement between the accreted upper portions and lower subducted portions would form (Moores, 2002). Wide accretionary prisms composed of the upper portion of Archaean oceanic crust would be stacked up along numerous thrust faults as "flakes", "shingles", or "panels" of limited lateral extent

3.8. Limpopo Belt o f Southern Africa

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(Hoffman and Ranalli, 1988). Voluminous pillowed or massive basalts would rapidly form wide accretionary prisms characterised by relatively low-grade metamorphism and with a limited range of magmatic and "emplacement" ages, stratigraphic and/or structural continuity, and sedimentary component. These features are found in many Archaean greenstone belts and may represent a distinctly different style of preserving oceanic crust, rather than a unique style of plate tectonics operative prior to the Proterozoic. Subsequent to the Archaean, reduction in the thickness of oceanic crust, and ophiolite preservation processes similar to modern ones may have resulted in enhanced recognition of ophiolites in the geologic record. If these temporal trends are real, they point to significant changes in plate tectonic processes and ocean crustal thickness with time (section 3.6). In addition, the relative importance of ophiolite emplacement in the generation of continental crust may have been much greater in the past than is commonly recognised. Implicit in each of the scenarios is a mechanism by which oceanic crust can avoid subduction and become accreted or obducted (emplaced) onto continental and/or island-arc margins. The proposed temporal changes in the thickness of the oceanic crust may have many important environmental implications (Moores, 1986, 1993, 2002). The explanations of these changes, however, remain speculative and ultimately tied to the Earth's thermal evolution and the historical development of convective patterns in the Earth's mantle, as well as to the presence or absence of wholemantle convection (e.g., Kellogg et al., 1999) (sections 3.2, 3.4, and 3.6).

3.8.

THE LIMPOPO BELT OF SOUTHERN AFRICA: A NEOARCHAEAN TO PALAEOPROTEROZOIC OROGEN

A.J. BUMBY AND R. VAN DER MERWE Introduction

The ENE-trending Limpopo belt of southern Africa, a granulite-grade metamorphic belt (see also section 3.9) believed to represent one of the earliest preserved examples of a collisional orogeny (see sections 3.6 and 3.7), passes through the northern-most part of South Africa and the southern-most part of Zimbabwe, westwards into eastern Botswana (Fig. 3.8-1). It has a length of c. 550 km and is c. 250 km wide. The orogenic event reflects a collision between the northern edge of the Kaapvaal craton and the southern edge of the Zimbabwe craton (present day orientation; e.g., Light, 1982; Treloar et al., 1992), with the involvement of a third exotic terrane, which has become sandwiched between the two cratons. After collision, the Kaapvaal and Zimbabwe cratons formed the Kalahari craton. The Limpopo belt is divided into three sub-parallel ENE-trending zones (Fig. 3.8-1). The Southern Marginal zone (SMZ) is considered as part of the greenschist- to amphibolite-grade northern Kaapvaal craton, though it is generally at a higher (granulite facies) metamorphic grade, and appears to have been exhumed from deeper levels. The southern The Precambrian Earth: Tempos and Events Edited by EG. Eriksson. W. Altermann, D.R. Nelson, W.U. Mueller and O. ('atuneanu

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Fig. 3.8-1. Simplified geological map and cross-section showing features of the Limpopo Belt (after Kr6ner et al., 1999; Roering et al., 1992).

edge of the SMZ, where it borders the Kaapvaal craton, is marked by the southwardsvergent Hout River shear zone. This thrust zone is thought to have accommodated the exhumation of the SMZ, and may bear the ortho-amphibole isograd in its hangingwall (Van Reenen et al., 1987; C.A. Smit et al., 1992). The SMZ is separated from the Central zone (CZ) by the Palala shear zone, which consists of a c. 15 km wide mylonitic zone in the type area of the Koedoesrand window (Fig. 3.8-1). The Palala shear zone generally has a sinistral sense of shear, though it has been reactivated locally in an opposite sense (McCourt and Vearncombe, 1987, 1992). Schaller et al. (1999) argue that the Palala has a

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predominantly dextral sense of shear. Eastwards, the Palala shear zone becomes covered by younger Proterozoic sedimentary rocks; in the Blouberg area (Fig. 3.8-1), outcrops beneath the younger cover show that the Palala shear zone is occupied by migmatitic gneiss, presumably reflecting deeper crustal levels of the shear zone (Bumby, 2000). Still further eastwards, the Palala shear zone may be represented by the Tshipise straightening zone (Bahnemann, 1972), which runs parallel to, and north of the Proterozoic-aged Soutpansberg Group (Fig. 3.8-1). The contact between the CZ and Northern Marginal zone (NMZ) is marked by the Triangle and Magogaphate shear zones, though much of their length is covered by Phanerozoic (Karoo) strata. The Triangle shear zone has a dextral sense of shear (McCourt and Vearncombe, 1987, 1992). Like the SMZ, the NMZ consists of high-grade, reworked granite-greenstone material of the adjacent (Zimbabwe) craton (Fig. 3.8-1), and is locally intruded by charnockites and enderbites. The NMZ is bound to the north by the northwards-vergent North Marginal thrust zone. Thus, a north-south cross-section of the Limpopo belt can be considered as being almost symmetrical in structure, though the Palala shear / Tshipise straightening zones dip much more steeply than the Triangle / Magogaphate shear zones (de Beer and Stettler, 1992) (Fig. 3.8-1). The Central zone contains entirely different lithologies than either of the Marginal zones, and is thus considered as a separate terrane, somehow incorporated into the orogen (see below). It is composed of a range of rocks, including supracrustal (continental platform) strata at granulite-grade, known as the Beit Bridge Complex (S.A.C.S., 1980; Van Reenen et al., 1992). The latter consists of metapelitic gneiss, marble, calc-silicate rocks, quartzite and magnetite quartzite (Van Reenen et al., 1992). In addition, there are several suites of quartzo-feldspathic and tonalitic and trondhjemitic grey gneiss, which may be basement to (e.g., Sand River Gneiss) or intrusive into (e.g., Bulai Gneiss) supracrustal gneiss of the Beit Bridge Complex. Mafic metamorphic rocks are relatively rare in the Beit Bridge Complex, and are intrusive into the earlier formed gneisses (Brandl, 2001). Timing of the Limpopo Event

The Limpopo belt has, in the past, generally been regarded as having formed due to collision of the Zimbabwe and Kaapvaal cratons during the late-Archaean, at c. 2.65 Ga. This view is based on ages determined from granitic gneiss bodies within the Central zone, which are intrusive into the Beit Bridge Complex (e.g., the Bulai Gneiss). The age of the Bulai is important as it cross-cuts, and contains xenoliths of foliated granulite-grade Beit Bridge rocks. As such, the intrusion of the granitic magma was thought to post-date peak granulite-grade metamorphism (McCourt and Armstrong, 1998). The Bulai Gneiss has been interpreted to have been intruded as a result of decompression melting during exhumation following the collisional event, thus the age of the granitic gneiss provides a minimum age for collision. Typically, ages between 2570 and 2664 Ma are derived from granitic gneisses in the Central zone (e.g., McCourt and Armstrong, 1998). However, Holzer et al. (1998) argue that although the Bulai Gneiss does indeed cross-cut high-grade metamorphic fabrics in the Beit Bridge Complex, the peak granulite event post-dates the

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Bulai Gneiss, and has an age of c. 2.01 Ga. Such an age for peak metamorphism suggests an early Proterozoic age for the timing of the main Limpopo collision. Recent work using alternative dating methods on metamorphic mineral assemblages (Pb stepwise leaching of metamorphic silicates and SHRIMP) has strengthened these new interpretations of a Proterozoic tectonic history for the Limpopo belt. Evidence for not one, but three high-grade metamorphic events is recorded within the Limpopo belt, at 3.2-3.1 Ga, 2.65-2.52 Ga and 2.0-+-0.05 Ga (Barton et al., 1994; Holzer et al., 1998; KrOner et al., 1999). Such data suggest that the evolution of the Limpopo belt was considerably more complex and occurred over a greater period of time than was thought previously. Each zone within the Limpopo belt appears to have a reasonably distinct tectono-metamorphic history (Holzer et al., 1998; Kr6ner et al., 1999), which is difficult to reconcile with the premise that adjacent terranes might be expected to have a common history after their tectonic juxtaposition. Importantly, Jaeckel et al. (1997), Holzer et al. (1998), and KrOner et al. (1999) all suggest that of these three high-grade events, the c. 2.0 Ga event reflects the main metamorphism and deformation in the Central zone, and thus the age of the collisional event. Holzer et al. (1998) do not draw conclusions regarding the relative positions of the Kaapvaal, Central zone and Zimbabwe provinces throughout the complex evolution of the Limpopo belt, but they reconstruct geological events in each of the three zones through time. The middle Archaean event (3.2-3.1 Ga), recorded only in the Central zone, reflects earliest metamorphism; as it has no correlates in either of the marginal zones, it can be viewed as an event which occurred when the Central zone was an isolated terrane. The late Archaean event (2.65-2.52 Ga), is recorded in all three zones, and reflects regional granulite-facies metamorphism, though this may have occurred up to 100 My earlier in the SMZ than in the CZ or NMZ. However, Barton et al. (1994) argue that the SMZ granulite-grade event was at 3.15 Ga (Pb-Pb analyses), rather than at c. 2.6 Ga. Finally, the Proterozoic event (2.0-t-0.05 Ga) is reflected by granulite-grade metamorphism in the CZ, low-grade overprinting in the NMZ, with no evidence for Proterozoic metamorphism in the SMZ (Holzer et al., 1998; Kreissig et al., 2001). Holzer et al. (1998) suggest that the clockwise p-T-t paths determined for the Proterozoic event most likely correspond to the actual dextral, transpressive orogeny between the three terranes at c. 2.0 Ga. Opponents of the Proterozoic collisional age rather view the 2.0 Ga event as reflecting reactivation within the previously (late Archaean) assembled Limpopo belt (e.g., McCourt and Armstrong, 1998; Bumby et al., 2001a), possibly related to similarly-aged tectonic events elsewhere on the Kalahari craton (e.g., the Magondi orogen to the northwest of the Zimbabwe craton; McCourt and Armstrong, 1999; Bumby et al., 2001a). However, such reactivation interpretations are difficult to reconcile with the presence of concordant leucosomes and cross-cutting melt patches throughout the Central zone, which have ages between 2.03 and 2.01 Ga (e.g., Jaeckel et al., 1997), which implies more substantial tectonism than mere reactivation. The strong evidence for the Proterozoic age of the Limpopo orogeny has been questioned by Bumby et al. (2001 a), who studied a sequence of unmetamorphosed sedimentary strata deposited above the Palala shear zone in the Blouberg area. Here, the high-grade

3.8. Limpopo Belt of Southern Africa

221

Limpopo gneiss is non-conformably overlain by the Blouberg Formation (sandstones and conglomerate), which is in turn non-conformably overlain by the Waterberg Group (sandstones and conglomerate), in turn succeeded by the c.1.85 Ga Soutpansberg Group (volcanics and quartzite) (Barton, 1979; Cheney et al., 1990). Such a complex and long-lasting tectono-depositional history recorded in these non-metamorphosed sediments, to be accomplished prior to Soutpansberg volcanism at c. 1.85 Ga, is not easy to reconcile with such a young (i.e., c. 2 Ga, Proterozoic) age for Limpopo collision tectonics (Bumby et al., 2001a); they thus favour an older (Neoarchaean) age for the assembly of the Limpopo belt. Evidence for a c. 2.0 Ga event is interpreted from syn-tectonic sedimentation and southwards-vergent brittle faulting in the lowermost Blouberg Formation, though such structures cannot be related to a high-grade collisional event. Similarly, Kreissig et al. (2001) find little evidence for Proterozoic high-grade events in the Southern Marginal zone, just to the south of Blouberg. However, given the imprecise dates of the Soutpansberg strata (Rb-Sr whole-rock; Barton, 1979), and the overwhelming evidence in favour of a c. 2.0 Ga high-grade event in the Central zone, it seems necessary to fit the long tectonosedimentary record in the Blouberg area into a period post-dating the 2.0 Ga event. Gerya et al. (2000) estimate that exhumation from a burial depth of c. 30 km for the Central zone of the Limpopo belt (which needs to be accomplished prior to the unconformable deposition of the Blouberg Formation) was accomplished within only 10 My.

Models for the Evolution of the Limpopo Belt Van Biljon (1977) and Light (1982) suggested Tibetan-style collision for the evolution of the Limpopo belt. Light (1982) considered the belt to have formed as a response to southover-north-directed oblique collision, with a southwards-dipping subduction zone located beneath the Kaapvaal craton. In this model the Central zone is considered as a southern peripheral margin of the Zimbabwe craton. The subsequent identification of the CZ as an entirely separate terrane led to further refinements of this model. Watkeys (1984) suggested that the Zimbabwe craton and CZ had first collided at c. 2.7 Ga, and that this was followed by juxtaposition with the Kaapvaal craton along the Palala shear zone at c. 1.9 Ga. McCourt and Vearncombe (1987, 1992) proposed that the CZ had been emplaced westwards/south-westwards as a giant nappe over the previously accreted Kaapvaal and Zimbabwe cratons at about 2.7 Ga. This was based upon the sinistral and dextral sense of movement in the Palala and Magogaphate-Triangle ("Tuli-Sabi") shear zones, respectively, which are thought to have acted as lateral ramps. Alternatively, van Reenen et al. (1987) proposed that all three terranes had collided together at about 2.7 Ga. In this model, the Palala and Triangle/Magogaphate shear zones are viewed as being caused by later, post-Bushveld Complex (< c. 2050 Ma) lateral shearing within the Limpopo belt, and not as tectonic sutures themselves. De Wit et al. (1992) envisage the Limpopo belt as having formed by tectonic juxtaposition during two late-Archaean episodes. Initially, the CZ was thrust northwards onto the Zimbabwe craton, accompanied by northwards thrusting of the Kaapvaal craton. This was followed by a period of southwards-directed thrusting with east-west transcurrent fault-

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Chapter 3: Tectonism and Mantle Plumes Through Time

ing accompanying "pop-up" type tectonics. Similarly, Roering et al. (1992) consider the Limpopo belt to have formed as a result of Himalayan-style continental collision, followed by the occurrence of a regional-scale pop-up, accommodated along the various Limpopo shear zones. Rollinson (1993) considered the Limpopo belt to have formed as a response to gradual accretion onto the Kaapvaal craton. The SMZ, CZ, NMZ and Zimbabwe cratons are considered as separate terranes, accreted southwards in a piecemeal manner onto the northern edge of the Kaapvaal craton. This is different to most other models which envisage the Limpopo belt to have resulted from a single collision between stable continental blocks. In contrast to these various perpendicular collisional models, Kamber et al. (1995) and Holzer et al. (1998) suggest that the Limpopo belt may have formed as a result of dextral transpression. Amongst the high-grade rocks of the Southern Marginal zone are numerous ultramafic lithologies (Smit, 1984), some of which contain podiform chromite deposits (e.g., Lemoenfontein at 23~ S; 29~ ' E). Such podiform chromites are typical of oceanic crustal material (e.g., Arai, 1997) and so may record the presence of ophiolite material (see also Smit, 1984) within the SMZ. The presence of an ophiolite (section 3.7) at such deep crustal levels as that in the SMZ implies deep burial after obduction (R.K.W. Merkle, 2002, pers. comm.). Such burial may relate to tectonics after collision of the cratons (e.g., westwards-vergent nappes emplacing the CZ; McCourt and Vearncombe, 1987, 1992), or to much earlier accretion during the Archaean assembly of the Kaapvaal craton. It seems plausible that reconciliation of such varying opinions regarding the timing and tectonic models may be reached by considering the Limpopo belt to be a composite orogen, having undergone collision at two separate times. Evidence from the SMZ shows less support for a high-grade Proterozoic event, but rather favours earlier collision at c. 2.65 Ga (Kreissig et al., 2001), with a degree of reactivation at 2.0 Ga. This may suggest that the Central zone and Kaapvaal craton collided in the Archaean. Schaller et al. (1999) found strong evidence for a 2.0 Ga high-grade dextral reactivation in the Palala shear zone, which appears to overprint the older, sinistral fabric (Broekhuizen, 1998). Evidence from the Palala shear zone and CZ, along with data from the Triangle shear zone and NMZ therefore shows strong support for a high-grade 2.0 Ga dextral transpressive collisional event (e.g., Holzer et al., 1998), which may reflect the collision of the previously juxtaposed CZ and Kaapvaal craton together with the Zimbabwe craton (as proposed by Roering et al., 1992; Fig. 3.8-1). However, high-grade dextral reactivation in the Koedoesrand area (Schaller et al., 1999) is not easy to reconcile with southwards-vergent brittle tectonics proposed only 70 km eastwards in the Blouberg area, only about 50 My later (Bumby, 2000; Bumby et al., 2001 a). At present, thus, it is difficult to reconcile either the timing of the Limpopo belt or the tectonic setting(s) of the deformation with any modern style of plate tectonic collision. An examination of other potential tectonic models which may have driven collisional orogens up to and during the early Proterozoic, and how they might relate with modern plate tectonic processes therefore seems pertinent (see also sections 2.4 and 3.5-3.7). As potentially one of the oldest well-preserved Precambrian orogenic belts, the Limpopo re-

3.9. G e o d y n a m i c CrustaI Evolution

223

gion may hold the answers to questions regarding the similarity between Proterozoic and modern-style tectonic processes.

3.9.

GEODYNAMIC CRUSTAL EVOLUTION AND LONG-LIVED SUPERCONTINENTS DURING THE PALAEOPROTEROZOIC: EVIDENCE FROM GRANULITE-GNEISS BELTS, COLLISIONAL AND ACCRETIONARY OROGENS

M.V. MINTS AND A.N. KONILOV Introduction

Proterozoic palaeomagnetic poles from the major shields point to a single apparent polar wander path (APWP) (Piper, 1983), which supports a possible single coherent continental lithospheric plate from c. 2.9 Ga to c. 1.1 Ga. However, the APWP method has intrinsic problems, such as large uncertainties in palaeopole ages and large gaps in the APWP record (e.g., Buchan et al., 1996). Consequently, geodynamic reconstructions of the history of early Precambrian supercontinents are based mostly on geological considerations; however, models reflect significantly different understandings of key geological structures, especially orogenic belts (e.g., Ga~il, 1992; Rogers, 1996; Condie, 1998). Geochronological data demonstrate episodicity in Palaeoproterozoic geological evolution, preceded by a prominent 2.7 Ga peak in the geochronological record, postulated to reflect creation of the first supercontinent (section 3.2) or a small number of composite continents. Palaeoproterozoic crustal evolution encompassed at least incomplete disruption of the supercontinent(s) (Khain and Bozhko, 1988; Mints, 1998; Condie, 2002a), commencing at c. 2.5 Ga. Reassembly at c. 1.75-1.65 Ga followed increased production of juvenile continental crust, which began at c. 1.9 Ga, followed by rapid accretion of arc systems at 1.88-1.84 Ga. Geochronological data also indicate a prolonged period of very low magmatic activity within continental areas between 2.45 and c. 2.1 Ga (Condie, 1998) (this is supported by an apparent lack of large igneous provinces at c. 2.4-2.2 Ga; section 3.3). Palaeoproterozoic juvenile assemblages dominate within two types of mobile belt: (1) low-grade (greenschist to low-temperature amphibolite facies) volcano-sedimentary and volcano-plutonic belts; analogous Archaean belts are generally termed greenstones e.g., sections 2.3, 2.4, 3.6, 4.3 and 4.4), and (2) high-grade (high-temperature amphibolite to ultra-high temperature granulite facies) "granulite-gneiss" belts (see also section 3.8). The former belts are interpreted as sutures (collisional orogens) or collapsed continental rifts. Extended volcano-plutonic assemblages at the margins of ancient continents are usually termed accretionary orogens (e.g., Hoffman, 1989c; Windley, 1992) (see also section 3.6). However, ideas on the nature and tectonic and geodynamic significance of granulitegneiss belts remain controversial. Structural constraints indicate that many large-scale Palaeoproterozoic granulite terranes evolved within a broadly collisional context (e.g., The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, I).R. Nelson, W.U. Mueller and O. Catuneanu

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section 3.8). Charnokite and enderbite intrusions geochemically resembling granitoids of the calc-alkaline and tonalite-trondhjemite series have been ascribed in many cases to arc environments (e.g., Van Kranendonk, 1996; Jackson and Berman, 2000). Consequently, granulite-gneiss belts have been interpreted as sutures/collisional orogens also (e.g., Hoffman, 1989c; Rosen et al., 1994; Daly et al., 2001). However, many features of granulite-gneiss belts are in conflict with this interpretation; reassessment of existing geodynamic models of the granulite-gneiss belts and their relationships with low-grade belts is thus one of the aims of this paper. A new evolutionary model is discussed here, emphasising the interaction of plume- and plate tectonic-related crustal-forming processes during the Palaeoproterozoic (see also section 3.6). Magmatic and thermal activity during the early Palaeoproterozoic was largely concentrated within Laurentia (defined here as comprising North American and Fennoscandian cratons; Condie, 1990). The Fennoscandian crustal segment forms the northwestern part of the present-day East European craton, which also includes the Volgo-Uralian and Sarmatian segments (see Fig. 3.9-2b). The latter segments existed as separate cratons by c. 2.05-2.0 Ga; their amalgamation with Fennoscandia and Laurentia as a whole may have occurred at c. 1.7 Ga (Bogdanova et al., 2001a). In contrast, late Palaeoproterozoic assemblages are distributed within all continents. The Laurentian and Siberian cratons, where post-Palaeoproterozoic reworking was limited, are utilised here as type examples. The temporal and spatial distribution of the main Palaeoproterozoic assemblages within mobile belts of Laurentia and Siberia are shown in Fig. 3.9-1. Geological Evolution of Laurentia The supercontinent interior 2.51-2.44 Ga superplume event and initial rifling of the Archaean Supercontinent. The Palaeoproterozoic evolution of Laurentia (Figs. 3.9-1 and 3.9-2) at c. 2.5 Ga was marked by

Fig. 3.9-1. Correlation between main events in the evolution of the Palaeoproterozoic mobile belts in Laurentia and Siberia. lmpoorly known continental crust of suggested arc-related affinity; 2mmafic dykes; 3--alkaline mafic-ultramafic intrusions; 4--layered mafic-ultramafic bodies; 5--gabbro-anorthosites; 6---sedimentary successions with volcanic intercalations, within platform cover and rift-related basins; 7--rift-related volcano-sedimentary successions; 8--granulite-gneiss complexes; 9--within-plate granites; 10---MORB-like volcanic rocks and ophiolite assemblages; 11--oceanic island mafic-ultramafic rocks; 12~island arc volcanic rocks; 13--arc-related granitoids; 14-15~time lines: first-order (14) and second-order (15). Shading of columns: blank -- granulite-gneiss belts; light grey -- accretionary orogens; grey = low-grade volcano-sedimentary belts; dark grey = passive margins. Abbreviations of orogens: T-T-Taltson-Thelon; CBcB--belts about the Cumberland Batholith; TH--Trans-Hudson; EA---eastern North American realm (K, Ketilidian; P, Penokean; Y, Yavapai-Mazatzal); Sv-ac~Svecofennian accretionary orogen; BSB--Belarus-South Baltica; S-p--Svecofennian passive margin; PV-CKB--Pechenga-Varzuga and Circum-Karelian belts; LGB~Lapland granulite belt; A--Akitkan belt; StanmStanovoy Province.

3.9. Geodynamic Crustal Evolution

Fig. 3.9-1.

225

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Chapter 3: Tectonism and Mantle Plumes Through Time

Fig. 3.9-2. Reconstruction of the Laurentian part of the Palaeoproterozoic supercontinent (the North American map is modified after Hoffman, 1989). (a) suggested 2.51 Ga Superior-Karelia junction and East European craton displacement to 2.45 Ga (the 2.45 Ga position is based on Mertanen et al., 1999). emplacement of mafic-ultramafic bodies and dykes. Layered peridotite-gabbro-norite bodies intruded into the upper crust at 2.51-2.44-2.40 Ga (Alapiety et al., 1990; Amelin et al., 1995; EE Mitrofanov, 2002, pers. comm.). Gabbro-anothosites intruded at 2.49-2.43 Ga into lower crust (Frisch et al., 1995; Mitrofanov et al., 1995; Kislitsin et al., 2000), being accompanied by granulite and high-grade amphibolite facies metamorphic assemblages. Emplacement of the 2.46 Ga Kolvitsa massif (Kola Peninsula) in an extensional setting was followed immediately by multiple mafic dyke injections. All the dykes contain xenoliths of sheared and metamorphosed gabbro-anorthosite (Balagansky et al., 2001). Metamorphic conditions of 700-900~ and 10-12 kbar, corresponding to depths of c. 35-43 km, have been determined for the gabbro-anorthosites together with hosting mafic granulites, and for small gabbro and diorite bodies in adjacent areas (Bogdanova, 1996; Glebovitsky et al., 1997; Balagansky et al., 2001). Initial rifting of the Archaean crust along the southeastern margin of the Superior craton was associated with unconformable deposition of the Huronian Supergroup (section 5.6). The lowest unit consists of flood basalts, felsic lavas and arkosic metasediments. The base of this volcano-sedimentary prism is cut by 2.49-2.46 Ga, within-plate, gabbro-

3.9. Geodynamic Crustal Evolution

227

Fig. 3.9-2 (continued). (b) the Palaeoproterozoic mobile belts in the Laurentian Craton. Archaean cratons: K, Karelia; Su, Superior. Palaeoproterozoic belts and tectonic zones: BB, Belarus-South Baltian; BS, Boothia Peninsula-Somerset Island; CB, Cumberland batholith; CK, Circum-Karelian; CS, Cape Smith; F, Foxe (Piling and Penrhyn Groups); H, Huron; Ke, Ketilidian; L, Lapland (branches: LI, Lapland belt sensu stricto; L2, Kolvitsa-Umba; L3, Moscow); LH, Lake Harbour; MC, Midcontinent; M-L, Makkovik-Labradorian; N, Nagssugtoqidian (Ussuit complex); NQ, New Quebec; P, Penokean; PV, Pechenga-Varzuga; R, Rinkian (Karrat Group); S, Snowbird; T, Taltson-Thelon (T~, Taltson; T2, Thelon); TH, Trans-Hudson; To, Torngat (Tasiuyak complex); Sv, Svecofennian; YM, Yavapai-Mazatzal. anorthositic, mafic-ultramafic and granite intrusions, accompanied by the 2.47-2.45 Ga Mattachewan dyke swarm. Upper Huronian units, consisting of glaciogenic conglomerate, mudstone, siltstone and carbonate rocks, capped by cross-bedded sandstone, are cut by a 2.22 Ga diabase suit (Hoffman, 1989c; Corfu and Easton, 2000). Heaman (1997) suggested a Superior-Karelia connection, plume-related rifting at 2.45 Ga, further spreading of the Mattachewan ocean and corresponding breakup of the pre-2.5 Ga continent, with transformation of the Huronian sedimentary basin to a rifted passive margin. Although palaeomagnetic data do not support this model (Mertanen et al., 1999; Buchan et al., 2000), we note the following aspects: (1) a striking uniformity of

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the 2.51-2.44 Ga Laurentian mafic-ultramafic magmatism, concomitant with its restriction to the Superior and Karelia areas (Vogel et al., 1998); (2) analogy of the 2.49-2.22 Ga Huronian and 2.5-2.3 Ga Lapponian (Karelia) shelf-type sequences; (3) the structural style of the reconstructed palaeocontinent includes oval, concentric mobile belts that permit speculation on inherited initial plume geometry (Fig. 3.9-2); (4) evidence for a SuperiorKarelia connection at 2.51-2.50 Ga, predates the palaeomagnetic data of Mertanen et al. (1999) by c. 50-70 My; (5) displacement and rotation permit superposition of Fennoscandia above the 2.44 Ga position based on palaeomagnetic data (Fig. 3.9-2).

2.44-2.0 (2.1l) Ga quiescent within-plate development.

The period from 2.44 to about 2.0 Ga was tectonically quiescent not only in Laurentia but also worldwide (Condie, 1998) (see also, sections 3.3 and 3.7 for indications of superplume quiescence). Sedimentary-volcanic successions with c. 2.44-2.4 Ga, predominantly mafic eruptive rocks occur in the eastern Fennoscandian Shield, within the Pechenga-Varzuga and Circum-Karelian belts (Melezhik and Sturt, 1994; Mints et al., 1996; Puchtel et al., 1996; Sharkov and Smolkin, 1997). Recent seismic data indicate that these belts are formed by trans-crustal overthrust-underthrust packages (Mints at al., 1996, 2001, 2002; Smithson et al., 2000). Initial volcanism in the eastern part of the Pechenga-Varzuga belt comprised low-K20 tholeiites associated with shallow water sandstones, succeeded by lacustrine quartzites and predominant low-Ti andesite-basalts, less abundant low-Ti and high-Mg (komatiitic) basalts, and rare andesites and dacites. A suite of 2.44-2.42 Ga dacitic rhyolites and 2.32 Ga subaerial andesites terminates this part of the sequence (Mitrofanov et al., 1991; Amelin et al., 1995; Mitrofanov and Smolkin, 1995). A suite of 2.41 Ga komatiitic basalts and andesite-basalts predominates in the Vetreny belt (southeastern branch of the Circum-Karelian belt) (Puchtel et al., 1996). Geochemistry of these various mafic lavas suggests plume or T-MORB-related sources, although some features can be attributed to subduction-related processes or to crustal contamination. The second part of this period commenced at c. 2.2 Ga with extrusion of alkaline basalts, trachybasalts with subordinate alkaline picrites, and rhyolites. They are associated with red epicontinental sediments, thick greywacke turbidites, ironstones, siltstones, black shales, dolomites, stromatolitic carbonate rocks, and phosphate-bearing and Mn-rich lithologies within the Pechenga structure (Mitrofanov and Smolkin, 1995), at the rifted passive margin of the Superior craton, in the New Quebec belt and in the Trans-Hudson orogen (Hoffman, 1989c). Between 2.3 and 2.1 Ga a significant part of the Karelian craton was covered by continental shelf sediments which were intruded by diabase sills at 2.2-1.97 Ga (Vuollo, 1994). Similarly, 2.22-1.99 Ga dyke swarms are common in North America (Ernst and Buchan, 2002a, and references therein).

(2.11) 2.0-1.95 Ga superplume event.

The second Palaeoproterozoic superplume event (sections 3.2 and 3.3) was more widespread. Plume-related rift-like extension was followed by transition to oceanic spreading in areas of future collisional orogens. Indisputable evidence for ocean growth and closure comes from the 2.0 Ga Purtunq (Scott et al., 1991) and 1.95 Ga Jormua ophiolites (Kontinen, 1987) (northern margin of the Superior craton and western Karelia craton, respectively) and the 2.0 Ga Kittil~i ophiolite-like com-

3.9. Geodynamic Crustal Evolution

229

plex (see also section 3.7) in northern Finland (Hanski et al., 1998). The 2.11-1.96 Ga T-MORB-type pillowed tholeiites in the Pechenga structure are interlayered with picritic lavas and felsic ash-flow tufts that are thought to have erupted from oceanic island volcanoes at 1.99-1.96 Ga. Subvolcanic, Cu- and Ni-bearing gabbro-wehrlite bodies of the same age cut the lower part of the MORB-type succession, but mostly are tectonically included in the accretionary prism known as the "Productive Layer" of the Pechenga ore field (Melezhik and Sturt, 1994; Mints et al., 1996; Sharkov and Smolkin, 1997). The submarine flood basalt province originated at the southeastern margin of the Karelia craton at 1.98 Ga (Puchtel et al., 1998b). A younger generation of rift-related gabbro-anorthosite intrusions, known at present within the Lapland granulite belt, was emplaced at 2.0-1.95 Ga at the same crustal level as the above-mentioned older generation (Bernard-Griffiths et al., 1984; Kaulina, 1999; Nerovich, 1999; Mints et al., submitted).

1.95-1.75 (1.71) Ga combined plume- and plate tectonic-related evolution. The youngest, 1.87-1.86 Ga assemblages of the Pechenga-Varzuga belt include rhyolites, dacites, andesites, picrites, high-Mg basalts of suprasubduction affinity and subordinate N-MORBlike basalts associated with volcaniclastic sediments, black shales and sporadic cherts (Melezhik and Sturt, 1994; Sharkov and Smolkin, 1997). Circa 1.96 Ga felsic calc-alkaline gneisses have been discovered in the southern vicinity of the Pechenga-Varzuga belt (Daly et al., 2001). The Ticksheozero alkaline mafic-ultramafic-carbonatite complex intruded in northern Karelia at 1.85 Ga (Belyatsky et al., 2000). Abundant subduction-related assemblages are known within the c. 500 km-wide Trans-Hudson belt. Arc magmatism from c. 1.92 to 1.88 Ga was followed by the early outboard accretion of oceanic elements at c. 1.87 Ga, and extensive plutonism between c. 1.86 and 1.83 Ga (Hoffman, 1989c; Stern and Lucas, 1994) (section 3.5). An extensive basin formed approximately at the same time in the area between the Pechenga-Varzuga and North-Karelian volcano-sedimentary belts. The metasedimentary and meta-igneous rocks deposited therein can be observed presently within the Lapland granulite belt sensu lato (LGB). Ages of detrital zircons from the metasediments range from 2.71 to 1.88 Ga; together with Sm-Nd data, this indicates that the youngest sediments were deposited at c. 1.9 Ga (Huhma and Meril~inen, 1991; Sorjonen-Ward et al., 1994; Balagansky et al., 1998; Daly et al., 2001; Bridgwater et al., 2001). Taking into account the 2.0-1.95 Ga age of a younger gabbro-anorthosite generation, sedimentation must have begun not earlier than 2.0 Ga. The relationship of this succession to subduction processes in the neighbouring volcano-sedimentary belts remains unclear. High-grade metamorphism followed immediately after emplacement of the gabbroanorthosites of each generation. The corresponding events (Mll at 2.46-2.43 Ga and an M12 at c. 1.95 Ga; Kaulina, 1996, 1999; Nerovich, 1999) at similar conditions of 860-960~ 10.3-14.0 kbar (Fig. 3.9-3) and restricted to the lowest part of the belt, were practically coeval with the intrusions. The M2 event was characterised by lower p - T conditions: 800-860~ 5.8-12.4 kbar. For the most pervasive M3 metamorphism, the p - T conditions were 640-770~ at 4.8-10.7 kbar. The M2 and M3 events, which affected both igneous and sedimentary rocks, date at 1925 Ma and 1917-1902 Ga, respectively

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(Bibikova et al., 1993; Kaulina, 1996, 1999; Kislitsyn et al., 1999). The age of thrusting and related exhumation was c. 1.89-1.87 Ga (Bibikova et al., 1993; Alexeyev et al., 1999; Tuisku and Huhma, 1999; Daly et al., 2001). Rapid overthrusting of the hot crustal slices during the M4 stage (550-650~ at 4.5-8.4 kbar) at 1.87 Ga resulted in inverted metamorphic zoning in the "para-autochthonous" rocks. Deep-seated garnet granulite-hosted xenoliths of norite and gabbro-norite to gabbro-anorthosite composition (Neymark et al., 1993; Kempton et al., 1995; Vetrin, 1998), found in Devonian pipes (Kempton et al., 2001) and which are akin to the mafic granulites of the LGB, suggest that the granulite-facies metamorphism affected the rocks within a thick crustal section from c. 70 to c. 25 km depth (see Fig. 3.9-3). Tectonic decoupling of the continental crust occurred during the Palaeoproterozoic collision: a crustal slice c. 40 km thick was detached and overthrust, whereas the lowest crustal rocks were preserved in their original position (Mints et al., 2000a, submitted). Geophysical data indicate that the continuation of the LGB beneath the platform cover of the East European Craton forms an extended arch-shaped system of belts approximately 2000 km long (Fig. 3.9-2). Near Moscow the thrust-nappe structure of these belts was recognised recently from reflection seismic profiling (Berzin and Mints, unpublished data). Petrological and geochronological studies of drill core samples yielded a 1.98 Ga magmatic age of the leuconorites and enderbites that were emplaced shortly before metamorphism. A preliminary estimate of peak metamorphic parameters of c. 1000~ and c. 10-12 kbar, as well as the age data (Bogdanova et al., 1999) match almost exactly the characteristics of the LGB assemblage in the Kola Peninsula. The Taltson-Thelon orogenic belt of northern Canada, consisting of the Taltson Magmatic zone (TMZ) and Thelon orogen (TO), resembles the Lapland granulite belt in many aspects of age, composition and metamorphism. The TMZ is dominated by 1.99-1.96 Ga I-type and 1.95-1.93 Ga S-type granitoids, associated with granulite-facies metasedimentary rocks (6-8 kbar and > 900~ It has been suggested that evolution of the TMZ involved subduction of oceanic crust beneath the Archaean Churchill Province at 1.99-1.96 Ga, followed by collision between this province and the 2.4-2.0 Ga Buffalo Head terrane (Hoffman, 1989c; McDonough et al., 1995; Ross and Eaton, 2002). However, there is new evidence that both I- and S-type granitoids had an exclusively intra-crustal origin (De et al., 2000). Together with the high-temperature metamorphism, this supports a within-plate (Chacko et al., 1994; Farquar et al., 1996) and possible plume-related origin for the TMZ. The TO was interpreted by Hoffman (1989c) as a product of dextral oblique collision between the Slave (foreland) and Rae (hinterland) Archaean provinces. Predominant granodiorites intruded into the high-grade country rocks (Hoffman, 1989c; Thompson, 1992), and the high-grade terranes of Boothia Peninsula-Somerset Island and Devon-Ellesmere Islands, which form the northern extension of the TO, are built of mainly Palaeoproterozoic ortho- and paragneisses (graphitic metasediments and marbles) with lenses of mafic and ultramafic granulites. The 1.9 Ga metamorphism at 740-900~ and 6-8 kbar was accompanied by syenitic magmatism and anatexis (Kitsul et al., 2000). Similarly, a branching system of Palaeoproterozoic mobile belts in northeastern North America and southern Greenland, centred on the 1.87-1.85 Ga charnockitic Cumber-

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Fig. 3.9-3. p--T evolution of the Lapland granulite belt and corresponding deep crustal section (after Mints et al., submitted). land batholith and comprising the Nagssugtoqidian, Rankain, Foxe and Torngat belts and some tectono-stratigraphic units (Lake Harbour Group, Narsajuaq arc, Ramsay River orthogneisses; St-Onge et al., 1999), can be reassessed in the light of recent geochronological and petrological studies (Taylor and Kalsbeek, 1990; Kalsbeek and Nutman, 1996; Van Kranendonk, 1996; Kalsbeek et al., 1998; Nutman et al., 1999; Scott, 1999; Jackson and Berman, 2000). These belts are formed mainly by granulite gneisses with inferred metamorphic temperatures having reached 950~ and pressures from c. 4 to c. 12 kbar. Protoliths of the lower parts of the metasedimentary sequences were predominantly platform- and rift-related rocks with subordinate evaporitic deposits, mafic and ultramafic volcanics and sills, and anorthositic bodies. The terrigenous metasediments were derived from 2.4-1.93 Ga juvenile Palaeoproterozoic precursors of unknown provenance with sig-

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nificant admixture of Archaean detritus. Sedimentation commenced at c. 2.0 Ga and was completed by 1.95-1.93 Ga or a little later. Earliest high-grade metamorphism in the lower crust is marked by the Sisimiut charnockite intrusion at 1.92-1.90 Ga, that is during or very soon after deposition of the sediments. The main pulse of high-grade metamorphism occurred at 1.85-1.80 Ga or a little earlier, and estimates of the completion of thrusting and exhumation vary from 1.85 to 1.74 Ga.

Accretionary orogens Svecofennian. From c. 2.5 to 2.0 Ga and up to 1.95 Ga, the western (present geographic coordinates) slope of the Kola-Karelia Province was covered by clastic sediments, dolomites and black shales intercalated with tholeiitic lavas (Lapponian and Kalevian Groups), characteristic of rifted passive margins. Thereafter, the Svecofennian accretionary orogen was formed over a short period of time, mainly by 1.93-1.86 Ga subduction-related mafic, intermediate and felsic volcanics, terrigenous, carbonate and volcaniclastic sediments, large granite plutons and, locally, by the high-Ti within-plate tholeiites that indicate rifting of mature arcs (Ga~il and Gorbatschev, 1987; Pharaoh and Brewer, 1990; Korsman et al., 1999 and references therein). Enhanced heat flow along the margin of the Karelia craton was responsible for the 1.89-1.81 Ga high-grade metamorphism (up to 800~ and 4-5 kbar) of some of the turbiditic rocks, possibly deposited in a back-arc environment (H61tt~i, 1988; Korsman et al., 1999). Early- and late-orogenic granitoids, formed at 1.90-1.87 and 1.83-1.77 Ga, were succeeded by minor granite intrusions considered to be early post-orogenic (Ga~il and Gorbatschev, 1987). The latter were followed by 1.70-1.54 Ga gabbro-anorthosite-rapakivi granite magmatism, which can be interpreted as anorogenic in the context of Palaeoproterozoic evolution. The southern (Belarus-South-Baltian) part of the Svecofennian accretionary orogen is formed by a succession of alternating arcuate low- and high-grade belts (Gorbatschev and Bogdanova, 1993). The arc-related magmatism started earlier than in the main Svecofennian area; ages of volcanic and intrusive rocks vary between 2.10 and 1.80 Ga, becoming younger westwards. The granulite-facies metamorphism in the high-grade belts (up to 900~ and 8-10 kbar; Scridlaite and Motuza, 2001; Taran and Bogdanova, 2001 b) occurred at 1.82-1.80 and 1.79-1.78 Ga and, like the arc-related magmatism, tends to become younger westwards. The 1.63-1.61 Ga metamorphic event is inferred to be linked to anorogenic magmatism. The mostly juvenile metaterrigenous component of the granulitegneiss assemblages contains detrital zircons with ages between 2.45 and 1.98 Ga, and minor admixture of Archaean detritus (Claesson et al., 2001). Some of the detrital material was derived from a rather old Palaeoproterozoic source, which is unknown in this area. It is suggested that high-grade metamorphism was linked with back-arc and/or post-collisional extension of thickened crust. The thermal history includes a 1.55-1.45 Ga imprint related to anorogenic anorthosite-rapakivi granite magmatism (Bogdanova et al., 200 lb). Incorporation of the Belarus-South-Baltian juvenile terrane into the main body of the Palaeoproterozoic crust started before 1.96 Ga (Claesson et al., 2001).

Penokean, Makkovik-Ketilidian, Labradorian and Yavapai-Mazatzal. The Penokean orogen is accreted to the southern edge of the Neoarchaean Superior craton (Fig. 3.9-2).

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It contains island arc and back-arc terranes built of deformed volcanic and sedimentary rocks and coeval 1.89-1.84 Ga gabbroic to granitic plutons. The Niagara and Eau Pleine suture zones between the Penokean orogen and the Archaean continent are bracketed between 1.86 and 1.835 Ga. In turn, the Makkovik-Ketilidian and Labradorian orogens extending along the southeastern margins of the Neoarchaean Herne, Rae and Nain Provinces, favour an Andean-type model. The lower succession, of passive margin affinity, contains terrigenous sediments intercalated with metabasalts and intruded by 2.13 Ga dykes and 1.91 Ga granodiorites. The upper sequence of active margin type, dominated by felsic and intermediate tufts and flows, was formed between 1.86 Ga and 1.81 Ga. Northwest-directed shortening accompanied by 1.81-1.80 Ga granodiorite and tonalite intrusions was completed by c. 1.79-1.76 Ga, immediately before rapakivi granite emplacement at 1.76 Ga. Collapse of the thickened crust that followed resulted in deposition of a within-plate volcano-sedimentary succession and bimodal anorogenic magmatism including gabbro-anorthosite-monzonite complexes, accompanied by 1.71-1.63 Ga granulitefacies metamorphism of the host rocks. The late Palaeoproterozoic juvenile crust in the southern Midcontinent and in the southwestern United States (Yavapai-Mazatzal-Midcontinent orogen)evolved in two general stages. Firstly, 1.79-1.71 Ga calc-alkaline volcano-plutonic terranes, interpreted as former island arcs and inter-arc basins, were amalgamated by about 1.70 Ga. Subsequently, subaerial felsic volcanism was followed by emplacement of 1.64-1.62 Ga post-tectonic granites and by Meso- and Neoproterozoic anorogenic magmatism (Hoffman, 1989c).

Wopmay. The presently west-facing Wopmay accretionary orogen extending along the western boundary of the Archaean Slave Province is formed by the poorly-known 2.4-2.0 Ga crust of the Buffalo Head and Hottah terranes, that were possibly accreted to the Archaean supercontinent at 2.4-2.3 Ga and between 1.9 and 1.7 Ga, respectively (Goff et al., 1986; Chacko et al., 2000; Ross and Eaton, 2002). This accretionary orogen includes a tectonically shortened passive margin sedimentary sequence along the western boundary of the Slave Province. Shelf-type sedimentation started at c. 1.97 Ga, and at 1.90-1.88 Ga rift-related bimodal magmatism occurred mainly along the off-shelf boundary. Some magmatic arc terranes consisting of 1.95-1.91 and 1.88-1.86 Ga juvenile crust were accreted successively to the Archaean continent. Some of them are underlain by cryptic 2.4-2.0 Ga (possibly 2.3-2.1 Ga) crust. Accretion was followed by post-orogenic, 1.86-1.84 Ga syenogranites and 1.71 Ga anorogenic rift-related diorites, gabbros, and subordinate anorthositic and syenitic rocks in the Yukon at the northwestern edge of the Wopmay orogen (Thorkelson et al., 2001). The process as a whole was terminated by a final phase of deformation, from 1.84 to 1.66 Ga (Hoffman, 1989c; Ross et al., 1991). Geological Evolution of Siberia The important difference between Siberia and Laurentia lies in the lack of Palaeoproterozoic low-grade volcano-sedimentary belts within the Siberian craton (Fig. 3.9-4), with the possible exception of the 2.03-1.82 Ga Akitkan belt (Gusev and Peskov, 1992; Rosen et al.,

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1994; Nozhkin, 1999). In contrast, Palaeoproterozoic granulite-gneiss belts are widely distributed within all tectonic provinces in Siberia. Rosen et al. (1994)argued that the origin of these belts was linked with the Palaeoproterozoic amalgamation of the Archaean microcontinents and collisional stacking of the crust. Frontal portions of these belts marked by blastomylonites, anorthosites and high-grade metamorphism have been directly interpreted as suture zones (Fig. 3.9-4). This concept coincides in the main with Hoffman's (1989c) model for the Palaeoproterozoic evolution of North America. The Anabar and Olenek Provinces comprise Archaean and Palaeoproterozoic rocks, the latter of which form the Daldyn and Hapschan granulite-gneiss belts, characterised by magmatic ages of 2.55 Ga for gabbro-anorthosite and 2.42 Ga for granulite-facies mafic metavolcanics. Metagreywackes and metacarbonates of the Hapschan belt were deposited in a shelf-related setting of a passive continental margin between 2.44 and 2.08-1.97 Ga. Based on Sm-Nd data, the source areas were composed of mature Palaeoproterozoic continental crust, which is not observed at the present-day erosion level (Zlobin et al., 2002). The high-grade metamorphic overprints affecting both Archaean and juvenile Palaeoproterozoic assemblages occurred at 2.18, 1.97, 1.94-1.90 and 1.80-1.76 Ga. Predominantly subalkaline granitoids intruded at 1.84-1.80 Ga (Rosen et al., 1994, 2000). Palaeoproterozic evolution in the Aldan Province began with emplacement of 2.49-2.40 Ga, within-plate, A-type granites (Rosen et al., 1994; Mints et al., 2000b and references therein). There are no precise data on other events during the early Palaeoproterozoic; however, recently acquired Nd model ages of Palaeoproterozoic metasedimentary granulites and igneous rocks range from 2.5 to 2.0 Ga (Kotov et al., 1995; Kovach et al., 1999; Rosen et al., 2000, 2002). The 9-12 km thick, 2.18-1.95 Ga Udokan Group, filling the intracontinental or passive margin sedimentary basin in the western part of the Aldan Province, consists predominantly of copper-bearing quartz arenites with intercalations of black shales, marine carbonates and molasse conglomerates. Emplacement of the Katugin alkaline granite at 2.01 Ga was related temporally and spatially to Udokan Group deposition. The calc-alkaline and subalkaline metavolcanic and metasedimentary granulites in the central and eastern parts of the Aldan Province were derived, based on Sm-Nd data, from both Archaean and Palaeoproterozoic sources. Soon after deposition, they underwent highgrade metamorphism (800-970~ at 7.0-10.7 kbar) together with the adjacent Archaean rocks. The 2.01-1.92 Ga granulite-facies metamorphism was broadly coeval with emplacement of a 2.04-2.01 Ga tonalite-trondhjemite complex and 1.99-1.90 Ga granitoids of various compositions. Later crust-forming episodes occurred after east-west compression as well as thrusting associated with intrusions of 1.92 Ga mafic-ultramafic dykes, 1.90-1.80 Ga A-granites and 1.87-1.77 Ga granulite-facies metamorphism. The final

Opposite: Fig. 3.9-4. The Palaeoproterozoic mobile belts in the Siberian craton (modified after Rosen et al., 1994). Palaeoproterozoic belts, tectonic zones and massifs: Ae, Aekit; An, Angara; CA, Central Aladanian; D, Daldyn; Dt, Dzheltulak; Dz, Dzugdzur gabbro-anorthosite; EA, East Aldanian (Uchur); H, Hapchan; S, Sutam; U, Udokan sedimentary basin.

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evolutionary stage saw 1.74-1.70 Ga within-plate magmatism (rapakivi-type batholiths, K-rich ash-flows, alkaline granites) on the eastern flank of the Aldan Province. In the Stanovoy Province, directly south of the Aldan Province, juvenile Palaeoproterozoic volcano-plutonic rocks are poorly known. The 1.74-1.70 Ga Dzhugdzur gabbroanorthosite massif is the only exception (Sukhanov and Zhuravlev, 1989; Neymark et al., 1992). Archaean granite-greenstone units contain fragmented linear belts and blocks of Archaean rocks, which bear the imprints of Palaeoproterozoic high-grade metamorphism at 2.2-2.0 Ga and 1.98-1.84 Ga (Mints, 2000b and references therein). The latter event preceded westwards-directed crustal shortening. The Palaeoproterozoic evolution terminated after 1.7 Ga with northwards overthrusting of the Stanovoy Province and its juxtaposition with the Aldan Province. The peak granulite-facies conditions (up to 1000-1100~ and 9.5-10.0 kbar; Karsakov, 1978), characteristic for the deepest crustal section (seen in outcrop within the Sutam unit, situated close to the boundary between the Aldan and Stanovoy Provinces), may have occurred at 1.98-1.84 Ga or later, synchronously with emplacement of the Dzhugdzhur gabbro-anorthosite massif and immediately before AldanStanovoy juxtaposition. Sm-Nd studies indicate that all or at least most of the Palaeoproterozoic juvenile magmatic rocks examined in the Siberian craton, including the gabbro-anorthosites, crystallised from melts strongly contaminated by Archaean crustal material (Kotov et al., 1995; Kovach et al., 1999; Rosen et al., 2000, 2002).

Geological Evolution of Other Continental Areas The above discussion of the assemblages in the Laurentian and Siberian cratons illustrates the main lines of crustal evolution of, and most important events in, the Palaeoproterozoic. High-grade complexes formed in the early Palaeoproterozoic are also known from India and Antarctica (Raith et al., 1990, 1999; Harley, 1998; Asami et al., 2002). In contrast, both low-grade volcano-sedimentary and high-grade belts of 1.95-1.90 Ga and younger ages are distributed sporadically in various continental areas of Asia, Australia, Africa and South America (e.g., de Almeida et al., 2000; Martin et al., 2000; Teixeira et al., 2000; Ernst and Buchan, 2002a; Wei, 2002). The period from 2.44 to about 2.0 Ga was a quiescent one not only in Laurentia but also in African cratons (e.g., Eriksson et al., 1999) and possibly worldwide.

Granulite-Gneiss Belts: Geodynamic Interpretation It is seen from the above review that the granulite-gneiss belts, besides their metamorphic grades, have a number of specific features, which are not shared by the low-grade volcano-sedimentary and volcano-plutonic belts. (1) The high-grade metamorphism is commonly predated or accompanied by crust-contaminated gabbro-anorthosite intrusions, high-temperature, "dry", within-plate granites and emplacement of enderbite-charnockite. (2) Lower parts of the high-grade sequences are usually formed by rift-related volcanics and meta-arenites derived from the Archaean basement. In contrast, the bulk of the upper

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metasediments ("khondalites" or "metagraywackes") was derived from juvenile Palaeoproterozoic rocks, whose provenance remains unknown or poorly constrained. The ages of detrital zircons and Sm-Nd data point to the existence of abundant juvenile, crustcontaminated felsic rocks of early to mid-Palaeoproterozoic ages, which are not only unknown in the present vicinity of the granulite-gneiss belts, but are rare in distant areas too. (3) The high-grade assemblages form thrust-nappe ensembles; inverted metamorphic zoning caused by heating from overthrusted hot tectonic slices can be observed in many para-autochthonous complexes. (4) The total thickness of the crustal sections that were affected by a single high-grade event may reach 50 kilometres. (5) Sediments in the uppermost part of the sequences were metamorphosed almost immediately after deposition and before transformation of the sedimentary basin into a thrust-nappe and fold belt. This suggests that intensive heat influx and related high-grade metamorphism affected the lower crust and then expanded into the volcano-sedimentary fill of the depositional basins. Final sedimentation, high-grade metamorphism and thrusting lasted some tens or at most, less than 50 My. Ideas on the nature and tectonic and geodynamic significance of granulite-gneiss belts remain controversial. In many studies, granulite-gneiss belts have been interpreted as sutures or collisional orogens (e.g., section 3.8). The thermal modelling of metamorphic evolution within collisional orogens calls for significant heating of the thickened crust up to granulite-facies conditions (England and Thompson, 1986; Thompson and Ridley, 1987). On the other hand, the features of granulite-gneiss belts discussed here are in obvious contradiction with such a geodynamic interpretation. Harley (1992) argued that collisionrelated homogeneous lithospheric thickening can not explain the very high temperatures typical of most granulites for reasonable limitations of critical model parameters, such as thermal conductivity, crustal heat productivity and basal heat flux. Regional granulite metamorphism requires input of more heat than that available from a thickened crust-lithosphere system. The rather slow cooling deduced for many granulite areas requires that the heat source must be external to the crust undergoing metamorphism; in most regional granulitefacies assemblages the high-grade metamorphism and magmatism are probably both consequences of lithospheric-scale thermal processes (Harley, 1992). Harley's arguments are especially relevant in the light of the inferred enormous thickness of granulite-facies crustal sections. Another key factor in the formation of granulite-facies terranes is low a H20 conditions resulting in relatively anhydrous ("dry") mineral assemblages, not only in the metasedimentary granulites but in enderbite-charnockite magmas also (Touret and Hartel, 1990). An increased water activity, which is characteristic for supra-subduction settings, causes partial melting and temperature stabilisation at amphibolite-facies conditions. Regional desiccation of crustal rocks is unlikely in the supra-subduction environment, but may be linked to back-arc extension. Sandiford (1989b) and Harley (1989) considered the possibility of granulite-facies metamorphism during continental rifting and suggested that granulites are forming today beneath the North American Basin and Range province. Similarly, the Early Cretaceous mafic and intermediate granulites of the Fiordland terrane (New Zealand) associated with the mid-Palaeozoic Black Giant anorthosite pluton (Gibson and

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Ireland, 1995, 1999) are localised in a back-arc extensional setting (Bradshaw, 1989). The observed abundance of rift-type assemblages in the lower portions of the Palaeoproterozoic granulitic sequences fits with the inferred rift- or back-arc-related origin of modem and young granulite assemblages. In summary, thus, the granulite-facies metamorphism was very likely caused by heat influx of mantle provenance and occurred before collisionrelated thrusting and thickening of the crust. Recently published data on detrital zircon ages provide new insights into the origin of granulite successions. Two explanations for the common absence of the source rocks for the granulite-facies metasediments may be suggested. (1) These source rocks originated close to rapidly subsiding basins that were entirely eroded, following which the erosion products were transported to adjacent basins and, soon after deposition, underwent highgrade metamorphism. (2) The metasediments may, at least in part, have been ash-flow deposits, which filled extensive calderas and volcano-tectonic depressions. Although the first explanation seems appropriate for back-arc basins, it appears to be valid in some cases only, as most of the Palaeoproterozoic granulite belts are situated far from arc-type igneous assemblages. Moreover, in some areas with a number of extensive granulite-gneiss belts (e.g., in Siberia, India) the requisite volcano-sedimentary or volcano-plutonic belts are rare or absent. The second suggestion seems more attractive because many features of crustal state, geodynamic setting and lithology are equally characteristic for both granulite protoliths and ash-flow deposits. Most important among those features are: (1) close link with back-arc and rift-related settings (Yarmolyuk and Kovalenko, 1991); (2) great volume and high rate of eruption, caldera collapse and filling (Smith, 1979); (3) "dry" or water-undersaturated, high-temperature magmatic conditions (up to 940~ e.g., Sutton et al., 2000); (4) correlation with granulite-facies metamorphism of mafic cumulates provided by deep crustal xenoliths (Smith et al., 1996); (5) common occurrence of ortho- and clinopyroxene, olivine and in some cases garnet crystals in the ash-flow assemblages and associated intrusives (e.g., Beddoe-Stephens and Mason, 1991; Sutton et al., 2000). Thus, we consider that granulite-gneiss belts resulted from: plume-induced heating; magmatism; emergence of riftogenic basins and volcano-tectonic depressions, their filling with rift-type sediments and juvenile but strongly contaminated lavas and ash-flow deposits; high-grade recrystallisation of the lower- and mid-crustal assemblages including the intraplate and back-arc basin-fills; and final thrusting and exhumation of high-grade assemblages caused by collision-related tectonism. The granulite-gneiss assemblages form intraplate belts of regional extent and some local inclusions within both accretionary and collisional orogens. Speculations on the Interaction of Palaeoproterozoic Plumes and Plate Tectonics

Reassessment of the nature of granulite-facies metamorphism leads to the recognition of the within-plate and plume-related origin of major granulite-gneiss belts. It seems probable that the start of plume-related evolution could be the same in both high- and low-grade mobile belts. Furthermore, commencement of Red Sea-type spreading resulted in rapid heat discharge via spreading ridges and cessation of the high-grade metamorphism in the

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adjacent continental crust. Although the real reasons for the different evolution of the belts remain unknown, it is clear that this difference played a fundamental role in crustal evolution. This new understanding will result directly in decreasing the number of orogenic belts that can be interpreted plausibly as sutures. Considering the worldwide decrease of magmatic activity from c. 2.45 to 2.2 Ga, Condie (1994b) argued that deep-seated heat generation processes in the Earth could not have been significantly lower during that period. As the ocean ridges are the predominant conduits for heat loss, it is extremely unlikely that plate tectonics should have stopped and restarted again several times (see also section 3.6). In view of this, a feasible explanation for the inferred stability of the supercontinent during the early Palaeoproterozoic and for the scarcity of subduction events along its margins, may lie in the assumption of intensive recycling of juvenile oceanic and young arc-related crust within the oceanic hemisphere (cf. de Wit's intra-oceanic model; e.g., 1998). Local post- and anorogenic events (bimodal magmatism including gabbro-anorthosite and rapakivi-granite intrusions, and high-grade metamorphism) in different places within a reborn Palaeoproterozoic supercontinent developed at various times after termination of the collisional and accretionary processes, generally after 1.7 Ga. In a certain sense, they can be interpreted as the beginning of a new Meso-Neoproterozoic evolutionary cycle. Palaeoproterozoic tectonic and crust-forming processes were instigated mainly by 2.51-2.44 Ga and 2.0-1.95 Ga mantle plumes of global significance ("superplumes") (sections 3.2 and 3.3). The plume-related fiftogenic and spreading processes within the continental areas can be attributed to "weak attempts" to disrupt the supercontinent (Mints, 1998). The fundamental change in the Earth's history at the Archaean-Palaeoproterozoic boundary was linked to the transition from Archaean "microplate tectonics" to Palaeoproterozoic "supercontinent tectonics" (or "micro-ocean tectonics" having in mind the limited size of the predominantly Red Sea-type oceans that originated within a partially disrupted supercontinent) (sections 3.4 and 3.6). The creation of the first supercontinent at the end of the Archaean (section 3.2), covering a significant part of the Earth's surface, must have played an essential role in the reorganisation of the convection cell system in the underlying mantle. From this perspective, the Palaeoproterozoic era witnessed only incipient breakup of the inferred supercontinent and revival of multi-cell convection in the mantle (see also section 3.6). On the other hand, the style of crustal evolution during the Palaeoproterozoic differs significantly not only from that in the Archaean, but also from that in the Phanerozoic. We thus postulate that Palaeoproterozoic history can be subdivided into five periods: (1) 2.51-2.44 Ga superplume activity and displacement of Fennoscandia; (2) 2.44-2.0 (2.11) Ga quiescent within-plate development complicated by local plume- and plate tectonics-related processes (see also section 3.7); (3) a 2.0-1.95 Ga superplume event (see also sections 3.2 and 3.3); (4) 1.95-1.75 (1.71) Ga combined plume- and plate tectonicsrelated evolution, resulting in the partial disruption of the continental crust, and formation of accretionary orogens along some margins of the supercontinent, and rebirth of the supercontinent entity, and (5) < 1.75 Ga post- and anorogenic magmatism and metamorphism.

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3.10.

FORMATION OF A LATE MESOPROTEROZOIC SUPERCONTINENT: THE SOUTH AFRICA-EAST ANTARCTICA CONNECTION

H.E. FRIMMEL

Introduction Major periods of continental growth have been linked with the formation of supercontinents and interpreted to reflect superplume events in the mantle (Maruyama, 1994), with periods of particularly high inferred continental growth rates at 2.7 and 1.9 Ga (Condie, 1998) (sections 3.2, 3.3, 3.6, 3.7 and 3.9). While the formation of a supercontinent at c. 1.2 Ga is not questioned, a causal link with a superplume event is doubtful, as the rate of continental crust formation then was not unusual (e.g., Condie, 2001b). The exact configuration of this Mesoproterozoic supercontinent remains controversial. Many workers have adopted the Rodinia (section 3.11) hypothesis, in which Laurentia and East Antarctica take a central position, whereas others prefer a configuration that is not dissimilar to Pangaea and is therefore referred to as Palaeopangaea (Piper, 2000). One of the major differences between the various proposed configurations lies in the position of the Kalahari and Congo cratons relative to East Antarctica (e.g., cf. Hoffman, 1991, and Dalziel et al., 2000). The assembly of this supercontinent is believed to have taken place during the "Grenvillian" orogeny, for which a time span from 1.35 to 1.0 Ga is given. In order to assess the probability of a Grenvillian link between southern Africa and East Antarctica, the tectono-thermal evolution of the intervening Grenvillian tectonic belts (Namaqua-Natal and Maud belts) is compared here, based on reliable and new U-Th-Pb single zircon age data.

The Namaqua-Natal Belt Although burial beneath Phanerozoic sedimentary rocks masks the inferred link between the Mesoproterozoic rocks in Namaqualand in the northwestern part of South Africa (Namaqua belt) and those in Kwazulu-Natal near the South African east coast (Natal belt), broad similarities in their respective metamorphic and kinematic history and geochronology, as well as the continuation of geophysical anomalies between the two belts, have been used to argue for a contiguous tectonic belt of Grenville age (Jacobs et al., 1993). The Natal belt was thrust towards the north onto the Archaean Kaapvaal craton (Matthews, 1972), whereas the Namaqua belt was thrust towards the east onto the Palaeoproterozoic Kheis belt along the western margin of the Kaapvaal craton (Humphreys and Van Bever Donker, 1987).

The Namaqua belt A series of major structural discontinuities within the Namaqua belt have been used to discriminate between several terranes (Fig. 3.10-1): Kaaien, Areachap, Kakamas and Bushmanland. The Areachap terrane is interpreted as a juvenile magmatic arc that accreted from the west or northwest onto the Kalahari craton (= Kaapvaal and Zimbabwe cratons, The Precambrian Earth: Temposand Events I~itcd by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

3.10. Late Mesoproterozoic Supercontinent

241

Fig. 3.10-1. Distribution of Palaeoproterozoic (Eburnian) and Mesoproterozoic terranes in southern Africa (modified after Thomas et al., 1993).

sutured along the Limpopo belt; see section 3.8) prior to the culmination of Mesoproterozoic orogeny in the Kakamas and Bushmanland terranes. Age constraints for the Areachap arc range from 1350 to 1280 Ma (Cornell et al., 1990), with the most reliable date being a U-Pb single zircon age of 1286 + 6 Ma (Evans et al., 2002). An older age of 1377 -+-5 Ma has been reported for a tonalitic gneiss beneath the volcano-sedimentary Sinclair Group in southern Namibia (Hoal and Heaman, 1995). Most of the high quality age data come from the Bushmanland terrane. There, a Palaeoproterozoic 2.0-1.8 Ga basement (Robb et al., 1999), floors a succession of high-grade metasedimentary and intercalated mafic and felsic metavolcanic rocks. This basement is comparable to a c. 2.0-1.8 Ga low-grade Andean-type magmatic arc (Reid, 1997) that is preserved as the Richtersveld terrane, which forms a basement wedge between the Kakamas and the Bushmanland terranes (Fig. 3.10-1). The northern contact of the Richtersveld terrane is a wrench-thrust, but the southern contact, although shown as a thrust on published maps, must be a normal fault, because it has a northwards dip with low-grade metamorphic rocks overlying high-grade ones. Previously, the supracrustal rocks in the Bushmanland terrane, which comprise

242

Chapter 3: Tectonism and Mantle Plumes Through Time

metapelite, metapsammite, quartzite, marble, calc-silicate, iron-formation, and bimodal, acid-dominated metavolcanic rocks, and which host vast stratiform exhalative massive base metal sulphide deposits (Moore et al., 1990), were believed to be roughly 2.0-1.6 Ga (Reid et al., 1997). A recent single zircon U-Th-Pb isotope study (Raith et al., 2003) revealed, however, that at least parts of the supracrustal successions are considerably younger. Detrital zircon grains in some of the quartzo-feldspathic metasedimentary rocks were derived from 1.25-1.20 Ga source rocks, whereas younger metapelite contains detrital zircon grains as young as 1136 + 21 Ma. Westwards in the Namaqua belt, granitic rocks become more abundant and the metamorphic grade increases from upper amphibolite to granulite facies. Whereas Willner (1993) arrived at a clockwise p-T path for the low-pressure amphibolite facies region in the northeastern part of the Bushmanland terrane, Waters (1989) established an anticlockwise p-T path for the low-pressure granulites in the central part of the terrane. From recent ion microprobe data on complex zircon grains, it appears that the thermal peak was reached more or less at the same time throughout the entire terrane. U-Pb single zircon ages of 1032 + 12 Ma (Robb et al., 1999) and 1027 • 14 Ma (D. Cornell, unpubl, data) in the northern part of the terrane are within error of the age of 1022 -+-5 Ma obtained for the southern Bushmanland terrane (Raith et al., 2003). Extensive recrystallisation masked previous mineral assemblages and textures during this high-grade metamorphism, which reached temperatures as high as 870~ at 5 kbar (Nowicki et al., 1995). However, an older metamorphic event, probably associated with an early phase of pre-tectonic granite plutonism (Little Namaqualand Suite), dated at 1212 i 11 and 1199 + 12 Ma (Robb et al., 1999), has been suspected by many workers, and has recently found support from U-Pb ages of 1187 -4- 15 Ma obtained on metamorphic zircon rims around 1200-1250 Ga detrital zircon cores (Raith et al., 2003). The exact nature of this metamorphism remains unclear, but it was probably accompanied by isoclinal folding and thrusting, recorded by locally preserved intrafolial fold relics that predate the penetrative deformation associated with the younger metamorphic event. The lack of Mesoproterozoic ages older than 1250 Ma, and the widespread emplacement of granitoids around 1200 Ma, followed or accompanied by high-grade metamorphism and probably intense deformation, suggest a major crust-forming event between about 1250 and 1200 Ma. Formation of a robust magmatic arc on a Palaeoproterozoic basement with subsequent accretion of that arc against the older arc of the Areachap terrane is indicated. Deposition of some of the supracrustal rocks has been bracketed between c. 1250 and 1187 + 15 Ma (Raith et al., 2003), from which a sediment source within the arc is inferred. This period overlaps with the time of deposition of the volcano-sedimentary Sinclair Group around 1216 Ma (Hoal and Heaman, 1995). Virtually undeformed and only very low-grade metamorphosed bimodal volcanic and coarse-grained siliciclastic rocks, that rest non-conformably above 1.8 Ga metamorphic rocks of the Kheis belt, are found in the Koras Group. Its relationship to the Sinclair Group and inferred sedimentation in the Bushmanland terrane is unclear. An age of 1171 -+- 7 Ma obtained for a quartz porphyry near the base of the Koras Group sets not only an upper limit on the age of sedimentation

3.10. Late Mesoproterozoic Supercontinent

243

but also a minimum age constraint on arc accretion, deformation and metamorphism of the Areachap arc (Gutzmer et al., 2000). In the Bushmanland terrane, magmatic arc formation was followed by the intrusion of mafic bodies of variable scale, ranging from metagabbroic and noritic plutons to amphibolitic or two-pyroxene granulitic dykes and sills, one of which yielded a Pb-Pb age of 1168 4-9 Ma (Robb et al., 1999). Arc formation was also succeeded by large volumes of equigranular to megacrystic granitic augen gneiss that is locally rich in biotite (> 10 vol.%), fine-grained and contains variable proportions of decimetre-sized alkali feldspar megacrysts. Contacts with adjacent semipelitic gneiss typically are gradational and marked by a gradual decrease in alkali feldspar augen. These relations are best explained as the product of partial decompression melting of biotite-rich assemblages with subsequent K-diffusion in the lower crust. One example of this gneiss yielded U-Pb single zircon ages of 1109 4- 7 and 1117 -4- 24 Ma (de Beer et al., 2002; Raith et al., 2003). Most of the metabasites associated with the augen gneiss are tholeiitic in composition, thus supporting their derivation from a mantle peridotite source by fractional crystallisation, as can be expected for a continental rift or back-arc basin (Reid et al., 1987; Moore, 1989; Raith and Meisel, 2001). An overall extensional regime at c. 1.1 Ga is supported further by the finding that some of the supracrustal rocks in the Bushmanland terrane were deposited some time between 1136 4- 21 Ma and 1065 + 2 Ma (Raith et al., 2003). Syn- to post-tectonic suites occur as massive, K-feldspar megacrystic granitoids (Spektakel Suite) and complex, in places highly cupriferous, anorthosite, diorite, leuconorite, norite and hyperstenite (Koperberg Suite). SHRIMP U-Pb zircon ages of 1064-+-31 and 1057 4- 8 Ma (Robb et al., 1999) as well as of c. 1030 Ma (Duchesne et al., 2001) have been reported for these suites. Associated charnockite yielded a U-Pb zircon age of 1063 -+- 18 Ma (H.E. Frimmel, unpubl, data). High-grade metamorphism probably began around that time and lasted until c. 1020 Ma. Considering their large volume, the rising melts most likely provided the heat transfer mechanism for high-temperature metamorphism to occur at relatively shallow crustal depths of not more than 15 km; this has been explained by regional mafic underplating (Waters, 1989). Granulite-facies metamorphism post-dates all major intrusive suites and obliterated most tectonic fabrics (Waters, 1989). Although a series of deformation events, with polyphase isoclinal folding, formation of a penetrative foliation and later open folds, and formation of synformal and antiformal structures can be distinguished, the exact age relationships of these events remain problematic (de Beer et al., 2002). A strong gneissosity imprinted on the older granitic bodies, including the 1109 Ma augen gneiss, indicates that the main fabric-forming deformation must have been younger and was probably contemporaneous with the emplacement of undeformed to mildly deformed intrusive suites around 1.07-1.05 Ga. Notwithstanding evidence for local compression (Kisters et al., 1996) an overall extensional regime is suggested for this thermal event, with down-faulting of the cooler Eburnian basement in the Richtersveld terrane relative to the Bushmanland terrane. In the latter, large volumes of rising felsic and mafic melts caused not only a thermal high but also contrasting p-T paths for the intruded supracrustal rocks, depending

244

Chapter 3: Tectonism and Mantle Plumes Through Time

on their proximity to the intrusive bodies (clockwise in the amphibolite facies and anticlockwise in the granulite facies rocks) (cf. section 3.9). The development of large-scale dextral wrench shear zones (Fig. 3.10-1) outlasted the metamorphic peak but still took place under high-grade conditions, and mark the final juxtaposing of the Mesoproterozoic crust with the Kaapvaal craton. Following the thermal peak around 1020 Ma, exhumation of the Namaqua Belt was fairly rapid. The high-grade metamorphic rocks cooled below about 320~ at 1006 -+- 4 Ma, as indicated by Ar-Ar biotite data (Frimmel and Frank, 1998). Renewed heating of the Namaqua crust is indicated by alkaline shallow intrusive activity, dated at 833 + 2 Ma, which heralds crustal thinning that eventually led to continental break-up at the beginning of the Pan-African tectonic cycle around 750 Ma (Frimmel et al., 2001). The Natal belt

In contrast to the Namaqua belt, no older basement rocks have been recognised in the Natal belt, which is therefore interpreted as a juvenile orogen (Eglington et al., 1989). Three tectono-stratigraphic units, the Tugela, Mzumbe and Margate terranes are distinguished (Thomas, 1989; Fig. 3.10-1). The Tugela terrane, adjacent to the Kaapvaal craton, comprises predominantly mafic metavolcanic and intrusive rocks, interpreted as a former ocean arc complex. Arc plutonism has been dated at 1209 -+- 5 Ma, and subsequent crustal thickening and tectonic burial are older than about 1180 Ma, the age of late- to post-tectonic mafic intrusive rocks and metamorphism (Johnston et al., 2001). Obduction (see also section 3.7) of the terrane onto the Kaapvaal craton was accompanied by high-pressure granulite facies metamorphism (Johnston et al., 2002) and occurred at around 1155 4- 1 Ma, the age of syn-tectonic granite and quartz monzonite sheets and dykes. Detrital zircon grains yielded a maximum age of 1276 • 10 Ma, which is the oldest age reported to date from the entire Natal belt (Johnston et al., 2001). These authors inferred from this latter age that the sediments were derived from within the Tugela terrane, which at that stage must have been far removed from the Kaapvaal craton. A similar history has also been established for the Mzumbe terrane. The oldest rocks recognised comprise 1235 + 9 Ma arc-related, felsic to mafic metavolcanic and minor metasedimentary supracrustal gneisses; a tonalite-trondhjemite suite of calc-alkaline I-type character intruded the supracrustal rocks at 1207 + 10 Ma (Thomas et al., 1999). Northeastdirected thrusting of the three terranes onto the Kaapvaal craton led to an inverse metamorphic stacking with granulite-facies rocks occurring in the structurally highest Margate terrane in the south. In both the Mzumbe and Margate terranes, an intervening extensional period, marked by the emplacement of a mafic dyke swarm, was followed by renewed plutonism (Thomas et al., 1999). Leucogranite intruded around 1100 Ma and was followed by late-tectonic 1065 -+- 15 Ma rapakivi granite and charnockite, accompanied by low to moderate pressure, high-temperature metamorphism. Minor intrusive activity continued until about 1025 Ma in an overall sinistral oblique strike-slip regime that outlasted the thermal peak until about 980 Ma, as derived from Ar-Ar data (Jacobs et al., 1993, 1997).

3.10. Late Mesoproterozoic Supercontinent

245

The Maud Belt

The widespread occurrence of 1.3-0.9 Ga isotopic ages throughout East Antarctica led to the inference of a continuous Grenville-age mobile belt, linked with the Grenville belt of North America to form the main suture in Rodinia configurations (e.g., Hoffman, 1991). However, subsequent age determinations around 0.5 Ga in this part of Antarctica, at least partially related to a major Pan-African tectonic overprint (Shiraishi et al., 1994; Dirks and Wilson, 1995; Fitzsimons et al., 1997; Jacobs et al., 1998), make a continuous Grenvillian Rodinia suture around East Antarctica questionable (Fitzsimons, 2000). The Maud belt in western and central Dronning Maud Land (Fig. 3.10-2) exemplifies the two-stage thermal history that is typical of the Mesoproterozoic to Early Cambrian high-grade metamorphic belts around East Antarctica. It consists of polyphase deformed upper amphibolite to granulite facies metamorphic rocks and a variety of intrusive rocks. The belt is separated by large glaciers from the Grunehogna Province to the northwest, comprising Archaean granite overlain by the Mesoproterozoic volcano-sedimentary Ritscherflya Supergroup (Wolmarans and Kent, 1982; Moyes et al., 1995). This province is generally believed to represent the easternmost portion of the Kaapvaal craton, detached during the breakup of Gondwana (e.g., Groenewald et al., 1995). The Ritscherflya Supergroup comprises a regressive sedimentary sequence of shallow marine to braided river deposits, with a proximal source in the southwest (Ferreira, 1988). The sedimentary strata are capped by basaltic to andesitic lava flows and minor volcaniclastic deposits. All of these deposits were intruded by thick mafic and ultramafic sills (Borgmassivet Intrusive Suite) of continental tholeiitic geochemical affinity (Krynauw et al., 1991); one of these sills yielded a U-Pb zircon age of 1107 Ma (M. Knoper, unpubl. data). The extent of deformation and metamorphism is generally very low but increases to the southeast near the Maud belt. Throughout the Maud belt, the oldest rocks identified are banded, probably metavolcanic paragneisses, with U-Pb single zircon ages between 1160 and 1140 Ma (Arndt et al., 1991; Harris et al., 1995; Jacobs et al., 1999; Paulsson and Austrheim, 2001; Board, 2002). The oldest intrusive rocks, mainly granitoids, are of similar age, ranging from 1140 to 1130 Ma (Jackson, 1999; Jacobs et al., 1999; Board, 2002). Based on their overall calcalkaline composition, these rocks have been interpreted as a magmatic island arc, whereas a back-arc basin is envisaged for the structurally overlying metasedimentary rocks, which comprise metapelitic migmatite, paragneiss, calc-silicate and minor marble, all of which are intruded by mafic dykes (Groenewald et al., 1995). U-Pb single zircon ages between 1100 and 1060 Ma have been obtained for various granitic gneisses and pegmatite in the northeastern half of the Maud belt (Harris et al., 1995). A megacrystic augen gneiss with a strike length of at least 300 km locally displays alkali feldspar megacrysts with resorption rims, suggesting that they are older than their matrix. SHRIMP U-Pb ages of 1127 -1- 12 and 1061 4- 14 Ma were obtained for this rock (Harris et al., 1995). The older age possibly records the event associated with crystallisation of the alkali feldspar megacrysts, whereas the younger one could reflect the time of matrix recrystallisation during metamorphism. The latter overlaps with a zir-

246

Chapter 3: Tectonism and Mantle Plumes Through Time

Fig. 3.10-2. The Maud belt, consisting of outcrops at Heimefrond]ella (HF), Kirwanveggen (KV), H.U. Sverdrupfjella (SF), Gjelsvikfjella (GF) and Central Dronning Maud Land (CDML), within a Gondwana framework (modified after Kr6ner, 2001).

con age of 1073 +15 -9 Ma from an orthogneiss in Heimefrontfjella (Arndt et al., 1991). In Kirwanveggen (Fig. 3.10-2), bimodal magmatism might be indicated by the close association of 1074-+- 10 Ma megacrystic granitic to granodioritic gneiss and metabasites (Jackson, 1999). High-grade metamorphism around 1.1 Ga is documented by a leuco-

3.10. Late Mesoproterozoic Supercontinent

247

some age of 1098 -+- 5 Ma in Kirwanveggen (Harris et al., 1995) and by metamorphic zircon growth at 1104 + 5 Ma in Heimefrontfjella (Arndt et al., 1991). Younger metamorphic zircon ages in Kirwanveggen are 1050-+- 10 Ma and 1031 4-6 Ma (Jackson, 1999). Similarly, metamorphic zircon overgrowth around 1035 Ma has also been recorded in H.U. Sverdrupfjella (Board, 2002). Whether these ages reflect discrete metamorphic pulses or a continuum of metamorphic recrystallisation from 1100 to 1030 Ma, or during parts of this time span, remains unclear. The penetrative fabric throughout most parts of the Maud belt is the product of top-to-northwest thrusting, for which a Grenvillian age has been inferred in the past (Grantham et al., 1995; Jackson, 1999). More recently, petrological analysis and U-Pb dating of syn-tectonic monazite revealed that this intense deformation is of Pan-African age, i.e., 530 4- 8 Ma, and was accompanied by upper amphibolite facies metamorphism following decompression from eclogite facies conditions (Board, 2002). Eclogitic mineral assemblages are found in pre-kinematic boudins, and it remains uncertain whether they are Grenvillian or early Pan-African in age. Obliterated in most areas by the intense PanAfrican re-foliation, an older fabric that pre-dates a cross-cutting granitic dyke with an age of 1011 4- 8 Ma, has been documented from the central part of the belt (Jackson, 1999). Rotation of the older fabric elements by the Pan-African deformation makes the reconstruction of the Grenvillian kinematic history difficult, but in Kirwanveggen, Jackson (1999) identified passive shear folds with a southeast-plunging kinematic vector for this earlier phase of deformation. The degree of Pan-African metamorphic overprinting of Mesoproterozoic crust increases towards the northeast, and this resulted in high-grade metamorphism and resetting of most isotope systems to c. 530-540 Ma, partial melting of the older crust, and subsequent formation of post-orogenic granite and orthopyroxene-rich syenite around 520-500 Ma in central Dronning Maud Land (Jacobs et al., 2002). No evidence of juvenile Pan-African crust has been identified in the Maud belt. Age of arc volcanism and subsequent bimodal magmatism The presence of volcanic and volcaniclastic deposits in the Grunehogna Province that escaped the intense polyphase deformation and metamorphism in the adjacent Maud belt offers an opportunity to further constrain the time and nature of crust formation in the Maud belt. Zircon grains were therefore separated from two tuff samples from the upper Ahlmanryggen Group (lower Ritscherflya Supergroup) and analysed for their U-Th-Pb isotopic compositions (for analytical details see Frimmel et al., 2001; isotope data available upon request from author). Two of four zircon grains from a tuff (sample B.VE1) are close to concordant (Fig. 3.10-3a), with the other two showing some discordance. All four grains define an upper intercept of 1130 + 7 Ma (MSWD - 0.63). This age is within error of the 2~176 ages of the two least discordant grains and is therefore considered the best estimate of the age of zircon crystallisation. Similarly, four zircon grains from another tuff (sample B.BP3) are almost concordant, yielding an upper intercept age of 1131 4- 7 Ma on a concordia diagram (Fig. 3.10-3b). Again, this age is indistinguishable from the mean 2~176 age and is interpreted as representative of the age of zircon crystallisation, and thus also of fel-

248

(a)

Chapter 3: Tectonism and Mantle Plumes Through Time

0.205

B.VE1 0.195 l140

J ~} 0.185

.Q

1060xr. /

a,, SlD

/

0.175

O

.Y 0.165

/

/

Intercepts at 1130 + 7 Ma and 228 + 250 Ma MSWD = 0.63

0.155 t_ 1.65

1.75

1.85

1.95

2~

2.05

2.15

U

(b) B

m

B.BP3

0.197

1150

0.193

1140

.Q a..

~D

0.189 Intercepts at 1131 + 7 Ma and 364 + 500 Ma MSWD = 1.7 0.185 t - z 1.96

I

I 2.00

I

I

I

I

2.04

2~

2.08

I

I 2.12

U

Fig. 3.10-3. Concordia diagrams for single zircon grains from two tuff layers, B.VEI (a) and B.BP3 (b), in the upper Ahlmanryggen Group, Ritscherflya Supergroup, western Dronning Maud Land.

3.10. Late Mesoproterozoic Supercontinent

249

sic volcanism in the hinterland of the Ahlmanryggen. This age is in good agreement with ages obtained previously for the earliest magmatic phase in the Maud belt. The Ritscherflya Supergroup is therefore interpreted as representing the fill of a basin that was positioned between an Archaean craton and a 1.13 Ga island arc to the southeast. It remains unclear, however, whether this basin was in a foreland or back-arc position. To assess the timing of bimodal magmatism that followed the formation of the 1160-1130 Ma magmatic arc, zircon grains from a metamorphosed quartz syenite and a metagabbro body from southernmost Kirwanveggen (mapped by S. Helferich, unpubl. data) were also analysed for their U-Th-Pb isotopic compositions (data available from author upon request). A concordia diagram for the zircon analyses from the metagranitoid (Fig. 3.10-4a) clearly shows the existence of two age groups. Six discordant analyses define an upper intercept age of 1079-t- 17 Ma, whereas one analysis plots close to concordia with a 2~176 age of 1024 -+- 4 Ma. The upper intercept for the six zircon grains is indistinguishable from the mean 2~176 age of 1078 -4- 5 Ma for these grains, which is interpreted as the age of granitoid emplacement. The younger age probably reflects the age of a high-grade metamorphic overprint. The zircon analyses from the metagabbro show some scatter on a concordia diagram (Fig. 3.10-4b). Ignoring the two analyses with particularly high analytical errors, a group defining a regression line with an upper intercept of 1073 + 19 Ma may be distinguished from a second one with an upper intercept of 1093 -+- 9 Ma. Both regression lines go through the origin. As the two upper intercept ages are within error of each other, the statistical significance of this distinction is questionable. However, the corresponding mean 2~176 ages of 1073 -+- 2 Ma and 1094 -t- 4 Ma, respectively, support the existence of two zircon populations. The older age is interpreted as the time of gabbro crystallisation and the younger age, which is within error of that obtained for the quartz syenite, might reflect renewed heating during granitoid emplacement.

Implications for Late Mesoproterozoic Supercontinent Formation Continuation of aeromagnetic trends from the Namaqua-Natal belt into the Maud belt and Coats Land (Gose et al., 1997; Golynsky and Jacobs, 2001) points to juxtaposition of the two belts in the late Mesoproterozoic. The contemporaneity of extensive, partly rift-related, predominantly tholeiitic 1105 Ma magmatism on the Kalahari craton (Umkondo Suite) and similar magmatism in Laurentia (e.g., Midcontinent rift) and in the Grunehogna Province (Borgmassivet Suite) has been used to infer proximity of the Kalahari/Grunehogna and Laurentian cratons at that time (Hanson et al., 1998). Dalziel et al. (2000) suggested indentation of the Kalahari craton into southern Laurentia during collisional orogeny and explained the Midcontinent rift system of Laurentia and the Ghanzi-Chobe rift of the Kalahari craton as collision-induced impactogens. The apparent contiguity of the Grenvillian tectonic belts from southern Laurentia, around the Kalahari craton into East Antarctica formed an important cornerstone in the reconstruction of the inferred Rodinia supercontinent. From the new geochronological data summarised here it becomes apparent, however, that continental growth around the Kalahari

250

Chapter 3: Tectonism and Mantle Plumes Through Time

(a) 0.182

Meta-quartz syenite

0.178t

0"174I ~ /

/ 'Y/

0.170 I--

........

..~/J"

0.166 1.68

/

1.72

1.76

1.80

1.84

2~

(b)

,

[intercepts at 1079 + 17 Ma

I

1.92

1.88

u

0.190

0.186

- Metagabbro

~0/

0.182

,~

o

0.178

_

,o4o_Jx

0.174

~" "/ /

0.170

/

",q Intercepts at 1105+ 22 Ma I and 156 + 44 0 Ma [ MSWD = 13

Intercepts at 1073 + 19 Maand-1 + 1500 Ma MSWD = 2.3

0.166 0.162 / , , 4 1.68

.

I 1.72

I

I 1.76

I

I 1.80

I 1.84

1

l 1.88

I

I 1.92

I

I 1.96

I 2.00

2o7pb/235U

Fig.3.10-4. Concordiadiagramsfor singlezircongrainsfrommetamorphosedquartzsyenite(a) and metagabbro(b) fromthe Maudbeltin southernmostKirwanveggen,westernDronningMaudLand.

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craton (including the Grunehogna Province) in Grenvillian times did not take place simultaneously but occurred in several stages over a time span of some 200 million years. The improved resolution in the geochronology of these belts reveals clearly that their formation is not related to a single supercontinent-forming collision event (Fig. 3.10-5). The first major Mesoproterozoic crust-forming event recorded in the Namaqua, Natal and Maud belts is the formation of a magmatic arc. Only in the northern Namaqua belt is this arc floored by a pre-Grenvillian (Eburnian) 2.0 Ga basement that was possibly accreted onto the Kalahari craton during the 1.8 Ga Kheis orogeny. Eburnian ages are widespread from the area between the Kalahari and Congo cratons, in West Africa and along the eastern side of the Amazonian and Rio de la Plata cratons in South America. There, the craton margins are defined by a major 2.0 Ga orogenic front (Teixeira et al., 2002). It may be speculated that the cratonic blocks of South America and Africa were already amalgamated at that time, with the opening of a Mesoproterozoic ocean more or less along the Eburnian suture (Fig. 3.10-6a). The first stage of Grenvillian continental growth of the Kalahari/Grunehogna craton affected the northwestern (in today's co-ordinates) margin of the craton. There, a 1350-1280 Ma volcanic arc accreted together with Eburnian basement around 1280 Ma. This evolution corresponds in time with that recorded in the southwestern part of the Grenville orogen in Texas (Mosher, 1998), but it is younger than the 1370 Ma Kibaran orogeny in East Africa (Tack, 2002). The second stage of continental growth (Fig. 3.10-6b), between 1250 and 1200 Ma, involved the formation of a second magmatic arc in the Namaqua belt and the first and only arc recorded in the Natal belt. Continuity from the Namaqua to the Natal belt is now evidenced by geochronological data that indicate a major arc-continent collision event in both belts between 1190 and 1170 Ma. This event led to continental growth along the southwestern and southern margin of the craton and caused not only widespread metamorphism and contractional deformation but also a hiatus in sedimentation. A link between the southwestern Grenville orogen and the Namaqua-Natal belt, as suggested by Dalziel et al. (2000), is doubtful. Apart from palaeomagnetic data that indicate separation of Laurentia and Kalahari by as much as 30-t- 14 ~ latitude at 1105 Ma (Powell et al., 2001), no evidence of continental collision exists from the Namaqua-Natal belt for the time of continental collision in the Llano-West Texas region (1150-1120 Ma). The third stage of continental growth around the Kalahari/Grunehogna craton affected its eastern margin, with the accretion of a 1160-1130 Ma magmatic arc in the Maud belt (Fig. 3.10-6c). The kinematic history of that belt is only poorly constrained, because of the intense Pan-African tectonic overprint. Deposition of the metasedimentary rocks in the Maud belt was roughly contemporaneous with arc formation and deposition of the volcanosedimentary successions of the Ritschersflya Supergroup, for which a back-arc setting is favoured. The isolated eclogite boudins in the supracrustal succession of the Maud belt may have formed during arc accretion but this still awaits confirmation or refutation by radiometric age data. In a Gondwana configuration, the expected continuation of the Maud Belt should be found in the Mozambique Belt of East Africa. As pointed out by KrOner (2001), evidence for magmatic arc formation and/or accretion between 1.25 and 1.1 Ga in the Mozambique belt and in Madagascar is scarce. Only a few zircon ages point to granite

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Fig. 3.10-5. Comparative summary of tectonothermal and sedimentary episodes in the Namaqua, Natal and Maud belts (from data discussed in the text) and the southwestern part of the Grenville orogen in southern Laurentia (after Mosher, 1998).

Opposite: Fig. 3.10-6. Three stages of continental growth (dark grey) of the Kalahari/Grunehogna craton (K) at around 1300 Ma (a), 1200 Ma (b) and 1130 Ma (c) within a Rodinia framework (similar to the reconstruction by Dalziel et al., 2000). Approximate position of Eburnian magmatic arcs/orogenic belts in South America and southern Africa (cross-hatched) is shown to illustrate a possible link prior to Grenvillian times. Kibaran belt shown in transverse hatching. All other cratons are undifferentiated (light grey). Some modem continent boundaries are shown for clarity. Although a Rodinia configuration is shown, note that the three-stage continental growth evolution of the Kalahari/Grunehogna craton would be compatible also with a Palaeopangaea model. F, Falkland Plateau; RP, Rio de la Plata craton.

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emplacement between 1.1 and 1.0 Ga in the Mozambique belt and zircon xenocryst ages for that period have been obtained from early Pan-African gabbro and granite plutons in the Antananarivo tectonic block in southern Madagascar (KrOner, 2001). While arc-continent collision took place on the (south-) eastern side of the Kalahari craton, crustal thinning along the western edge of the craton led to intra-continental sediment deposition, as recorded by the ~< 1171 Ma Koras and the 1105 Ma Ghanzi-Chobe (Modie, 1996) rifts in southwestern Africa, and to decompression melting in the lower crust and upper mantle between c. 1150 and 1070 Ma. Whether these rifts represent backarc basins in response to the 1250-1200 Ma arc on the southwestern side, or are a far-field effect of 1160-1130 Ma arc accretion on the eastern side of the craton, or are unrelated altogether to the tectonic processes around the craton margin, remains unclear. The latter possibility is favoured in view of the significant age differences between the accretion events around the craton and this continental rifting phase. Hanson et al. (1998) have already suggested that the Ghanzi-Chobe rift is the western part of a much larger, regionally developed igneous province (Umkondo province) and thus reflects a craton-wide igneous event. All around the craton, from the Namaqua to the Maud belt, this period was followed by low pressure, high temperature metamorphism and bimodal magmatism in the lower crust, that lasted until 1020 Ma. The anti-clockwise p-T path established for some of the granulite facies metamorphism (Waters, 1989) might reflect thickening of the crust due to the emplacement of large volumes of late-tectonic granitoids, with corresponding clockwise p-T paths experienced by the mid-crustal rocks that sank to greater depths due to gravitational redistribution, similar to the model of Gerya et al. (2000). In spite of the difference in the timing of late Mesoproterozoic accretion and collision, extension-related basin development in the upper crust and high-temperature, low-pressure metamorphism and partial melting of the lower crust and upper mantle took place more or less simultaneously throughout southern Africa, the southern -part of the Grenville belt in Laurentia (e.g., Llano uplift) and Western Dronning Maud Land. If this extension were related to crustal thinning consequential upon lithospheric delamination or orogenic collapse, the delay from the orogenic peak should be similar in all belts. This is not the case. Thus a common cause is sought that is independent of the preceding tectonic history of each area, i.e., independent of crustal processes, such as a thermal event in the mantle (see also sections 3.2-3.4) as previously proposed to explain the Umkondo magmatism (Hanson et al., 1998). If decompression melting of the upper mantle and lower crust were decoupled from crustal plate movements, it could explain the widespread occurrence of similar and coeval magmatism and high-grade metamorphism in a variety of different types of continental crust over a large area. Development of one or more plume heads below a thick, stable, consolidated, relatively small craton would lead to lateral escape of heat along the base of the craton and elevated heat flow along the craton margin, thus promoting the type of magmatism and metamorphism observed around the Kalahari craton between 1.1 and 1.0 Ga. Did the speculated plume(s) beneath the inferred Mesoproterozoic supercontinent eventually coalesce to form a superplume that triggered the break-up of the supercontinent? If the term superplume is used in the sense of a number of mantle plumes rising to the base

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of the lithosphere over a short period on the order of 50 million years (Condie, 2001b) (section 3.2), the answer must be negative, as supercontinent breakup along the Mesoproterozoic sutures around the Kalahari craton began only around 750 Ma (Frimmel et al., 2001). If the inferred late Mesoproterozoic thermal mantle anomaly was of longer duration, it may have played the triggering role in early Neoproterozoic continental breakup.

3.11.

A MECHANISM FOR EXPLAINING RAPID CONTINENTAL MOTION IN THE LATE NEOPROTEROZOIC

J.G. MEERT AND E. TAMRAT Introduction

The notion that large continental plates (> 2 x 107 km 2) with thick tectospheres might undergo extremely rapid motion (> 20 cm yr -1) via "normal" plate driving mechanisms is viewed with scepticism. Available constraints on post-Mesozoic plate motion (for which we also have oceanic floor records) indicate that the best-documented case for rapid motion of a continent with significant size is the c. 20 cm yr-l northwards migration of India in late Mesozoic to early Cenozoic time (Klootwijk et al., 1992). However, the Indian continent is a relatively small piece of continental crust (an order of magnitude smaller than Laurentia) and its rapid motion is often explained by invoking both a warmer mantle beneath India and the increased pull of the ancient Tethyan oceanic slab (Fig. 3.11 - 1; see also Forsyth and Uyeda, 1975; Gordon et al., 1979). In fact, the analysis of plate-driving forces by Forsyth and Uyeda (1975) indicated that the motion of large continental plates would be "slowed" by excess asthenospheric drag at the base of the plate (Fig. 3.11-1; see also Meert et al., 1993; Gurnis and Torsvik, 1994). Recently, Becker and O'Connell (2001)discussed the problems faced by geodynamicists trying to model the relative contributions of each of these forces. Although there must exist a theoretical upper "Plate Tectonic Speed Limit", it is not explicitly stated in the literature. To avoid this problem some authors simply placed a limit on the speeds of continental plates during the development of their palaeogeographic models (e.g., Scotese et al., 1999) or advocate alternative mechanisms to account for the apparent high velocities (Kirschvink et al., 1997; Evans, 1998). While the absolute limits for the rate of plate motion are not explicitly stated, one can imagine that they are ultimately related to the thermal regime underlying the plates. In fact, variations in the thermal regime of the Earth (whether due to plumes, mantle insulation or subduction) (see also sections 3.2 and 3.5-3.7) have all been used in geodynamic models to describe the enhancement or inhibition of plate velocities (Gurnis, 1988, 1990; Gurnis and Torsvik, 1994). Gordon et al. (1979) analysed the pre-Tertiary motions of continents using palaeomagnetic data and concluded that Palaeozoic continental plate motions were significantly faster than those observed today. All pre-Mesozoic plate velocity estimates are regarded as minima since we cannot, without correlative seafloor anomaly data, account for longitudinal motion using palaeomagnetism. Gordon et al. (1979) documented motions on the order The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altcrmann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Fig. 3.11-1. "Normal" plate driving/inhibiting forces. The figure is modified from Forsyth and Uyeda (1975) and Becker and O'Connell (2001). Positive (--driving) forces are in italics and negative (-- inhibiting) forces are underlined.

of c. 5-6 cm yr-I for Laurentia, Gondwana and Eurasia and argued that the presently observed continental velocities should not be viewed as limits to the maximum speeds of large continental plates. Ullrich and Van der Voo ( 1981) also noted rapid pulses of latitudinal velocities for several continents, but their analysis was hindered by a lack of well-constrained ages/poles for the continents in the Proterozoic. Nevertheless, they suggested that plate motions in the past might have included pulses of rapid motion. Subsequently, Meert et al. (1993) provided evidence that both Laurentia and Gondwana underwent phases of very rapid plate motion during the late Neoproterozoic and middle Palaeozoic, respectively. Later, Meert and Van der Voo (1997) noted that Gondwana's late Neoproterozoic to early Palaeozoic motion approached minimum velocities of up to 24 cm yr - l , and Kirschvink et al. (1997) noted that the apparent polar wander paths (APWPs) from several large continents showed nearly 90 degrees of motion over a 15 Ma time span (c. 60 cm yr -l). A number of explanations were proposed to explain this rapid motion including a warmer mantle beneath the Neoproterozoic supercontinent (Gurnis and Torsvik, 1994) (see also section 3.10), true polar wander (TPW; Evans, 1998), inertial interchange true polar wander (IITPW; Kirschvink et al., 1997) and a combination of both TPW and warmer mantle conditions (Meert, 1999). None of the explanations has been wholly satisfying and both the observations and explanations for these fast plate motions are hotly debated. For example, proponents of IITPW argue that the palaeomagnetic data support rapid motion of nearly 90 degrees during the Tommotian-Toyonian interval (523-508 Ma; Kirschvink et al., 1997; Evans et al., 1998) whilst opponents of the idea note the discordance in length of appar-

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ent polar wander paths and the non-synchroneity of the observed motion (Torsvik et al., 1998; Meert, 1999; Torsvik and RehnstrCm, 2001). Rapid plate motion (~> 10 cm yr -1) is observed elsewhere in the Proterozoic record (late Mesoproterozoic; Meert and Torsvik, in review), but is best documented in the late Neoproterozoic interval. This section examines the first-order observations supporting the rapid plate motion in the Neoproterozoic and the subsequent evolution of the continents involved in the rapid motion. The argument is made that the rapid motion resulted from a thermal instability beneath the lithospheric plates generated via a deep-seated mantle plume (sections 3.2, 3.3 and 3.10). We acknowledge that TPW can also generate similar effects, but argue that TPW is not absolutely required to explain the rapid motion of large continents.

Previous Models True polar wandering Kirschvink et al. (1997) argued for an episode of extremely rapid apparent polar wander from the Tommotian through Toyonian interval of the Cambrian (c. 523-508 Ma). Their analysis, if correct, indicates a motion of the entire lithosphere at rates of 66 cm yr -1. Kirschvink et al. (1997) did not view this motion in terms of conventional plate tectonics and instead argued that the entire mantle and lithosphere tumbled through 90 degrees as the intermediate and maximum moments of inertia "interchanged" (see Fig. 3.11-2). They cited, in addition to their analysis of the palaeomagnetic data, the observation that other planets (such as Mars) may also have undergone similar processes, based on the large, observable mass excesses located in the equatorial regions (e.g., Olympus Mons). Later, Mound et al. (1999) modelled the effects of IITPW on sea level changes by using a 25 My duration for the proposed inertial interchange and the palaeogeographic models of Kirschvink et al. (1997). The models suggested that the sea level change was dependent on the location of the continent undergoing the rotation and the duration of the IITPW event. The models qualitatively supported the IITPW hypothesis although the model itself was limited due to available sea level change records for the continents and the inability of the model to account for other possible changes influencing sea level. In more or less the same vein, Evans (1.998) argued that true polar wander is an inherent consequence of supercontinental assembly (sections 3.2, 3.7 and 3.10). Evans notes, as have others (Anderson, 1998; Richards et al., 1999) that the prolate contribution of the non-hydrostatic geoid (spherical harmonic degree 2; ~ = 2) is a long-lived feature associated with supercontinent assembly. In an ideal case, the great-circle trend of TPW could mark the centre of the former supercontinent although an offset of nearly 40 degrees exists in the present dataset. Evans (1998) used the extant palaeomagnetic database from Gondwana and Laurentia in support of his hypothesis that rapid TPW occurred in an oscillatory fashion following the breakup of the supercontinent Rodinia. Meert and Torsvik (in review) also discuss the possibility of TPW during the final assembly of the Rodinian supercontinent (c. 1100-900 Ma). According to the Evans (1998) hypothesis, this TPW episode may have been triggered by instabilities associated with the pre-Rodinian supercontinent of Columbia (Rogers and Santosh, 2002). The suggestion

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Fig. 3.11-2. True polar wander and inertial interchange true polar wander explanations, adapted from Meert (1999). Top figure shows normal true polar wander (after Spada et al., 1992) for both an isoviscous and stiff lower mantle case. The shaded ball represents a mass excess (subduction) and the migration of the spin axis is shown for both cases. The bottom figure shows intertial interchange true polar wander (Kirschvink et al., 1997) with a view along the Imin axis. If the magnitude of the lin t axis exceeds the/max axis, the mantle and lithosphere will tumble through 90 degrees as the axes interchange.

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that subduction around the edges of a large supercontinent surrounding mantle upwellings beneath the supercontinent (e.g., lower mantle mass excesses) will ultimately draw the supercontinent towards the equator (via TPW) is consistent with the known palaeogeographies of Columbia, Rodinia and Pangaea. At the same time, it is important to note that the palaeomagnetic data supporting the older configurations of Columbia and Rodinia along with their latitudinal positions are poorly resolved (Meert, 2002; Meert and Powell, 2001) (section 3.10). Kent and Smethurst (1998) proposed that non-dipole contributions to the geomagnetic field would result in a low-latitude bias of palaeomagnetic data; however, they also acknowledge repeated cycling of landmasses to the equator would also produce the same inclination bias. Finally, Torsvik and Van der Voo (in press) assert that the effects of a persistent octupolar geomagnetic field result in artificially high velocities in the Gondwana palaeomagnetic dataset. Correcting for these non-dipolar fields actually reduces the magnitude of Gondwana's lower Palaeozoic latitudinal motion by 10-12 cm yr- 1. Thus, it is unclear whether or not TPW is an inherent consequence of supercontinental assembly or if other explanations can work in concert with TPW to produce the observed rapid motion of continents described above. Thermal mechanisms

The thermal budget beneath a supercontinent may also play a role in the "speed" at which the continents may move during breakup (Gurnis and Torsvik, 1994; Meert, 1999), although Evans (1998) notes that the observed minimal velocities generated using Neoproterozoic palaeomagnetic data are "more easily reconcilable with TPW" than arguing for changes in specific mantle conditions. However, Gurnis (1988) showed that these special mantle conditions are the expected expression of a supercontinent that forms an insulating lid on the mantle (sections 3.2 and 3.10). Furthermore, Honda et al. (2000) showed that a high-viscosity raft on top of a convecting mantle (i.e. a supercontinent) results in the growth of a mantle plume beneath the supercontinent on time scales ranging from 200-2000 million years. The large variability in the estimates was the result of (a) the type of geodynamic model employed and (b) the initial Rayleigh values. Higher Rayleigh values (107) in three-dimensional rectangular box models resulted in the shortest time period for the generation of mantle plumes. Thus, while the temporal arguments regarding the generation of thermal anomalies are debatable, these anomalies appear to be a natural consequence of supercontinental aggregation (sections 3.2 and 3.9). Eide and Torsvik (1996) argued that rapid continental motion could also be driven by the subduction of old oceanic crust during supercontinent formation. They noted that the formation of high-pressure and ultra-high pressure rocks was often preceded by "bursts" in plate velocities. Thus, in their analysis, the continental plate is pulled towards a cold spot in the mantle (section 3.2). This argument is similar to the explanation for India's rapid migration towards Asia in that continental plates attached to old oceanic slabs will move at higher relative speeds (see also Forysth and Uyeda, 1975; Gordon et al., 1979; Ullrich and Van der Voo, 1981). Meert (1999) combined the two models and argued that rapid motion away from the long wavelength geoid high produced by mantle upwellings and towards long wavelength

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geoid lows produced by mantle downwellings (see also section 3.2) might be able to produce the rapid motion observed in the latest Neoproterozoic.

The H. O. G. Hypothesis While oscillatory true polar wander remains a viable explanation for the observed speed of continents in the Neoproterozoic (Evans, 1998), it lacks testability on a fine temporal scale because of the current poor resolution of the palaeomagnetic database (Meert, 1999). The proposal made here is that warmer mantle conditions coupled to mantle plume activity (sections 3.2, 3.3 and 3.6) can also account for the rapid plate motions observed during the final breakup of the Neoproterozoic supercontinent. Figs. 3.11-3a-c show the disposition of the continents at three distinct times, beginning at c. 800 Ma, again at c. 570 Ma and during the middle Cambrian (c. 510 Ma). While there is some disagreement about the exact configuration of the continents in these reconstructions (see Meert and Powell, 2001; Meert and Torsvik, in review; section 3.10), they form the starting point for the analysis given herein. We propose that rapid continental drift (on the order of 15-25 cm yr -l ) can be driven by thermal buoyancy generated via mantle plumes and the increased heat beneath the supercontinental lid (Gurnis, 1988; Gurnis and Torsvik, 1994) (section 3.2). Figs. 3.11-4a-d show the hypothetical process for generating this rapid drift. Fig. 3.11-4a shows the assembly of a supercontinent containing regions of thick tectospheres some 100 My after formation. As described in Gurnis (1988), the continent is "anchored" in place by subduction zones. The two-dimensional models employed by Gurnis (1988) assumed (largely for computational ease) that the supercontinental lid was either stationary or moving very slowly (1 cm yr-l ) prior to its breakup. On the real Earth, it is likely that the rate of motion is greater than zero although it may have been relatively slow (c. 2-3 cm yr-1). The slowmoving supercontinent serves to accentuate the thermal regime in the underlying mantle through a blanketing effect and may produce a maximum temperature excess on the order of 200 K (Gurnis, 1988). Anderson (1.998) suggests that the thick tectospheres control the tomographic ~ = 6 geoid whilst the supercontinent and associated subduction zones are mainly responsible for the features of the s = 1 and ~ = 2 geoids (see also Scrivener and Anderson, 1992). In Fig. 3.11-4b, the supercontinent has covered the underlying mantle for nearly 200 million years and the mantle plume that began nucleation in Fig. 3.11-4a now begins its ascent through the heated mantle. This ascent, along with the increased thermal buoyancy beneath the tectosphere enhances the tension in the supercontinent and elevates the ~ -- 1,2 geoid. In Fig. 3.11-4c, some 200-400 million years after supercontinent formation, the plume impinges upon the thickened lithosphere. Large scale volcanism, igneous

Opposite: Fig. 3.11-3. (a) Supercontinent Rodinia at c. 800 Ma. The areal extents of plume volcanism at c. 800 and c. 600 Ma are shown by the shaded regions. (b) 580 Ma reconstruction showing the south polar location of eastern Laurentia and southern South America. (c) c. 500 Ma reconstruction showing the eastern margin of Laurentia at c. 30 ~ S and southern south America at c. 30 ~ S.

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intrusion and radiating dyke swarms form in zones of weakness which are exploited under the tensional regime generated by the thermal buoyancy (see also Courtillot et al., 1999) (sections 3.2 and 3.3). In Fig. 3.11-4d, the supercontinent begins to break apart. Although the tectosphere acts as an inhibiting force during upper mantle convection, the plume "exploits" the tectosphere and accentuates the motion of the continent off the thermally buoyant geoid high (analogous to the model of Gurnis and Torsvik, 1994). This can be envisioned by assuming that the mantle plume acts as a hand which uses the tectosphere to "grab" the continent and "throw" it off the superheated region. One can playfully call the plume "The Hand O' God", or H.O.G. for short. The continent is then pulled towards the s = 1,2 geoid lows (see also section 3.2). The models used by Gurnis (1988) indicated peak velocities for supercontinental breakup of up to 7 cm yr - l . This is significantly less than the 15-25 cm yr -1 observed in the Neoproterozoic interval. In the Gurnis and Torsvik (1994) model, the augmented velocities were dependent on dimensions of the lithospheric root and the lateral temperature contrast arising from a heat source originating in the lower mantle (see two definitions of mantle plumes, section 3.3). They estimated a maximum temperature contrast of 160 K over 500 My that resulted in a c. 6 cm yr-1 augmented velocity (i.e., above the background velocity). Gurnis and Torsvik (1994) also noted that if the drift was driven additionally by the presence of a mantle cold spot, then this might also supply a 10 cm yr-l augmented velocity to the continental plate. The article did not supply an absolute limit to plate velocities but suggested that speeds of up to 20 cm yr-lmight be expected. Campbell and Griffiths (1990) calculated that the excess temperatures in plumes could reach temperatures of up to 200-300 K above ambient. Thus, a combination of the thermal anomaly associated with supercontinental insulation (200 K) in addition to the thermal anomaly of the mantle plume (100 K) would produce a total lateral temperature gradient of up to 300 K. According to the analysis made by Gurnis and Torsvik (1994), a 300 K lateral temperature gradient would result in an augmented increase in velocity of c. 10-12 cm yr-I for a 250 km thick lithospheric root. Our model is qualitative and is based upon the conclusions of previously published geodynamic models discussed above. The prediction that mantle plumes coupled with an enhanced thermal regime beneath supercontinents can drive plates rapidly away from the geoid highs and towards the geoid lows (see also section 3.2) can be tested by geodynamic models. Geologic consequences of the model are discussed briefly in the next section.

Neoproterozoic palaeogeography and H. O. G. Rodinia (Fig. 3.1 l-3a) was thought to have formed largely during the 1100-1000 Ma Grenvillian orogeny (Dalziel, 1997) although parts of the supercontinent may have coalesced earlier (see also section 3.10). The supercontinent began to break up at c. 800 Ma (Li et al., 1999, 2002) with the arrival of a mantle plume beneath south China (along the present-day western margin of Laurentia). Frimmel et al. (2001) (section 3.10) noted that extensive igneous activity in South Africa (Richtersveld Igneous complex) was broadly coeval with the south China event. They proposed that a c. 800 i 60 Ma megaplume stretched from S. Africa through Australia-south China and northernmost Laurentia

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Fig. 3.11-4. Note: The figures are not to true scale; LM represents a segment of the Lower Mantle and UM represents a segment of the upper mantle; the dashed line represents missing mantle between the UM and LM. (a) Supercontinent + 100 My after formation, with increased sub-lithospheric temperatures. Plume nucleation at the Core-Mantle Boundary (CMB) begins and a region of cold mantle downwellings may appear at the boundaries of the supercontinent as well as in regions of old oceanic crust. The s = 1, 2 geoid is shown in idealised fashion next to the figure. (b) The supercontinent +200 My after formation. The plume is now formed and is ascending in the warm mantle region beneath the supercontinent. The supercontinent is now under increased tension due to the elevated geoid caused by the thermal anomaly beneath the supercontinent.

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Fig. 3.11-4 (continued).

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~ee also Park et al., 1995; Wingate et al., 1998; Wingate and Giddings, 2000; Foden et al., 2002). If these estimates are correct, then the mantle plume(s) developed beneath the Rodinia supercontinent some 200-300 My after its formation. This timing is in agreement with one of the proposals by Honda et al. (2000). Unfortunately, we have difficulty estimating minimum plate velocities during the initial breakup of the supercontinent due to a (near-complete) lack of palaeomagnetic data during the 750-600 Ma interval (see Meert and Powell, 2001). The model given here would predict a short "burst" of rapid plate motion, but its testing must await further refinement of the palaeomagnetic database. The supercontinent was not fully disaggregated during this first rifting phase and it was followed some 200 million years later by a second pulse of plume activity (the so-called Sept Iles plume; Higgins and van Breeman, 1998). The presence of this plume has clear manifestations in eastern Laurentia and possible manifestations in Baltica (Bingen et al., 1998; Meert et al., 1998), but there is no clear evidence of the plume in the South American blocks. However, there is some controversy regarding the relationship between the South American cratons and eastern Laurentia (Tohver et al., 2002; Meert and Torsvik, in review). Palaeomagnetic data from Laurentia during the 570-510 Ma interval is estimated conservatively to give drift rates of 16-20 cm yr-l (Meert, 1999). Palaeomagnetic data from Gondwana show rapid motion of western Gondwana over the pole from c. 550-500 Ma (c. 24 cm yr-l ; Meert et al., 2001) although the magnitude of this motion may be reduced significantly if non-dipolar fields are considered (Torsvik and Van der Voo, in press). Conclusions and Possible Tests

Motion of Laurentia and Gondwana, as constrained by palaeomagnetic data, during the Late Neoproterozoic through Middle Cambrian interval was rapid. An absolute upper "speed limit" for large continental plates with thick tectospheres is not established although such a limit must exist. Geodynamic models of supercontinents (Gurnis, 1988; Gurnis and Torsvik, 1994; Honda et al., 2000) indicate that the mantle will warm beneath the supercontinent and plumes can form within 200-400 million years after assembly (section 3.2). Gurnis and Torsvik (1994) estimated that the plate velocity of continents with thick tectospheres can be augmented provided the mechanism was deep-seated. Honda et al. (2000) showed that mantle plumes are a natural result of supercontinent assembly and thus provide the deep-seated mechanism necessary for the rapid drift of continents (section 3.2).

Fig. 3.11-4 (continued). (c) The supercontinent at +400 My. The plume has now impinged upon the tectosphere and exploited previous weak zones (former sutures) producing flood basalts and an elevated ~ = 1,2 geoid. The increased tension coupled with the "cold" mantle downwellings begins to break apart the supercontinent. (d) Initial breakup of the supercontinent and rapid drift towards geoid lows. The plume becomes entrained in the moving continent and effectively "throws" the continent off the geoid high. A large piece of the supercontinent may still exist and serve to generate a second episode of thermal buoyancy and rifting.

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The velocity can also be enhanced as the continents drift from the long wave length geoid highs towards cold (= low) spots produced by subduction. The presence of large plumes beneath the Rodinian supercontinent (see also section 3.10) and its vestiges are indicated at 800 + 60 and 585 + 30 Ma. The temporal development of these plumes is consistent with the models of Honda et al. (2000) and may provide the driving force necessary for the observed rapid drift of Laurentia and Gondwana. Due to a lack of palaeomagnetic data, we are unable to document minimum plate velocities during the initial breakup of Rodinia, but plate velocities during the Late NeoproterozoicEarly Cambrian interval are consistent with the models of Gurnis (1988) and Gurnis and Torsvik (1994). Long-lived subduction along the assembling (and assembled) margins of Gondwana (e.g., closure of the Mozambique, Brasiliano and associated oceans) may have resulted in a region of cold mantle that helped augment the motion of Gondwana. There are some testable side effects of the H.O.G. mechanism. For example, as noted by Gurnis and Torsvik (1994), flooding of continents (transgression) will occur as the continent moves off the geoid high towards cold regions of the mantle. In the case of Cambrian Laurentia, the Sauk transgression may have resulted from the rapid motion away from a regional of thermal buoyancy towards a geoid low. The lag between continental flooding and rifling in Laurentia is c. 50 Ma. We also envision that the earlier phase of rifling (800+60 Ma) in western Laurentia may show a similar lag between the onset of rifling and the development of marine facies. Second-order sea level effects (see also sections 8.1 and 8.2), predicted from geodynamic models (Gurnis, 1990) should also be represented in the sedimentary record (see also section 3.2). In contrast to Laurentia, the interior of Gondwana remained largely emergent during the Cambrian (Veevers, 1995). As noted by Veevers (1995), assembled Gondwana continued to blanket the mantle until its breakup (see also section 3.10). The thermal buoyancy produced by mantle insulation may explain the pulses of rapid motion for Palaeozoic Gondwana and ultimately, the massive outpouring of flood basalts during its Mesozoic breakup (Meert et al., 1993; Veevers, 1995; e.g., the Ferrar-Karoo-Parana provinces). Interestingly, the time period between final Gondwana assembly (c. 530 Ma) and the Mesozoic igneous events associated with its breakup (Fig. 3.11-5) is on the order of 400 My, consistent with the estimates of Honda et al. (2000). Thus, we note, as have many others, that the formation of a long-lived supercontinent may impose a thermal structure on the underlying mantle leading to its ultimate demise (section 3.2). Finally, we acknowledge that the combination of increased thermal buoyancy and longlived subduction beneath the margins of the assembled Gondwana continent may have led to inertial instabilities and true polar wander (Van der Voo, 1994; Evans, 1998). Excitation of true polar wander during the initial breakup of Rodinia (c. 750 Ma) is unlikely since the supercontinent straddled the equatorial regions, although its equatorial position may have resulted from TPW. Proposed episodes of Vendian-Cambrian TPW are possible given the geodynamic mechanisms proposed elsewhere (Kirschvink et al., 1997; Evans, 1998), but the palaeomagnetic record is currently too sparsely populated to provide a rigorous test of those hypotheses.

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267

Fig. 3.11-5. APWP for Gondwana during the three intervals of Neozoic rapid APW (550-520 Ma; Meert et al., 2001), 475--420 Ma (Meert et al., 1993) and 420-340 Ma (Meert et al., 1993). Major flood basalt provinces are shown: the Parana-Etendeka (133-131 Ma), the Karoo-Ferrar (184 Ma), and the Deccan (65 Ma); after Courtillot et al. (1999). The interval from final Gondwana assembly (taken here as 530 Ma) and the observed volcanism ranges from 350-465 My, in line with the estimates of Honda et al. (2000).

3.12.

COMMENTARY

EG. ERIKSSON AND O. C A T U N E A N U Geological evolution during the Hadaean remains speculative, but Trendall's (2002) "plughole" model provides for the possibility of a gradual transition (in the c. 4-2.5 Ga interval, and probably diachronous for the various early Archaean cratonic nuclei), from whole mantle convection and an Earth dominated by thermal processes to one where layered mantle convection enabled plate tectonics to become increasingly dominant (section 3.6). Although identification of Archaean ophiolites is difficult (partly due also to preferential preservation of the upper parts of Archaean ocean crust) and even controversial to some, those that are inferred suggest a genesis encompassing thickened ocean 7"he Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. ('atuneanu

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plateaus in concert with significantly higher heat flow (Moores, 2002). Identified or inferred ophiolites older than 1 Ga support temporal change in Earth's geothermal gradient, increasing mantle heterogeneity, and thinning of the ocean crust with time, with major breaks at c. 1 and 2.5 Ga; oceanic thinning in the Proterozoic may have led to the onset of conventional plate tectonics (Chiarenzell and Moores, section 3.7). Certainly, the formation of granite-greenstone crust (section 1.2, chapter 2) and the occurrence of komatiitic lavas are essentially Archaean in age (section 3.4). The widely accepted value of 2-3 times (modern-Phanerozoic) mantle heat flow in the Archaean suggests that a more chaotic mantle convection regime was active; consequently, catastrophic magmatic events could have been important in crust formation throughout most of the Archaean (Nelson, section 3.4). However, magmatism was probably relatively localised rather than global in expression during the Early Archaean (cf. Trendall, 2002; section 3.6). It is possible that the earliest gneissic and sialic protocratonic nuclei formed at c. 4 Ga, through essentially thermal-magmatic processes (Trendall, 2002). It is surmised that parts of many, or even most Archaean cratons were underlain by low Rb-Sr-type metasomatised lithospheric mantle, and that this laterally relatively widespread mantle enrichment resulted from fluid release during shallow subduction (section 3.5). Shallow subduction may thus have been reasonably common, at least in the Neoarchaean (Cousens et al., section 3.5) when plate tectonism had probably become important in global geodynamic processes. It should be noted that shallow subduction, when applied to an earlier onset of the plate tectonic paradigm than that envisaged in Trendall (2002), could explain, at least partially, the operation of plate tectonics in the Archaean (de Wit, 1998). A superplume in the usage of Ernst et al. (section 3.3) encompasses buoyant material rising through the mantle irrespective of depth of origin, whereas Condie (section 3.2) defines it as rising from the deep mantle. Large igneous province (LIP) events in the Archaean, identified on the basis of the presence of komatiite and possibly represented by certain classes of greenstone belts, are less certain; a decreased plume frequency prior to 2.8 Ga may merely be an artifact of analysis of available data (Ernst et al., section 3.3). A major change in Earth's geological evolution and its first-order thermal and tectonic controls is apparent in the Neoarchaean. LIPs (Ernst et al., section 3.3) and the mantle superplumes (Condie, section 3.2) inferred to be responsible for their origin, increase in frequency in the geological record at 2.8-2.7 Ga (Condie, 200 la; Ernst and Buchan, 2002a, b). Nelson (section 3.4) suggests that catastrophic mantle overturn events became global in scale at c. 2.7 Ga, as the transition to a plate-tectonically-dominated Earth became increasingly effective; large volumes of granite-greenstone crust formed on both the Yilgarn and Superior cratons between 2760 and 2620 Ma, including the eruption of komatiites in greenstone belts in both at 2705 Ma (Nelson, 1998a). Several relatively well preserved Palaeoproterozoic ophiolite complexes are recognised at the suture between the Churchill and Superior Provinces (section 3.7), and possible ophiolites occur in the c. 2.7 Ga Limpopo orogenic belt in southern Africa (section 3.8). A first supercontinent is postulated for c. 2.7 Ga ("Kenorland"; e.g., Aspler and Chiarenzelli, 1998) (see also section 3.9). Evidence for relatively abundant Neoarchaean greenstones with oceanic plateau-type geochemistry suggests that these were major con-

3.12. Commentary

269

tributors to this supercontinent (Condie, 1994b; Tomlinson and Condie, 2001), along with pre-2.7 Ga crustal fragments, and, possibly also, ocean arc systems (sections 3.2 and 3.6). The relative buoyancy of Archaean subcontinental lithosphere, being less amenable to subduction, also contributed (Condie, section 3.2). Identified and inferred ophiolite complexes cluster in time at c. 1-1.5, 1.8-2.3, 2.5-2.7 and at c. 3.4 Ga, suggesting a causative link with the supercontinent cycle; obviously ophiolite emplacement will precede continentcontinent collision (Chiarenzelli and Moores, section 3.7). This first supercontinental assembly may reflect the first catastrophic slab avalanche event at c. 2.7 Ga, as plate tectonics became significant and as slabs possibly reached a critical mass at the 660-km mantle discontinuity; avalanching into the lower mantle may have triggered the first superplume event (e.g., Peltier et al., 1997; Condie, 1998). Major superplume events were most likely associated with supercontinents close to their terminations, and the two major such events inferred, at c. 2.7 and 1.9 Ga, were associated with globally elevated sea levels (chapter 8), peaks in stromatolite (section 6.5) occurrence and diversity, and significant changes in ocean chemistry (Condie, section 3.2; see also section 5.2). The c. 2.7 Ga Ventersdorp continental flood basalts suggest a direct plume hit on the Kaapvaal craton, thus providing high freeboard (Eriksson, 1999) and leaving no direct evidence for global eustatic rise (Eriksson et al., 2002b). Although the age of the Limpopo orogenic belt is subject to debate, evidence of collision between the Kaapvaal and Zimbabwe cratons at c. 2.7 Ga is strong (section 3.8). Mints and Konilov (section 3.9) discuss evidence for plume-related origin for many Palaeoproterozoic high-grade mobile belts, with major superplumes affecting Fennoscandia at c. 2.52-2.44 Ga, and a more widespread superplume event at c. 2-1.95 Ga. The LIP record appears to be relatively constant, except for gaps at 615-720, 2220-2400 (Mints and Konilov, section 3.9, also suggest predominant quiescent within-plate geodynamic processes from 2.44-2 Ga) and 3000-3300 Ma, although these may merely reflect artifacts (section 3.3). Rates of approximately one continental LIP every 20 My from 2.5 Ga apply, and a weak cyclicity (c. 170, 330 and 730-600 My) is evident, but applied only to limited portions of the post-Archaean record (Ernst and Buchan, 2002a, b). The interplay of plate tectonics and mantle (super)plumes continued throughout the Precambrian (and later) as the predominant first-order control on Earth's lithospheric evolution, particularly with respect to the supercontinent cycle. Formation of a Neoproterozoic supercontinent (Rodinia) at c. 1.2 Ga is widely accepted, despite large differences of opinion on its configuration (sections 3.10 and 3.11). A causative link with a superplume event during its formation is questioned by Frimmel (section 3.10), who also notes episodic continental growth around the Kalahari craton; however breakup, which began at c. 750 Ma around the Kalahari craton, was related to a thermal mantle anomaly (section 3.10). Meert and Tamrat (section 3.11) discuss evidence for rapid post-breakup motion of large tectonic plates related to supercontinental blanketing of mantle heat (Gurniss, 1988; Gurniss and Torsvik, 1994; Honda et al., 2000), emphasising that warming of the mantle after assembly can be augmented by plumes which may form 200-400 My after assembly. The plumes provide the deep mechanism to enhance post-breakup velocities of plates with thick tectospheres (Gurniss and Torsvik, 1994; Honda et al., 2000). Large plate velocities

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can also be enhanced by supercontinental fragments drifting from long wave length geoid highs towards cold spots (cf. geoid lows; Condie, section 3.2) resulting from subduction, as illustrated by Meert and Tamrat (section 3.11) for the rapid drift of the Laurentian and Gondwana plates after breakup of Rodinia.

The Precambrian Earth: Tempos and Events Edited by RG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) Published by Elsevier B.V.

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Chapter 4

PRECAMBRIAN VOLCANISM: AN INDEPENDENT VARIABLE THROUGH TIME

4.1.

INTRODUCTION

W.U. MUELLER AND RC. THURSTON Volcanic rocks constitute prominent segments of Archaean and Proterozoic greenstone belts that are generated by plate subduction, which includes divergent margin rifting, and by mantle plumes. A shift in mantle conditions and geodynamic processes from plumegenerated komatiites (Campbell et al., 1989; Abbott and lsley, 2002b) and boninites, adakites and TTG (tonalite-trondhjemites-granodiorite) plutonic suites, a product of lowangle subduction tectonics (Chown et al., 1992; Polat et al., 2002; Wyman et al., 2002a) (section 3.5), during the Archaean-Palaeoproterozoic, to classic high-angle subduction and an absence of komatiite-generated magmas occurs as of the Mesoproterozoic. The question arises, can volcanic change through time be identified via mantle and subduction processes? If so, would volcanic successions provide an efficient way of tracing geodynamic evolutionary trends? Changes in Earth's evolution include: (1) the atmosphere with respect to oxygen and carbon dioxide contents (Kasting, 1993) (sections 5.2-5.5), palaeosols (Beukes et al., 2002) (section 5.10), or chemical alteration of sedimentary rocks (Nesbitt and Young, 1982) (section 5.10); (2) the hydrosphere (Holland, 1984) (sections 5.2-5.5); (3) the development of life (Schidlowski, 1988; Altermann, 2002) (chapter 6); (4) the influence of vegetation or the absence thereof on fluvial dispersal deposits (Schumm, 1968) (section 7.8); and (5) Earth and Moon tidal periodicity (Kvale et al., 1999; Eriksson and Simpson, 2000) (sections 5.9 and 7.5). Magmatism, and hence volcanism, generated both oceanic and continental crust, but also affected the atmosphere and hydrosphere with the emission of H20, CO2, SO2 and CO gases (e.g., section 5.2). Effectively, volcanism is an independent variable influenced by mantle and crustal processes; however, the repetition of mafic to felsic volcanic sequences in greenstone belts suggests cyclical magmatic behaviour. Still, volcanic stratigraphy must be taken with a grain of salt because the stratigraphic concepts of Walther's law (see Blatt et al., 1980), which states that adjacent depositional units in space occur sequentially in crustal profile, do not apply in sensu stricto. In optimal cases, the volcanic and sedimentary stratigraphy is comparable, but ancient volcanic sequences rely heavily on U-Pb age constraints because intrusions and sills/dykes may be unrelated to edifice construction (e.g., Mueller and Mortensen, 2002), and discerning between intrusive versus extrusive volcanic rocks can be difficult. This chapter focuses on the principal characteristics of volcanism during the Precambrian and on features requiring special consideration. The main themes are: (1) komati-

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Chapter 4: Precambrian Volcanism

ites, which represent a window on early Earth, (2) setting and evolution of Archaean and Proterozoic greenstone belts (sections 2.3 and 2.4), and (3) subaqueous explosions and Archaean calderas. In addition, volcanic terminology is discussed to facilitate our understanding of volcanic processes.

The Independent Variable: Volcanism The driving mechanisms of volcanism for early and modern Earth remain theoretically the same, with voluminous and devastating historical island arc eruptions o f Taupo (C.J.N. Wilson et al., 1995) or Krakatoa (Simkin and Fiske, 1983) having Archaean(Tass6 et al., 1978; Hudak et al., 2002a) and Proterozoic arc counterparts (Hildebrand, 1981). Similarly, plume-controlled continental volcanism exemplified by the Columbia River basalts or Yellowstone rhyolites (Anders and Sleep, 1992; R.I. Hill et al., 1992) have the 2.7 Ga Ventersdorp Supergroup flood basalt (Eriksson et al., 2002b) and ignimbrite analogues (van der Westhuizen and de Bruiyn, 2000). Evidently, explosive volcanism was operative, but is there a change in style of volcanism? An oxygen-poor atmosphere would not affect the eruption mechanism or the transport process, nor would the Precambrian hydrosphere change subaqueous pyroclastic transport mechanisms. Intuitively, the fundamental difference between early and modern Earth should be in the volume of magma generated at mid-ocean ridges and convergent plate margins. Supposedly, Archaean effusion rates were higher because of a higher geothermal gradient and rapidly colliding microplates (Bickle, 1978, 1986; Sleep and Windley, 1982; Galer and Metzger, 1998) (see also sections 2.8 and 3.6), but mass balance calculations by Dimroth (1985) suggested rates of magma emplacement for the Abitibi greenstone belt were similar to Mesozoic-Cenozoic arcs. Thurston (1994) drew similar conclusions. A c. 80 km spacing between Abitibi arc edifices compares favourably to modern arcs (Windley and Davies, 1978), and hence the notion of magma generation. Apparently arc systems are not the locus to discern a change through time. Why? Arcs, independent of age, are formed by plate motion with magma generated by subduction. Higher Archaean-Palaeoproterozoic temperatures (section 3.6) would only cause magma-generation at shallower levels. Boninites and adakites may be the response to shallow subducting plates (section 3.5), because high heat flow conditions with rapid subduction of young oceanic crust are required (Kerrich et al., 1998; Leybourne et al., 1999;Komiya et al., 2002; Wyman et al., 2002a). Alternatively, mid-ocean ridges represent a site of magma generation, which might have changed with time. According to Fisher and Schmincke (1984), ridges represent the primary locus of heat loss and generation of new oceanic crust. Although this is the case at present, and possibly for the Precambrian Earth, it is not possible to quantify the volume of extrusive volcanic rocks at these sites. The volume of effusive volcanism at oceanic ridges was probably higher during the Archaean but these assumptions are speculative at best (see also section 3.6). The physical volcanology of certain basalt sequences, as suggested by Wells et al. (1979), showing a dominance of massive lava flows, is most likely a function of vent proximity rather than higher effusive rates during the Archaean.

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273

So where is the difference? The principal breakthrough came with the recognition of komatiites, an extrusive rock with 20-30% MgO (Viljoen and Viljoen, 1969a, b). Komatiites are inferred to originate from mantle plumes (Campbell et al., 1989) and their abundance in the Archaean and absence in the present (see also sections 3.2-3.4) has major implications for Earth's evolution. Plume-generated volcanism produces voluminous eruption fields occurring over tens of millions of years and was more voluminous during the Archaean and Early Proterozoic as inferred by the notion of superplume events (Nelson, 1998a; Abbott and Isley, 2002b). Precambrian superplume events between 1.7 and 2.9 Ga (sections 3.2 and 3.3) may have completely resurfaced the planet (Abbott and Isley, 2002b), with magma volumes ten times larger than Phanerozoic plume counterparts. Extensive (radial) mafic dyke swarms are prime examples of plume influence (Fahrig, 1987). The second but less evident difference relates to a dearth of orogenic andesites in Archaean arc sequences. Abbott and Hoffman (1984) suggested that a hotter Archaean Earth would result in low-angle subduction of young, hot oceanic crust, which would result in production of more siliceous melts (Helz, 1976) and bimodal volcanism. A "cooler" mantle with highangle subduction zones favours the formation of orogenic andesites with plagioclase and clinopyroxene phenocrysts (Gill, 1981). The principal refinement in Precambrian volcanology is the notion of long-term coeval interaction of mantle plumes and subduction zones, as documented in Australia (Nelson, 1998a) and Canada (Dostal and Mueller, 1997; Kerrich et al., 1998; Wyman et al., 1999a).

4.2.

TERMINOLOGY OF VOLCANICLASTIC AND VOLCANIC ROCKS

W.U. MUELLER AND J.D.L. WHITE Naming a volcanic rock is a problem and volcanic terminology has been a source of intense discussion (Fisher and Schmincke, 1984; Cas and Wright, 1987; Stix, 1991; McPhie et al., 1993; White, 1994; Orton, 1996). The terminology used here is based on the criteria of Fisher (1961, 1966). By definition, clastic rocks containing abundant volcanic material irrespective of their origin or environment are referred to as volcaniclastic (Bates and Jackson, 1987, p. 715). The adjective volcaniclastic is therefore an umbrella term, which encompasses pyroclastic, autoclastic and epiclastic, and can be employed for Precambrian rocks in which particle origin remains enigmatic. Volcaniclastic Rocks

Nomenclature becomes a problem when authors use different criteria to assign a name to a volcaniclastic deposit, and this is especially evident in the usage of the terms "tuff" and "sandstone". These terms represent grain size classes in volcanology and sedimentology, respectively, and have important connotations concerning origin and transport process. Tuff implies a volcanic origin, in which particles derive from explosions or thermal granulation, and sandstone reflects an epiclastic origin, whereby grains originate from erosion of con771ePrecanzbrian Earth: Temposand Events Edited by P.G. Eriksson, W. Ahermann, I).R. Nelson, W.U. Mueller and O. Catuneanu

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Chapter 4: Precambrian Volcanism

solidated rock (Schmid, 1981; Fisher and Smith, 1991). The eruption mechanism, erosion process, transport process, depositional setting, transporting medium, type of constituents and their abundance, and grain size are all qualifiers that influence rock classification. The varying usage of these terms depends on what processes authors emphasise: (1) particle formation and fragmentation mechanism, i.e., tuff, or (2) the transport medium and process after deposition and lithification, i.e., sandstone. Fisher and Smith (1991) argued that transporting agents, such as wind and water, do not change the origin of the components. Because eruptions strongly modify surface environments and add large amounts of debris to sedimentary systems without an intervening weathering stage to temper supply rates (section 7.3), this origin is of fundamental importance even where grains are redistributed soon after eruptions by sedimentary processes. Consequently, unconsolidated, remobilised tephra transported via wind, fluvial currents or subaqueous sediment gravity flows, and with preserved pyroclasts are best characterised by Fisher's (1961) pyroclastic scheme, which recognises the eruptive origin of the deposit. A modifier of "reworked", or one specifying a depositional setting, should be added where such reworking is demonstrable (i.e., fluvial cross-bedded tuff). The Australian school (Cas and Wright, 1987; McPhie et al., 1993), in contrast, prefers a sedimentary classification scheme for any unconsolidated or consolidated deposit formed of pyroclastic particles that have, either possibly or demonstrably, undergone reworking by wind or water (i.e., crystal-rich sandstone or graded bedded sandstone), as well as for deposits where the nature of final transport and deposition is unknown. This extends to all deposits, in which final transport was by water, even those from subaqueous eruptions where the particles travel directly from the vent to the depositional surface. Defining the term "pyroclastic" is not only semantic. A redefinition of "pyroclastic" by Schmid (1981 and lUGS Subcommission on the Systematics of Igneous Rocks) included particles that are a "direct result of volcanic action" rather than "generated by disruption during volcanic eruptions". A fragment from phreatomagmatism is a "variety of pyroclast formed by steam explosions at magma-water interfaces, and also by rapid chilling and mechanical granulation of lava that comes in contact with water or water-saturated sediments" (Fisher and Schmincke, 1984, p. 89). Hyaloclastites are products of explosions and/or thermal granulation (Schmid, 1981), and would therefore be considered pyroclasts. Genetic terms such as ash-flow tuff, referring to deposits from gaseous pyroclastic flows, typical of caldera-forming eruptions, should be used only after detailed study because specific eruption processes, transport mechanisms and bedding features are implied. Fisher (1961, 1966) established a scheme based on grain size and deposit components. The scheme is "field-user friendly" because it accommodates both the historically important pyroclastic rock names (e.g., ash-flow tuff), and enables a comparison of rocks at the hand-specimen or thin-section scale. In Table 4.2-1 the Wentworth scale is the basis for a volcanic grain size classification that incorporates the notions of Schmid ( 1981), and Fisher and Schmincke (1984). It is extended to remove the anomaly of having only two grain size classes defined in the sub-2 mm size range, and applies "sand" grain size subdivisions for the tuff or ash (< 2 mm) range, while retaining lapilli (2-64 mm) and breccia or bomb categories (> 64 mm). This expanded scheme is an unobtrusive supplement to the ash/tuff size

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275

range. The redefinition of "fine ash" from a pan-fraction to between 0.125 and 0.25 mm, and ash particles < 0.0625 mm to mud-grade ash is a logical consequence. The lithified debris has sedimentary grain size counterparts with mudstone-siltstone, sandstone, or conglomerate. Furthermore, pyroclastic rocks in this sense can be named from drill core, without any requirement of knowing the specific transport and depositional processes involved in accumulation of the deposit. The abundance and size of pyroclastic components in bedding units is considered so that composite grain size names result, such as lapilli tuff, tuff breccia or lapilli tuff breccia, with individual components requiring >~ 25%. Volcanic breccia (> 2 mm) and sandstone (< 2 mm) are general terms describing volcanic rocks composed of unabraded lithic fragments, irrespective of their fragmentation origin (Fisher and Schmincke, 1984, p. 92) and are best used as a first-order field descriptor if outcrop quality and/or extent prohibit further assessment. To better understand the type of volcanic deposit, an adjective indicating the prominent transport process (i.e., turbiditic tuff), the composition (i.e., felsic tuff), prevalent component (i.e., vitric tuff), or physical nature (i.e., massive lapilli tuff breccia) will facilitate matters. Volcanic Rocks

The physical volcanology of lava flows is easier to assess because they are not dependent, as are pyroclastic rocks, on the transporting medium gas-air or water, and because flows and their structures are readily recognised. Terms such as hyaloclastite breccia, flow breccia, pillow breccia or flow top breccia should be employed for autoclastic fragmentation processes. A descriptive attribute indicating composition or a physical characteristic such as in "flow-banded rhyolite lava flow", "columnar-jointed basalt flow" or "feldspar-phyric rhyolite flow/lava" is useful if it represents a major portion of the flow. Identifying large scale volcanic structures is important but also problematic in areas of limited outcrop. Felsic domes are three-dimensional bodies composed of flows, which are massive, lobate and brecciated. These dome-flow complexes have been recognised readily in subaqueous Archaean sequences (de Rosen-Spence et al., 1980; Gibson et al., 1997; Lafrance et al., 2000; Mueller and Mortensen, 2002). The descriptive term lobe-hyaloclastite is generally associated with extrusive domes (de Rosen-Spence et al., 1980; Gibson et al., 1997). Discerning if domes or lobe terminations are extrusive or intrusive is often difficult. Endogenic (intrusive) domes (e.g., Goto and McPhie, 1998; Lafrance et al., 2000), also referred to as cryptodomes (McPhie et al., 1993), may cause significant inflation of the edifice (e.g., Mount St. Helen's prior to the May 18th 1980 eruption). In addition, gravitational collapse of domes causes catastrophic pyroclastic block and ash flows (e.g., recent Merapi eruptions and 1991 Unzen eruption), but these are difficult to recognise in the ancient rock record. With clear field relationships, autoclastic breccia deposits mantling domes are commonly called carapace breccias or flank breccias. Facies mapping based on phenocryst composition and their abundance permits distinction of endogenic lobes in thick breccia sequences. In contrast, mafic rocks favour the formation of thick sill-shaped bodies significantly inflating edifices or sequences. Each author conveys a different message with deposits, expressing either a fragmentation mechanism, or the transport process, or a geographic position on

Table 4.2-1. Expanded, wentworth-based, grain size scheme for pyroclastic rocks Grain size Finer than 4 phi (< 0.0625 mm) Between 4 and 3 phi (0.06254.125 rnm) Between 3 and 2 phi (0.125-0.25 mm) Between 2 and 1 phi (0.254.5 mm) Between 1 and 0 phi (0.5-1 mm) Between 0 and - 1 phi ( 1-2 mm) Between - l and -2 phi (2-4 mm) Between -2 and -4 phi (4-16 mm) Between -4 and -6 phi (1mm) Coarser than -6 phi (> 64 mm)

Schmid (1981). Fisher and Schmincke (1984) Fine ash1

Coarse ash

~a~illi~ Lapilli

Blocks and bombs

Unconsolidated deposit name Mud-grade ash

Rock name Mud-grade tuff

Very fine ash

Very fine tuff

Fine ash

Fine tuff

Medium ash

Medium tuff -?

Coarse ash

Coarse tuff

Very coarse ash

Very coarse tuff

Fine lapilli (lapilli bed4) Medium lapilli

Fine lapillistone

Coarse lapilli

Medium lapillistone Coarse lapillistone

Blocks and bombs

Breccia

Complete rock name Muds-grade tuff Very fine-grained tuff Fine-grained tuff Medium-grained tuff" Coarse-grained tuff5 Very coarse-grained tuff Fine lapillistone Medium lapillistone Coarse lapillistone Breccia

Notes: '"Ash" is an aggregate name; single particles are ash grains, or ash particles. 2's~apilli"is a plural particle name (singular is lapillus); aggregates of lapilli alone form a deposit, e.g., lapilli unit, lapilli bed. 'Deposits or rocks comprising a mixture of grains within a single major class, such as a lithified aggregate of fine to coarse ash, default to the class name, e.g.. "tuff" rather than "fine-medium-coarse tuff". 4 ~ e p o s i t or s rocks composed of a mixture of grain sizes are modified in the same way as are sedimentary rocks using the Wentworth scale, e.g., "lapilli a s h for ash containing r 25% lapilli and ash components (cf. pebbly sand), or "ash-bearing lapilli bed" for bed of lapilli with subordinate ash (cf. sandy [pebble] gravel). "Tuff breccia" is a rock containing > 25% blocks or bombs with a > 25% lithified ash matrix (cf. sandy conglomerate). 5 ~ h attribute e "-grained represents the full rock name in the tuff grade scheme and is comparable to "fine-grained sandstone".

s

8s P b

a

2 3 y

ij.

E 2

3.

3

4.3. K o m a t i i t e s

277

an edifice. Whatever rock description is preferred, it should be based on field observations and the reader should be informed how the terms are being used.

4.3.

KOMATIITES: VOLCANOLOGY, GEOCHEMISTRY AND TEXTURES

Komatiites are ultramafic effusive and intrusive volcanic rocks that are almost entirely restricted to the Archaean and Palaeoproterozoic. With inferred eruption temperatures of c. 1600~ (Nisbet, 1982; Nisbet et al., 1993a; Arndt, 1994), these low viscosity flows (0.1-1 Pas/1-10 Poises) reflect composition, temperature, and melting processes in the early Earth's mantle (see also section 1.2). This section considers the flow features of komatiites, addresses their geochemistry, and discusses specific komatiite and tholeiitic basalt textures. The petrology of flows, significant for source modelling and mantle conditions, requires further consideration that is not possible here (see Arndt, 1994; Parman et al., 1997).

4.3.1

Physical Volcanology of Komatiites

W. U. Mueller Komatiites have been identified historically and classified according to internal textural zoning in flows, referred to as A and B zones with Am-A3 and B l-B4 divisions (Pyke et al., 1973; Arndt et al., 1977; Fig. 4.3.1-1). Spinifex textures (Viljoen and Viljoen, 1969a, b) are confined to the A2-A3 divisions. The A-zone with polygonal joints (thermal contraction) and olivine or pyroxene spinifex reflects a cooling history, whereas the B zone, with olivine or pyroxene phenocrysts, indicating accumulation and settling, displays the crystallisation history (Lajoie and G61inas, 1978; Renner et al., 1994). Although spinifex flows receive most of the attention, they are subordinate compared to prominent massive flows and locally prevalent vesicular units. Focus is placed on the geometry of well-exposed flow fields as well as physical features and internal textural zones of individual flows in the Abitibi (Champagne et al., 2002) and Barberton (Dann, 2000, 2001) greenstone belts. Case studies and flow models, notably from Kambalda (Yilgarn Block, Australia) have been presented (Lesher et al., 1984; Thomson, 1989; Hill et al., 1995; Moore et al., 2000; Beresford et al., 2002), but this body of work is prominently based on drill core due to outcrop paucity so that interpretations are limited. Komatiite flow models of Hill et al. (1995), with an anastomosing pattern of channel and levee/overbank deposits, draw on previous subaqueous mafic flow models (e.g., Dimroth et al., 1978, 1985). But are these correct? Points of debate, which require further studies, are thermomechanical erosion surfaces, flow inflation and explosivity of komatiites. Komatiites are inferred to propagate under both turbulent and laminar flow conditions (Huppert and Sparks, 1985a; Cas et al., 1999), whereby flow turbulence is used to explain thermal erosion contacts between flows (e.g., The Precambrian Earth: Tempos trod Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuncanu

278

Chapter 4: Precambrian Volcanism

Lateral termination of tube-shaped flow ?

4,,

-y

4

Single flow unit

.

, oO;-O, Oo*

Sheet flow ~\~11 \ 1

9

'1111, 9149 9149

~.', ' ' 9 1 4 99 1 4 9

..'~

9

.o

9 ..

9

9

9

,~ o o..OL

9

}

FLOW DIVISIONS A1 Chilledflow top A2 Small random spinifex A3 Large aligned platy spinifex B1 Foliated skeletal spinifex B2 Fine-to medium-grained cumulate B3 Knobbycumulate

I

, 25 metres

I

B4 Fine- to medium-grained cumulate

Fig. 4.3.1-1. Komatiite flows with A-B zones and subdivisions at Spinifex Ridge (Champagne et al., 2002). The well-defined distribution of A and B zones on this outcrop suggests a prominent sheet flow morphology rather than tube-shaped flows (see statistics from Table 4.3.1-1 ).

4.3. Komatiites

279

Williams et al., 1998). Recently, Houl6 et al. (2001) suggested at Alexo Mine thermomechanical erosion of a basal andesite unit eroded by a komatiite flow containing andesite blocks. Komatiites have been presented as thick, high volume flows, yet the low viscosity should favour a thin, highly fluid runout. Hon et al. (1994) demonstrated from modern pahoehoe sheet flows on Hawaii that thin flows could inflate significantly. Dann (2001) in turn used these modern analogies to explain how Barberton komatiites inflated and thickened due to on-going flow pulses. Evidence of flow inflation was suggested by the accumulation of multiple vesicle-rich layers (Self et al., 1998; Moore et al., 2000; Dann, 2001). Finally, low-viscosity high-temperature flows should not favour explosive fragmentation of komatiitic magma, yet komatiitic tuff and lapilli tuff deposits (Saverikko, 1985; Nisbet et al., 1993b, pp. 146-147) and vesicular komatiite forming pyroclastic deposits with armoured and accretionary lapilli (Schaefer and Morton, 1991; Lowe, 1994a) have been observed. Although most fragmentation processes are probably autoclastic, volatile-rich komatiites interacting with seawater in a shallow-water setting may produce subaqueous to subaerial Surtseyan-type eruptions with resultant small-volume pyroclastic deposits. Abitibi greenstone belt

Mapping of Abitibi (see also section 2.4) komatiites (Fig. 4.3.1-2) sheds light on compound flow organisation, internal textural zones and flow field geometry (Mueller et al., 1999; Houl6 et al., 2001; Champagne et al., 2002). Lateral and vertical variations of textural zones in komatiite flows are commonly discernible at the outcrop scale. Komatiites of the Lamotte-Vassan Formation (Spinifex Ridge) and komatiite-komatiitic basalts of the Stoughton Roquemaure Group are discussed here. Champagne et al. (2002), dividing Spinifex Ridge flows into sheet flow (generally > 30 m width) and tube-shaped komatiites (< 10 m width; Table 4.3.1-1), noticed that the distribution and thickness of A and B zones is correlative with flow geometry. Tube-shaped komatiites display a prominent B zone (A zone 24 cm vs B zone 62 cm), whereas the A zone is dominant in sheet flow komatiites (A zone 44 cm vs B zone 24 cm; Table 4.3.1-1) so that A/B ratios may help discern flow type if outcrop exposure is incomplete (Fig. 4.3.1-1). The tube-shaped flows (Fig. 4.3.1-3a) have low aspect ratios < 10 (av. 6:1 width vs height), whereas sheet flows (Figs. 4.3.1-3b, c) have high aspect ratios > 10 (av. 16:1). For comparison, mafic Archaean pillowed flows have aspect ratios of < 2 (Sanschagrin, 1982). Although welldeveloped at Spinifex Ridge (Fig. 4.3.1-3d), komatiites generally lack Az-A3 spinifex zones and well-defined B zones. Thickness of A-B textural zones or their absence is attributed to flow morphology, lava velocity, effusion rate, and water access into the flow. Thermal quenching appears more effective around tube-shaped komatiites than sheet flows, because of their smaller size, with seawater ingress along fractures at the top and side going deep into the tube. Spinifex growth occurs from the roof downwards and in sheet flows thermal contraction factures are generally restricted to the initial 10 cm of the roof. Once the crust is formed, an efficient insulation barrier is achieved and komatiite underflow may continue or stagnate. Olivine spinifex growth as described by Shore and Fowler (1999) ensues with skeletal downwards growth into a cooling medium. Locally, rip-ups of A-zone spinifex from the lava roof may be found in the B zone, indicating dynamic flow

280

Chapter 4: Precambrian Volcanism

Fig. 4.3.1-2. Abitibi greenstone belt with the various volcanic and sedimentary cycles spanning c. 65 My. Note the divisions into Northern and Southern Volcanic zones. Numbered black stars indicate the locations of: (1) the Noranda caldera of the Blake River Group (see Fig 4.6-1 for local geology); (2) the Hunter Mine and Stoughton-Roquemaure Groups with outlined area indicated in Figure 4.3.1-4; (3) La Motte-Vassan Formation with komatiites of Spinifex Ridge; and (4) komatiites of Munro Township and adjacent areas. and cooling conditions (Barnes, 1985). In small tube-shaped komatiites spinifex growth is less pronounced or inhibited because a thick glassy zone with polygonal fractures forms not only at the top, but also the base and sides, restricting spinifex growth to the central segment. In contrast to the rounded margins of tholeiitic pillowed basalts, tube-shaped komatiites show sharp, low-angle lateral flow terminations indicative of a very low viscosity (Fig. 4.3.1-3a). Closely-packed, tube-shaped komatiites are overlapping flows that compare favourably with lobate pahoehoe flows. The tubes have a low-arching roof, either caused by flow inflation or more viscous flow. Drainage cavities (Fig. 4.3.1-3e) and pockets of breccia in the centres of tubes occur at the same height as the low-angle lateral flow terminations. Cavities and breccia are suggestive of several pulses of lava transport. Inflation must have been a process facilitating tube arching and flow thickening. Spinifex textures did not form in these tubes and amygdules constitute < 1%. The thin sheet flows (Fig. 4.3.1-3c), characterised by their lateral continuity, display classic A 1-3-B 1-4 divisions and lack drainage cavities. Sharp lateral flow terminations are indistinguishable from those of tube-shaped komatiites. Considering the thin nature of these flows (Table 4.3.1-1), an absence of vesicularity and a dominant A zone flow inflation were insignificant. Champagne et al. (2002) mapped an erosion surface between sheet flows at Spinifex Ridge with a thick flow down-cutting into an underlying A zone segment and forming a small scour (Figs. 4.3.1-3b, c). A thermo-mechanical erosion surface is advocated, possibly resulting from multiple flow pulses under turbulent flow conditions.

4.3. Komatiites

281

Fig. 4.3.1-3. Flow morphology and characteristics of Spinifex Ridge, Lamotte-Vassan Formation. (a) Closely-packed tube-shaped komatiites (Ts Kom) with sharp low angle terminations (Sh T). Central segments may have in situ brecciation. Scale: pen = 13 cm. (b) Two sheet flows ( 1 and 2) marked by an erosive contact. A zone is eroded by overlying flow. Scale: field book = 20 cm. Note lateral continuity of thin sheet flow. (c) A series of well-defined sheet flows with a sharp erosive surface between 1 and 2. (d) Details of a sheet flow with A 1, A 2 and A 3 divisions: A 1 division is a chilled glassy flow top with thermal contraction fractures (TC Fr), whereas mm-scale spinifex is developed in A2, and large cm-scale radiating spinifex is formed in A 3. Scale: pen = 13 cm. (e) Tube-shaped komatiite flow with a drainage cavity (D Cav) and abundant thermal contraction fractures (TC Fr). Scale: pen = 13 cm. Note in situ brecciation below cavity.

Table 4.3.1 - 1. Characteristics of Spinifex Ridge sheet flow and tube-shaped komatiites (synthesised from Champagne et al., 2002) Komatiite How morphology and features Sheet How

Sheet How features Prominent A1, A2, A3, B2, B j , B4 zones; tabular flows with low-angle lateral terminations Tube-shaped features Prominent A1, Aj, B2, B4 zones; flat to arched tubes with low-angle lateral terminations

Size

Average, m

Maximum, m

Width (W) Thickeness (T) Aspect ratio W/T

17.87 I .09 16.39

> 34.83

Width (W) Thickeness (T) Aspect ratio WIT

5.37 0.85 6.37

A zone (T)

B zone (T)

A/B ratio

43.96 cm (average)

(average)

(average)

Minimum, m

8.42

Cavities and ioints Drainage cavities absent; polygonal joints prominent at roof and flow margin

3

23.65 cm (average)

(average)

(average)

Drainage cavities locally present; polygonal joints prominent in flow

%2

a b

8 2

$-

5'

4.3. Komatiites

283

A thin, possibly recrystallised film at the flow contact may be due to reheating from the overlying flow (e.g., Burkhard, 2003). The 5-20 m-thick vertical (and lateral) facies architecture of sheet flow to tube-shaped komatiites indicates the change in lava flow dynamics, possibly related to a decrease in lava supply over a subdued topography (< 1-3 degrees). The subaqueous komatiite sheet flows delivered lava to the interconnected tube-shaped komatiites at the margins and flow front, as documented in pahoehoe sheet flows (Hon et al., 1994; Crown and Baloga, 1998). The closely packed tube-shaped flows (Fig. 4.3.1-3a) may be analogues to coalescing pahoehoe toes (e.g., Hon et al., 1994; Crown and Baloga, 1998). Kilauea sheet flows advanced at average velocities of 0.01-0.05 km hr-1 (maximum velocity 0.1-0.2 km hr -1) and in insulated tubes at 3-6 km hr-1 (Hon et al., 1994), so that if komatiites are comparable to pahoehoe flows, then propagation of flows should be comparable or possibly an order of magnitude higher considering their lower viscosity and higher temperature. Thick massive flows may be analogues of tube pahoehoe flows (Peterson et al., 1994) or master tubes, produced by flow inflation (Self et al., 1998). Once a flow network was established on the Archaean ocean floor or seamount, komatiite lava would have been transported to the flow front in an efficient manner with hardly any loss of temperature, as shown for basalts (loss of 0.5-1 ~ per km; Hon et al., 1994). The Stoughton-Roquemaure komatiites and komatiitic basalts (Figs. 4.3.1-4a, b), with large- to small-scale tube-shaped flows (Fig. 4.3.1-5a) and pillowed lava flows, complement the observations at Spinifex Ridge by displaying the transition from master to distributary tubes. Effusive sequences, c. 50-150 m thick, are divided into (1) complex master tubes, > 20 m wide and up to 5 m thick, (2) secondary distributary tubes, 5-20 m wide, (3) branching pillows and pillow tubes < 5 m wide, (4a) pillow fragment breccia, and (4b) hyaloclastite and pillow rind breccia. The striking difference from the Spinifex Ridge lava flows is the ubiquitous rounded, bulbous lateral margins of flows, suggesting higher viscosities. Master tubes are complex domal or flat-topped structures, massive to columnar-jointed (Fig. 4.3.1-5a) that locally display pillows and pillow selvages at the margins. Large pillows at the master tube margin indicate branching into distributary tubes (Fig. 4.3.1-5b). Thin sheet flows as observed at Spinifex Ridge did not develop. Numerous thermal fractures with chilled margins perpendicular to the cooling front commonly traverse thick flow units, possibly explaining the absence of large spinifex. The distributary tubes and pillows are flat to bulbous structures that on three-dimensional exposures display pillow budding and branching (Fig. 4.3.1-5c), longitudinal and transverse spreading cracks, and abundant polygonal jointing (Dostal and Mueller, 1997). Numerous tubes display multiple chilled margins possibly suggestive of flow inflation. The pillow fragment and pillow rind breccia located at or near the top of these effusive cycles, are the result of autoclastic process although implosions or even flow inflation may well have contributed to komatiite flow fragmentation. Varioles are common in the pillowed segment of the flow unit (Fig. 4.3.1-5d). Pillow rind breccias with multiple chilled pillow margins may in fact be a direct result of localised overpressure resulting in pillow tube breakout, comparable to pahoehoe flow front breakouts (Hon et al., 1994; Self et al., 1998).

284

Chapter 4: Precambrian Volcanism

Fig. 4.3.1-4. (a) Geology of the Lake Abitibi area, Abitibi greenstone belt. (b) Composite stratigraphy of the felsic Hunter Mine caldera and overlying Stoughton-Roquemaure komatiite~ (modified from Mueller and Mortensen, 2002).

4.3. Komatiites

285

Fig. 4.3.1-5. Characteristics of Stoughton-Roquemaure flows. (a) Contact between two effusive sequences. Pillowed flow units overlain by a columnar jointed master tube. Scale: book - 20 cm. (b) A master tube with pillows at the margin of the flow suggesting separation into secondary distributor tubes. (c) Branching of pillow tube with well-defined surface features. Scale: pen = 13 cm. (d) Pillow tube budding out of a longitudinal spreading crack. Scale: pen = 13 cm. (e) Varioles in pillow breccia of komatiitic basalt. Scale: pen tip = 2.5 cm.

286

Chapter 4: Precambrian Volcanism

Fig. 4.3.1-6. General geology of the Barberton greenstone belt (Swaziland) with study area in location A (Dann, 2001).

Barberton greenstone belt

Komatiites of the Barberton greenstone belt (Fig. 4.3.1-6) have been a centre of debate because the extrusive origin of these komatiites had been questioned (Parman et al., 1997). Facies mapping by Dann (2000, 2001) at the classic locality showed that these are lava flows. The Barberton komatiites, complementing the compound flow architecture at Spinifex Ridge, allow for the characterisation of the large-scale flow field geometry. The 1.7 km thick Lower Komati Formation, is composed of interlayered komatiites and komatiitic basalts (Fig. 4.3.1-7a) with 12-270 m-thick komatiite flow units, which are: (1) massive (61 -+- 9%), (2) spinifex-textured (37 -+- 10%) and (3) vesicular (2 + 0.5%). The flow units, traceable for 11 km, are interpreted to represent individual flow fields (Fig. 4.3.1-7b). Massive units without textural zoning or chilled glassy flow tops display local multiple cooling units suggesting multiple sheet flows. The massive units are overlain by sequences (< 100 m thick) of multiple spinifex flows with the classical A-B zones. The spinifex flows occur as sheets (< 8 m thick) and thinner lenticular units with pointed lateral margins, are interpreted as sheet flows and flow lobes, respectively (Dann, 2000). The spinifex-overmassive zoning of komatiite flow fields with intervening komatiitic basalts is repeated five times (Fig. 4.3. l-7a). Apparently, each effusive event started with the thick flows of olivinephyric komatiite and ended with thin flows of spinifex komatiite with reduced phenocryst loads, possible due to waning effusion rates.

4.3. Komatiites

287

Fig. 4.3.1-7. (a) General lava flow stratigraphy in study area of the Lower Komati Formation. (b) Overall geometry of a komatiite flow field with compound spinifex flows (Dann, 2001). Outcrop numbers such as K2-15 indicate area mapped in detail and used to describe flow inflation features in Fig. 4.3.1-8.

288

Chapter 4: Precambrian Volcanism

The minor vesicular komatiites shed new insight into behaviour of komatiite flows and occur at the boundary between massive and spinifex flows within flow fields (Dann, 2001). Vesicles are concentrated in the upper carapace of thick flows, with olivine-cumulate basal zones. Outcrop evidence of synvolcanic rotations and magma intrusion into the upper carapace, and a changing topography of the flow surface with internal structures and zoning are suggestive of inflation. Flow inflation structures had previously only been identified in subaerial pahoehoe (Hon et al., 1998; Self et al., 1998) and subaqueous pahoehoe (Umino et al., 2000) flow fields. To elucidate, the upper vesicular carapace, 15 m thick in one komatiite unit, underwent block rotations during the influx of magma that crystallised spinifextextured komatiite (Fig. 4.3.1-8a). The resulting domed structure, forming a tumulus, added c. 20 m of surface relief to the flow. The structure is comparable to the upper surface of smaller tumuli observed on the Loihi Seamount, Hawaii (Umino et al., 2000). Inflation with renewed flow breakout may also occur at the margin of lava rises or pits (Fig. 4.3.1-8b, part 1). The carapace displays fracturing and block rotations with intrusion by a network of dykes, some feeding new flows (Fig. 4.3. l-8b, part 2). The inflating segment causes a down flow depression in the flow surface filled by a thick volcanic breccia. This breccia includes fragments of a spinifex flow that may have been a breakout fed by dykes intruding the hinge area (Fig. 4.3. l-8b, part 3). Further along strike the carapace of the same vesicular komatiite (Fig 4.3.1-8c, part 1) was faulted, forming a shallow graben, intruded by komatiite dykes, and flooded by a massive flow (Fig 4.3.1-8c, part 2). The mobility of the cumulate interior of komatiite flows is well demonstrated by continued flooding and thus inflation of this flow top and the networks oi: dykes (Fig 4.3.1-8c, part 3). The vesicular units are direct evidence for devolatilisation of komatiite lava. Vesicles occur preferentially in the upper carapace of these flows, just as they do in basaltic flows. However, at Kambalda, vesicle layers occur within the basal cumulate zone and may be linked to particularly thick flows (Moore et al., 2000). In contrast, Dann (2001) suggested that such occurrences could represent foundering of the vesicular carapaces by mobilisation of the interior cumulate zone, a process well illustrated in outcrop, but not resolvable from drill core. Alternating layers of vesicular and spinifex komatiites in the upper carapace of the Barberton vesicular komatiites may represent episodes of lava influx during an early stage of inflation (Dann, 2001), processes known from subaerial, pahoehoe flows (Self et al., 1998). Textures of preserved bubble coalescence indicate that bubbles rose and accumulated beneath a downwards crystallising roof. Because volatile saturation is linked to ongoing crystallisation, vesicle-free spinifex zones may record the input of fresh lava

Opposite: Fig. 4.3.1-8. Flow features and inflation characteristics of Barberton greenstone belt komatiites (from Dann, 2001). (a) Formation of a large komatiite tumulus via numerous lava pulses. (b) Development of a lava rise or possible lava pond with flow breeching at the margin. (c) Komatiite flooding of a graben collapse structure. Note that synvolcanic fractures are an integral component of flow inflation or collapse of the roof of lava flows or tubes. These fractures served as pathways for magma/lava and are commonly masked by dyke emplacement.

4.3. Komatiites

289

Chapter 4: Precambrian Volcanism

290

and the overpressure that accompanies inflation. In thick tumuli, textures are particularly coarse-grained, preserving segregation vesicles that record multiple events of increasing internal pressure, possibly related to inflation. It should be noted that although the vesicular komatiites are unique for their high concentrations of vesicles, the spinifex komatiites also have vesicles, particularly along the upper glassy flow tops. To summarise, studies in Canada and South Africa show that simple flow models with channel and levee deposits require reassessment. The Spinifex Ridge and StoughtonRoquemaure compound flows represent segments of large-scale Barberton-type flow fields and compare favourably to Hawaiian pahoehoe flows. Because the Archaean strata are steeply dipping, outcrop zones represent a cross-section of inferred compound flows or flow fields. Volcanic facies mapping in conjunction with new observations from pahoehoe fields in Hawaii, can explain how komatiites inflated, how komatiite tumuli formed and how superposed flows may simply have been results of flow breakout. In the subaqueous setting, collapse structures or foundering of lava tubes is a common occurrence and Archaean komatiite flows should be no exception. The graben structure may be generated possibly by large scale master tube implosion, or collapse of a drained tube due to hydrostatic pressures in a deep-water setting. The Abitibi komatiitic basalts and komatiite flows define compound flows from smaller flow volumes and show the complex lateral and vertical changes in flow geometry. The work of Dann (2001) may suggest relative flow field locations: (1) medial to distal flow top and flow front breakouts, (2) proximal to medial tumuli, and (3) proximal to medial lava rises or ponds with lava breakout, or collapse and subsequent lava flooding. 4.3.2

Komatiite Geochemistry

J. Dostal and W.U. Mueller

The geochemical characteristics of komatiites (Table 4.3.2-1) are based mainly on the Barberton and Abitibi greenstone belts because of excellent stratigraphic control. Komatiites were defined as rocks derived from liquids with more than 18 wt.% MgO (Arndt and Nisbet, 1982b; see also section 4.3.3 for definition), and display elevated Ni contents, with very low TiO2, Na20, K20 and incompatible trace element abundances. Most of the chemical variations observed can be accounted for by crystallisation and accumulation of olivine. Olivine-rich cumulate rocks (B zone) at the base of flows contain 30-40% MgO, whereas the spinifex-textured upper parts (A zone) have 20-28% MgO (Smith and Erlank, 1982; Arndt et al., 1997). Glassy margins of the komatiitic flows with a minor amount of olivine phenocrysts (Fo94) and aphyric flows contain c. 28-30 wt.% MgO (Barnes et al., 1983; Arndt, 1986), suggesting this composition represents the magmatic liquid. An almost twofold increase in TiO2, A1203 and CaO from komatiites with 35-20 wt.% MgO can be accounted for by 30-50% crystallisation of olivine (Smith and Erlank, 1982). An evaluation of komatiite petrogenesis therefore requires the use of element ratios, and particularly ratios of incompatible elements not modified by olivine crystallisation. The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

4.3. Komatiites

291

Table 4.3.2-1. Major and trace elements of selected Archaean komatiites Sample

Al-undepleted komatiites

Al-depleted komatiites

SiO2 (wt.%) TiO 2 A1203 Fe20~ MnO MgO CaO Na20 K20 P20 5

M 664 45.82 0.36 6.39 10.99 0.18 29.59 6.20 0.35 0.10 0.02

5019 43.57 0.28 2.94 15.46 0.36 34.96 2.34 0.02 0.02 0.05

Total LOI Cr (ppm) Ni Rb Ba Sr Nb Hf Zr

M 712 Z2 478 4 4 . 4 1 49.65 48.39 0.21 0.43 0.33 3.74 9.01 7.04 8.78 1 3 . 0 4 10.63 0.16 0.21 0.19 39.58 1 7 . 5 9 24.48 2.82 8.79 8.61 0.27 1.19 0.26 0.01 0.06 0.03 0.02 0.03 0.04

100.00

100.00

100.00

100.00

100.00

7.30

10.15

2.30

4.90

1.78

2500.00 1550.00 5.85 14.00 23.10 0.55 0.45 16.00

1948.00 2506.00 0.99 1.50 4.00 0.35 0.23 8.80

1.30 12.30 33.90 0.77 0.57 19.90

0.90 2.70 5.10 0.69 0.59 20.90

B 14 47.69 0.43 4.32 12.86 0.19 26.13 8.24 0.10 0.01 0.03

B 15 46.10 0.30 2.98 11.78 0.20 33.92 4.70 0.00 0.00 0.02

95-12 47.72 0.73 8.23 13.52 0.19 18.98 10.19 0.38 0.01 0.05

100.00

100.00

100.00

5.09

8.76

4.60

5200.00 2427.00 1640.00 1.40 1.16 3.00 7.00 5.30 34.30 0.81 1:48 0.49 0.62 18.20 23.10

1942.00 1876.00 720.00 0.29 1.00 13.30 22.00 20.10 11.00 0.83 2.10 0.46 1.14 17.20 44.00

Komatiites constitute a small proportion of most greenstone belts (< 5% in the Abitibi belt; Sproule et al., 2002) and are associated with tholeiitic to komatiitic basalts (12-18 wt.%; Arndt and Nisbet, 1982a). Komatiitic basalts have similar flow facies and textures to komatiites, with pyroxene as the cumulate or skeletal spinifex mineral. Pyroxene spinifex-textured komatiitic basalts with MgO and Ni (typically < 1000 ppm; Barnes, 1983; Wyman et al., 1999b) have values lower than komatiites. Major and trace element geochemistry permitted the distinction into aluminium (Al)-depleted and aluminium (A1)-undepleted komatiites (Nesbitt et al., 1979; Smith and Erlank, 1982; Jahn et al., 1982; Figs. 4.3.2-la, b, c), but several new subtypes have since been identified. Al-depleted versus Al-undepleted komatiites Komatiites are divided into: (1) Al-depleted flows derived from greater depths with a lower degree of melting, and (2) Al-undepleted flows originating from shallower levels with a higher degree of melting. Al-undepleted or Munro-type komatiites feature (i) near-chondritic ratios of AlzO3/TiO2 (c. 20; Fig. 4.3.2-1b) and CaO/AI203 (c. 1; due to Ca mobility, ratio may not be reliable), and (ii) flat heavy rare earth element (REE) patterns with (Gd/Yb)n c. 1, and with flat to depleted light REE (LREE) patterns comparable to those of recent N-MORB, although concentrations in komatiites are significantly

292

Chapter 4: Precambrian Volcanism

Table 4.3.2- l (continued). Sample

Al-undepleted komatiites

Y Th La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

M 664 8.54 0.04 0.55 1.62 0.27 1.55 0.64 0.24 0.95 0.15 1.20 0.26 0.78 0.11 0.77 0.11

M 712 3.67 0.03 0.38 1.17 0.19 1.02 0.40 0.16 0.62 0.11 0.77 0.17 0.48 0.07 0.50 0.08

Z2 10.40 0.10 1.00 2.57 0.44 2.14 0.79 0.29 1.03 0.18 1.34 0.34 0.98 0.15 0.89 0.14

478 9.80 0.05 0.65 2.12 0.39 2.06 0.81 0.25 0.98 0.18 1.22 0.31 0.92 0.14 0.84 0.13

5019 4.40 0.11 1.82 4.00 0.54 2.30 0.62 0.23 0.78 0.13 0.81 0.18 0.51 0.08 0.43 0.06

Al-depleted komatiites B 14 8.71 0.15 1.67 5.22 0.75 4.02 1.26 0.35 1.42 0.25 1.55 0.33 0.92 0.13 0.83 0.13

B 15 5.93 0.08 1.21 3.62 0.46 2.56 0.83 0.31 1.07 0.19 1.18 0.25 0.68 0.10 0.61 0.11

95-12 l l.00 0.22 1.84 5.85 0.97 5.01 1.67 0.60 2.05 0.34 2.21 0.42 1.20 0.17 1.20 0.17

AI203/TiO2 (Gd/Yb)n

18 1.0

18 1.0

21 0.96

21 0.97

11 1.5

10 1.4

10 1.5

11 1.4

Al-undepleted samples: M 664---olivine spinifex lava, c. 2.714 Ga komatiite flow, Alexo, Abitibi greenstone belt, Canada (Lahaye et al., 1995; Jochum et al., 1990); M 712--olivine cumulate, c. 2.714 Ga komatiite flow, Alexo, Abitibi greenstone belt, Canada (Lahaye et al., 1995); Z 2--c. 2.7 Ga old komatiitic basalt, Zwishavane, Belingwe greenstone belt, Zimbabwe (Jochum et al., 1990); 478--c. 2.7 Ga komatiite from volcanic sequence at Kambalda, Western Australia (Jochum et al., 1990; Lesher and Arndt, 1995); Al-depleted samples. 5019-c. 3.5 Ga komatiite, Theespruit Formation, Onverwacht Group, Barberton greenstone belt, South Africa (Jahn et al., 1982; Jochum et al., 1990); B 14---c. 3.5 Ga olivine spinifex lava, komatiite flow, Komati Formation, Barberton greenstone belt, South Africa (Lahaye et al., 1995); B 15---c. 3.5 Ga olivine cumulate, komatiite flow, Komati Formation, Barberton greenstone belt, South Africa (Lahaye et al., 1995); 95-12--c. 2.724 Ga pyroxene spinifex lava, Stoughton-Roquemaure Group, Abitibi greenstone belt, Canada (Dostal and Mueller, 1997): n-chondrite normalised.

Fig. 4.3.2-1. (a) Primitive mantle normalised trace element abundances of komatiites with I, Al-depleted and II, Al-undepleted komatiites: normalising values after Sun and McDonough (1989). See Table 4.3.2-1. (b) Variations of MgO (wt.%) versus AI203/TiO2, A1203 (wt.%), TiO2 (wt.%), (Gd/Yb)n, Gdn and Lan in Al-undepleted komatiitic (o) and Al-depleted komatiitic (+) rocks; Al-undepleted komatiites: Abitibi (Arndt, 1986; Jochum et al., 1990; Lahaye et al., 1995; Dostal and Mueller, 1997; Sproule et al., 2002; Wyman et al., 1999b), Belingwe (Jochum et al., 1990) and Kambalda (Jochum et al., 1990; Lesher and Arndt, 1995) greenstone belts; Al-depleted komatiites: Barberton (Jahn et al., 1982; Smith and Erlank, 1982; Jochum et al., 1990; Lahaye et al., 1995; Byerly, 1999) and Abitibi (Dostal and Mueller, 1997; Sproule et al., 2002) greenstone belts; n-chondrite normalised. (c) Chondrite-normalised REE abundances in Al-depleted and Al-undepleted komatiites, averaged from samples in Table 4.3.2-1, are compared to N-type and E-type MORBs (examples and normalising values; Sun and McDonough, 1989).

293

4.3. Komatiites

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294

Chapter 4: Precambrian Volcanism

lower than in MORB (Fig. 4.3.2-lc). Al-depleted or Barberton-type komatiites have lower A1203 contents with AIzO3/TiO2 ratios of c. 10 (Fig. 4.3.2-1b), but higher CaO/AI203, more than 1, with values close to E-MORB. Likewise, these komatiites have typically higher contents of strongly incompatible trace elements (Th, LREE) relative to Al-undepleted types (Fig. 4.3.2-1 a). Al-undepleted komatiites are prominent in young, 2.7 Ga Archaean belts, which include the Abitibi (Canada) (see also section 2.4) and Belingwe (Zimbabwe) greenstone belts, and sparse in pre-3.0 Ga belts. In contrast, Al-depleted komatiites are common in the 3.5-3.0 Ga Barberton greenstone belt, South Africa and greenstone belts of the Pilbara craton, Australia (Herzberg, 1995) (see also sections 2.5 to 2.7). The combination of Al-depleted and -undepleted types is rare but has been documented in segments of the Abitibi greenstone belt (Cattell and Arndt, 1987; Dostal and Mueller, 1997). Furthermore, Jahn et al. (1982) identified rare Al-enriched komatiites, with high AlzO3/TiO2 and low (Gd/Yb)n ratios that are more abundant in pre-3.0 Ga belts (Arndt, 1994). Barnes and Often (1990) described Ti-rich komatiites from Norway, whereas Sproule et al. (2002) recorded Ti-enriched (with AlzO3/TiO2 < 15 and [Gd/Yb]n > 1.2) and Ti-depleted komatiites (with AlzO3/TiO2 c. 25-35 and [Gd/Yb]n c. 0.6-0.8) in the Abitibi and Swayze greenstone belts.

Komatiite isotope and trace element signature The Nd isotopic compositions of komatiites and associated tholeiitic basalts are probably best represented by the data set from the Abitibi greenstone belt where the basalts and komatiites as well as their constituent pyroxenes give, typically, ENd values of c. +2.5 to +3.8 (Machado et al., 1986; Lahaye et al., 1995). Some of the Nd isotope values for komatiites are higher, suggesting that their initial ENd values were probably affected by REE mobility during alteration (e.g., Lahaye et al., 1995). Similarly for the Sr isotopic data, in order to avoid the effects of alteration on the system, Hart and Brooks (1977) and Machado et al. (1986) analysed pyroxenes from mafic and ultramafic rocks of the Abitibi belt and obtained an average initial 87Sr/86Sr ratio of c. 0.701. Despite these uncertainties, isotope data imply depletion in the more incompatible elements relative to primitive mantle. Komatiitic magma may be contaminated by continental crust during ascent (Huppert and Sparks, 1985b; Jochum et al., 1990), but also by thick sedimentary sequences common to greenstone belts. Crustal contamination would lead to enrichment of light REE and Th relative to heavy REE and high field-strength elements, particularly Nb and Ta (Jochum et al., 1990). The contamination produces negative anomalies for Nb, Ta and Ti on the mantle normalised trace element patterns. Rocks contaminated by older continental crust also give lower ENd values (Arndt et al., 1997; Faure, 2001). Because of the large differences in Th and light REE concentrations and in Nb/Th and Nb/light REE ratios between komatiites and typical continental crust, these ratios, particularly in conjunction with Nd isotopic data, are very sensitive indicators of crustal contamination (Jochum et al., 1990). Abitibi greenstone belt, Canada The c. 2.7 Ga Abitibi greenstone belt (Fig. 4.3.1-2), the largest coherent Archaean supracrustal sequence in the world, has a well established stratigraphy, in which komatiite

4.3. Komatiites

295

successions evolved over c. 20 My (2724-2703 Ma) (see also section 2.4). Abitibi komatiites are linked to volcanic cycle 1 (2735-2720 Ma), volcanic cycle 2 (2720-2705 Ma), and volcanic cycle 3 (2705-2697 Ma) with each cycle spanning 8-15 My (Mueller et al., 1996; Mueller and Mortensen, 2002). There is an older, poorly documented cycle with komatiites, the Pacaud assemblage (2750-2735 Ma; Ayer et al., 2002; Sproule et al., 2002) but these rocks are part of the Swayze greenstone belt. The Abitibi greenstone belt formed in an oceanic setting and displays a continuum of events of arc formation, arc evolution, arc-arc collision and arc fragmentation (section 2.4). Abitibi tectonic evolution is intimately associated with komatiites, and their distribution is focused along two terrane zippers, the E-trending northern Destor-Porcupine Manneville and southern Cadillac-Kirkland Lake fault zones (Fig. 4.3.1-2). This relationship is not fortuitous, and may show how the inferred plume created zones of weakness that were subsequently exploited during shortening to form thrust and strike-slip zones. Volcanic cycle 1 in the Lake Abitibi area (Fig. 4.3.1-4a) has the 0.2-2 km thick Stoughton-Roquemaure Group (SRG) with tholeiitic basalt, komatiitic basalt and komatiite (Dostal and Mueller, 1997) conformably overlying the calc-alkaline 2734-2728 Ma Hunter Mine Group (Fig. 4.3.1-4b). SRG komatiites are critical to understanding the petrogenesis of komatiites, as Al-depleted and Al-undepleted komatiites are stratigraphically superposed and interlayered at the large scale to form 50-150 m-thick eruptive sequences (Dostal and Mueller, 2002). Both komatiite types and tholeiites have overlapping positive end values and basal SRG tholeiitic basalts resemble N-MORB. The komatiitic basalts, with low AIzO3/TiO2 ratios (c. 10) and fractionated heavy REE with (Gd/Yb)n of 1.4-1.8, correspond to Al-depleted komatiites, whereas komatiites in s e n s u stricto are all Al-undepleted with high AIzO3/TiO2 (c. 20) and unfractionated heavy REE with (Gd/Yb)n c. 1.1-1.2. The SRG stratigraphy shows Al-depleted komatiite overlying MORB-like basalts at the base of the komatiitic sequence followed by MORB-like basalts and Al-undepleted komatiites. The Stoughton-Roquemaure assemblage in Ontario (volcanic cycle 2) has transitional rocks in terms of A1203/TiO2 ratios (Al-depleted and Al-undepleted komatiites) but Al-depleted are prevalent (Sproule et al., 2002). The 2718-2710 Ma Kidd-Munro assemblage, Ontario (volcanic cycle 2) includes the classic 1000 m thick Munro komatiite flows (Arndt et al., 1977), which contain Al-undepleted komatiites with local Al-depleted komatiites (Cattell and Arndt, 1987). Spinifex Ridge of the La Motte-Vassan Formation (volcanic cycle 2) is Al-undepleted (Champagne et al., 2002). The Tisdale assemblage, Ontario (volcanic cycle 3) contains Al-undepleted komatiites with rare Al-depleted type komatiites (Sproule et al. 2002), whereas the 2703 Ma Jacola Formation with AIzO3/TiO2 of c. 14 is a Ti-enriched komatiite variety and is associated with inferred arc rocks (Champagne et al., 2002). Al-undepleted komatiites are the most common variety in the Abitibi greenstone belt. Komatiite sequences of all volcanic cycles display a wide range of MgO contents, from komatiitic basalts to cumulate with > 32 wt.% MgO. Volcanic cycle 1 has the largest proportion of komatiitic basalts relative to komatiites, as well as the largest abundance of Al-depleted rocks. Volcanic cycles 2 and 3 are dominated by the Al-undepleted variety,

296

Chapter 4: Precambrian Volcanism

and some are enriched in highly incompatible trace elements, and have low Nb/Th and high Th/Sm ratios suggestive of a crustal signature (Sproule et al., 2002).

Barberton greenstone belt, South Africa The Barberton greenstone belt (see also section 1.3) is divided into the prominent volcanic 3.3-3.5 Ga Onverwacht Group overlain by sedimentary 3.2-3.3 Ga Fig Tree and Moodies Groups (Fig. 4.3.1-6; Anhaeusser, 1971; Armstrong et al., 1990; de Wit et al., 1992; Lowe and Byerly, 1999b). The belt has undergone significant folding and thrusting so that age determinations have been significant in unravelling the stratigraphy. Komatiites and komatiitic basalts occur throughout the Onverwacht Group. The 3.48 Ga Komati Formation, the type locality, is a sequence of interlayered komatiites and komatiitic basalts. Komatiites and komatiitic basalts also occur in the overlying 3.47 Ga Hooggenoeg Formation, and the Mendon and correlative Weltevreden Formations. Tholeiitic basalts occur in the 3.47 Ga Hooggenoeg and 3.4-3.3 Ga Kromberg Formations. The Kromberg Formation only has thick sills of ultramafic composition. Smith and Erlank (1982) and Lahaye et al. (1995) identified a dominant Al-depleted type (e.g., Komati Fm.; AlzO3/TiO2 c. 10; Fig. 4.3.2-1b) and a minor Al-undepleted type komatiite (e.g., Mendon Fm.; AlzO3/TiO2 c. 20). Al-depleted komatiites have significantly lower A1203 and slightly higher TiO2 contents compared to Al-undepleted counterparts, but both have similar CaO contents for comparable MgO values. The Al-depleted komatiites have high CaO/AI203 (> 1; mean 1.33) and high (Gd/Yb)n, > 1 (mean 1.4; Jahn et al., 1982). The LREE patterns vary and in some rocks LREE are depleted; in other rocks they are flat or enriched relative to heavy REE. Al-undepleted komatiites have (Gd/Yb)n of c. 1 (Jahn et al., 1982), and inter-element ratios show no systematic variations with MgO contents. Smith and Erlank (1982) argued that variations within Al-undepleted komatiites can be explained by olivine fractionation, but Al-depleted komatiites require an additional phase. Locally, rare Al-enriched komatiites with heavy REE enrichment of (Gd/Yb)n < 1 and light REE enrichment with (La/Sm)n > 1, low CaO/AI203 (< 1; mean 0.6), and AI203/TiO2 of c. 40 were also recorded (Jahn et al., 1982). The Komati Formation (Fig. 4.3.1-6) is composed of both Al-depleted and minor Al-undepleted lavas (Smith et al., 1980; Smith and Erlank, 1982), whereas the overlying Hooggenoeg Formation contains only Al-undepleted komatiitic rocks (Williams and Furnell, 1979; Byerly, 1999). The Kromberg Formation includes minor Al-depleted and rare Al-undepleted komatiites that are intercalated with abundant tholeiitic basalts (Byerly, 1999; Vennemann and Smith, 1999). The Sandspruit Formation of uncertain stratigraphic position contains Al-depleted komatiites with AI203/TiO2 of c. 8 (Viljoen and Viljoen, 1969a; Jahn et al., 1982; Byerly, 1999) and the Mendon Formation has Al-depleted komatiites with AlzO3/TiO2 of c. 10 (Lahaye et al., 1995). The Weltevreden Formation contains abundant Al-undepleted and minor Al-depleted and Al-enriched komatiites (Anhaeusser, 1985; Byerly, 1999). Gruau et al. (1990) reported eyd values of c. 0 for the Barberton komatiites and tholeiitic basalts, whereas Lahaye et al. (1995) obtained a range of +0.6 to -+-2 for whole rocks but +2.3 for primary pyroxene from komatiitic basalts of the Weltevreden Formation.

4.3. Komatiites

297

Petrogenesis Komatiites are interpreted as products of anhydrous mantle melting (Arndt et al., 1998) and have very high volatile-free liquidus temperatures, particularly in comparison with basalts. Alternatively, they have been suggested to be hydrous melts (Parman et al., 1997; Asahara et al., 1998) (see also section 3.6). Assuming dry melting, experimental data suggest that komatiitic liquids are generated by adiabatic decompression melting of Archaean mantle at depths ranging from about 70-270 km (2-9 GPa; Herzberg and O'Hara, 1998). These melts originated from mantle plumes (Campbell et al., 1989). The low concentrations of incompatible trace elements and their ratios in both komatiitic types are, in general, consistent with their derivation from sources slightly depleted in incompatible elements relative to a primitive mantle composition. The high contents of MgO (up to 30 wt.%) coupled with low abundances of incompatible trace elements in non-cumulate komatiites suggest that the komatiitic magma was generated by a high degree of partial melting (Nesbitt et al., 1979). Most major element variations of komatiites are due to fractional crystallisation and/or accumulation of olivine + / - chromite. Fractionation could have reached about 50% to produce komatiitic basalts with about 12 wt.% MgO. Differences between Al-depleted and Al-undepleted komatiites cannot be accounted for by low pressure fractional crystallisation dominated by olivine and/or pyroxene, although both are important phenocryst phases in komatiites. Olivine incorporates limited amounts of elements such as A1, Ti and Ca as well as most incompatible trace elements. Pyroxene fractionation does not significantly modify the A1203/TiO2 ratio but it changes the A1/Sc ratio, which has, usually, near-chondritic values in many komatiites (Byerly, 1999). Likewise, both komatiite types have overlapping Ca contents and thus clinopyroxene fractionation cannot readily explain the differences. Herzberg (1995) and Arndt et al. (1997) explained the differences between Al-depleted and Al-undepleted komatiites as related to the role of garnet in the source. Al-undepleted komatiites have fiat HREE patterns, contain relatively uniform A1203/TiO2 (Figs. 4.3.2-1a, b) and have several refractory lithophile element (A1, Ca, Ti, Sc, Zr, Y) ratios (e.g., A1/Sc) with values close to those of chondrites. This suggests that garnet and clinopyroxene, which can fractionate these elements, were incorporated into the liquid during melting. The ratios would not be uniform and chondritic if garnet and/or clinopyroxene remained in the residue. Al-undepleted komatiites were probably produced by melting of a garnet peridotite, leaving only olivine ( + / - or orthopyroxene) in the residue, probably in the pressure range of 3-5 GPa (Herzberg and O'Hara, 1998). Alternatively, Al-undepleted komatiites may be generated by melting of a garnet-free source. Al-depleted komatiites have lower A1203/TiO2 (Figs. 4.3.2-1a, b) and near-chondritic CaO/TiO2 ratios indicating either a source depleted in A1 or that some mineral, such as garnet, remained in the melting residue. The latter is supported by the fractionated heavy REE pattern of the Al-depleted komatiites, which is attributed to the presence of garnet in the melting residue (Ohtani et al., 1989; Herzberg, 1995). Blichert-Toft and Arndt (1999) inferred from Lu-Hf isotopic data that Barberton Al-depleted komatiites were derived from a garnet-bearing source and that their residuum was garnet-rich. Experimental data indicate that Al-depleted komatiites were generated at 6-9 GPa leaving an olivine-garnet-

Chapter 4: P r e c a m b r i a n Volcanism

298

clinopyroxene residue (Herzberg and O'Hara, 1998; Walter, 1998). Both the Al-depleted and Al-undepleted komatiites were probably generated by a high degree of melting from comparable sources. A mantle plume (sections 3.2 and 3.3) composed of garnet peridotite could be the source. Subsequently the komatiitic magma underwent lower pressure fractional crystallisation dominated by olivine. 4.3.3

Textures in Komatiites and Variolitic Basalts

N.T. Arndt and A.D. Fowler

Komatiites and variolitic basalts are widespread in Archaean volcanic sequences. Spinifex is a spectacular bladed olivine or pyroxene texture that characterises komatiite and varioles are cm-scale leucocratic globular structures abundant in many Archaean basalts. These striking textures provide valuable information about conditions during emplacement of the host magmas, particularly about how magmas crystallised. Spinifex textures consisting of arrays of numerous subparallel olivine blades extend tens of centimetres to metres below flow tops. The habit of the strongly anisotropic crystals is suggestive of fast cooling near the flow margin, yet the crystals form deep within the flows. The large temperature difference between solidus and liquidus of komatiites (300--400~ provides a partial explanation. In addition, the blades are so orientated that their fastest growing faces were normal to the cooling contacts, suggesting growth in a strong chemical-potential gradient, in part created by the crystals, as they modified the composition and temperature of the liquid from which they crystallised. The term variole is useful in the field, particularly during the study of Archaean rocks because textures are often blurred by alteration. Varioles result either from blotchy alteration or magma mingling, or represent a form of plagioclase spherulite. The internal organisation and geochemistry is incompatible with the concept of quenched immiscible liquids (i.e., Grlinas et al., 1977). Most examples of varioles from the SW Abitibi greenstone belt are plagioclase spherulites. These are found within aphyric tholeiitic basalts, suggesting the magmas were superheated during eruption. The presence of komatiites and the widespread occurrence of plagioclase spherulitic basalts are indicative of unique Archaean thermal conditions. Variolites

Initially, varioles (Fig. 4.3.3-1 a) were defined as spherical masses, which may or may not be spherulites, found on the weathering surfaces of basalts and diabases (e.g., Lofgren et al., 1974). Commonly, varioles are considered to be the mafic counterparts of spherulites found in felsic volcanic rocks. Bates and Jackson (1987) give the following definition: "A pea-size spherule usually composed of radiating crystals of plagioclase or pyroxene. This term is generally applied only to such spherical bodies in basic igneous rocks". Fowler et al. (2002) recommend the initial usage because several different mechanisms give rise to cm-scale globular structures. Spherulites (Fig. 4.3.3-1b) are densely packed arrays of The Precambrian Earth: "Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Muellcr and O. Catuncanu

4.3. Komatiites

299

Fig. 4.3.3-1. (a) Light coloured spherical to amoeboid structures (Abitibi belt). Field of view 1 m. (b) Altered plagioclase spherulite from margin of pillowed basalt. Spherulite is circular in section and has a concentric structure (Abitibi belt). Field of view 2 cm under plane polarised light. (c) Altered plagioclase spherulites from a pillowed basalt (Abitibi belt). Spherulites emanating from a line classify as axiolites. Field of view 5 mm.

Chapter 4: Precambrian Volcanism

300

(d) Rate of crystal nucleation & growth vs.undercooling

"I"1

Undercooling

Tg

Fig. 4.3.3-1 (continued). (d) Nucleation and growth rate as a function of temperature. T 1 is the liquidus temperature and Tg is the glass transition temperature. Finite cooling must be achieved before nucleation occurs and nucleation is absent below Tg. fibrous crystals that emanate from a line or point. Each fibre has a crystallographic orientation slightly different from its neighbour and hence a "Maltese cross" extinction pattern under crossed-nicols. The term spherulite is misleading because they can be organised into linear forms that resemble sheafs or combs (Fig. 4.3.3-1c). Although feldspar spherulites are the most common, the habit has also been observed in quartz, pyroxene and high-polymers (e.g., nylon). Spherulites commonly range in scale from 1-10 mm, but rare metre-scale spherulites have been reported (Smith et al., 2001).

Nucleation and crystal growth A brief review of unusual crystal morphology in komatiites and variolitic rocks is warranted in order to better understand the kinematics of such processes. Lasaga (1998) has provided a detailed treatment. Mineral growth in lava is driven by a number of factors, including magma mixing and loss of volatiles, but here focus is placed on heat loss, the major factor in the rocks under consideration. At equilibrium, a crystal within a silicate melt neither grows nor dissolves. The entrenched term "equilibrium growth" is incorrect, as all growth proceeds away from equilibrium. Growth is an attempt to achieve equilibrium in accordance with conditions imposed upon the system. For thermally driven crystallisation from a melt, undercooling, i.e., cooling the system below the equilibrium or liquidus temperature for a particular phase is required. Under far-from-equilibrium conditions (sudden cooling), there is interplay between crystal growth and nucleation rates, influenced by the diffusion of growth constituents and heat, and by crystal growth anisotropies. This can lead to feedback or nonlinearities such that distinctive growth patterns spontaneously emerge. These patterns are termed self-organised (e.g., Ortoleva, 1994) and result from the growth kinetics, not growth on a pre-existing template. Examples of self-organised pattern formation in mineral growth include snowflakes, spherulites, spinifex, oscillatory-zoned crystals, and fractal olivine (Fowler et al., 1989). Crystal growth is initiated from a nucleus, a stable ensemble having a critical radius. These are either assembled in the melt by homogeneous nucleation, or are present as pre-

4.3. Komatiites

301

existing solid "impurities", so-called heterogeneous nucleation. For homogeneous nucleation, clusters with smaller than critical radii are unstable and will not grow. Nucleation does not start directly below the liquidus temperature due to a nucleation barrier below the glass transition temperature. The curve maxima indicate that as the temperature drops the silicate melt becomes more viscous and diffusion of growth species is inhibited. The effect is demonstrated at sill and dyke margins where chilled margins preserve numerous small crystals. Crystal growth rates near to, and far from, the liquidus temperature exceed nucleation rates (Fig. 4.3.3-1 d) because it is more difficult to assemble a stable nucleus of critical size from a silicate melt than to add growth species to a crystal. Once a stable nucleus forms, crystal growth may proceed. This may include the following processes: (1) diffusion of growth species to the surface, (2) attachment to the surface, (3) diffusion to specific growth sites, and (4) diffusion of heat. Under near-to-equilibrium conditions, for example very small undercoolings characteristic of large plutons, crystal growth proceeds by the orderly infilling of growth constituents on kinks, steps, or other non-planar crystal-surface irregularities. Filling of these sites is preferred as more free energy is expended, because here there is a greater energy loss due to bonds being formed at the ledge- and plane-face, rather than the face. At small undercoolings, crystal faces are atomically smooth but at increased undercoolings a roughening transition occurs so that the crystal face is no longer smooth at the atomic scale. Thus, under near-to-equilibrium growth conditions crystal surface diffusion and attachment kinetics favour compact, euhedral, compositionally homogeneous crystals. Continued undercooling facilitates the formation of compositionally zoned crystals. In general, crystal zoning only occurs for minerals that are members of solid-solution series, with plagioclase being the archetype (e.g., Shore and Fowler, 1996). At larger undercoolings, anisotropic, skeletal and dendritic habits develop. Strongly anisotropic textures such as plate spinifex, comb layering and cres-cumulate textures are characterised by elongated crystals. Often the long axes of the crystals are oriented normal to former cooling contacts. Nuclei that are oriented with their fastest growing faces normal to the cooling contacts grow preferentially, and starve less optimally oriented crystals from growth. Skeletal olivine crystals (Fig. 4.3.3-2a) form when corners and edges of crystals grow more than planar faces. Corners and edges subtend more solid-angle in the melt than plane-faces and when growth is rapid and diffusion in the melt sluggish, they can grow faster than the latter. Sections through skeletal crystals give the illusion that they are composed of disconnected though crystallographically oriented parts. Dendrites (Fig. 4.3.3-2b) have parabolic shaped crystal tips and an ordered morphology characterised by a regular arrangement of sidebranches along specific crystallographic axes. Sections through dendritic crystals are also skeletal, and the two habits are not easily discriminated. Spherulitic morphology occurs at still larger undercoolings below the point of roughening transition. Growth rate is rapid, diffusion in the liquid is slow and the crystallites become microscopically rough. Rapid growth promotes the accumulation of low melting temperature constituents adjacent to the crystallite-melt boundary, locally causing an effective undercooling. Protrusions on the rough crystallite project through the accumulation into the "undercooled" zone and grow to produce an organised array of fibrous crystals.

302

Chapter 4: Precambrian Volcanism

Fig. 4.3.3-2. (a) Spinifex texture in thin-section (Barberton belt), one of the freshest known Archaean komatiites (Nisbet et al., 1987). Note skeletal olivine crystals randomly oriented in a matrix of acicular augite crystals and altered glass. (b) Cr-spinel dendrite from a komatiite flow (Pyke Hill, Abitibi belt). Field of view 150 ~m in plane polarised light. (c) Detail of variole resulting from magma mingling (Abitibi belt). Field of view 2 mm. Altered phenocrysts of quartz and alkali feldspar with spherulitic overgrowths are rhyolitic in composition and are found as globules in mafic rocks. (d) Spinifex texture from type section of komatiites (Barberton belt, South Africa; Viljoen and Viljoen, 1969b). Textures formed below A1 flow top are a zone of A2 randomly oriented olivine blades and A 3 books of parallel elongate blades of olivine oriented at a high angle towards the flow top.

4.3. Komatiites

303

Spherulite growth occurs directly from the melt at high undercoolings. Similar forms also result from devitrification of glass, though probably only through the intervention of fluids because growth is strongly impeded at surface temperature (Manley, 1992). Spherulites are common in dacites and rhyolites because silica-rich compositions cause the melt to be highly polymerised. Under conditions of very high undercooling, crystal growth at the margins of a rapidly cooled aphyric lava flow produces non-compact branching crystals characterised by several orders of non-crystallographic branching. The crystals are fractal objects (Fowler et al., 1989) that can be modelled using the DLA algorithm (Fowler et al., 1989; Fowler and Roach, 1996). Crystal growth occurs in a steady-state field (e.g., invariant temperature gradient) and is dominantly controlled by the random diffusion of growth constituents in the melt. The constituents freeze the instant they collide with the growing crystal. Branching growth is favoured because random-walking growth constituents are more likely to collide with branch tips than to penetrate deep between the branches, and thus are self-propagating. Rapid cooling is not the only mechanism capable of producing far-from-equilibrium crystal morphologies in igneous rocks. A sudden loss of volatiles from magma abruptly decreases PH20 and increases the liquidus temperatures of its silicate minerals, producing an effective undercooling. This process was responsible for the formation of branching, skeletal olivine crystals in the Rum intrusion of Scotland (Donaldson, 1974). As shown by Lofgren and Russel (1986), the melt history may also play an important role in the development of rock texture. Superheating, which raises the system temperature above the liquidus, will destroy pre-existing nuclei, embryonic nuclei, and crystals. Cooling of superheated experimental charges produced non-equilibrium habits at lower undercoolings than charges not superheated due to the lack of nuclei. Varioles in volcanic rocks of the Abitibi greenstone belt

Early work on variolitic rocks from the Abitibi greenstone belt focused on pillowed, massive and flow-banded melanocratic volcanic rocks. The varioles, ranging from mm- to cm-scale in diameter, are generally leucocratic and weather recessively. Internal structures are inconspicuous at the macroscopic scale. Several types of phenomenon may give rise to varioles in these rocks. G61inas and Brooks (1974) concluded, using major element composition (roughly a low-K rhyolite) and shape, that these cm-scale varioles were produced by liquid silicate immiscibility, whereas mm-scale features were plagioclase spherulites. In contrast, Philpotts (1977) and Hughes (1977) favoured a spherulite interpretation for both. Fowler et al. (1986) argued against an immiscibility model using trace-element partitioning and detailed textural observations. They argued that the structures were plagioclase spherulites that grew directly from the melt, and present-day albite mineralogy fortuitously yields a chemical composition similar to "low-K rhyolite". Petrographic investigation of the texture located close to the pillow margin revealed the following transition: (1) altered glass and in situ breccia containing no crystals, (2) altered glass containing sparse mm-scale plagioclase spherical spherulites, (3)cm-scale more abundant and coarser spherulites, (4) arrays of mutually interfering axiolitic plagioclase and clinopyroxene spherulites, in which the spherulites are coarser than those near the cooling margins,

304

Chapter 4: Precambrian Volcanism

and (5) locally isolated skeletal crystals. Spherulites may coalesce, but planar boundaries between individual spherulites suggest growth from individual nucleation points. Larger pillows have spherulite-rich interiors due to flow differentiation of spherulites within lava tubes. Plagioclase spherulites in Abitibi basalts are restricted to aphyric tholeiitic lavas, consistent with superheating and an absence of nuclei. Basalt extrusion on the ocean floor caused rapid cooling, in which the few nuclei formed were rough at the atomic scale and grew rapidly to form spherulites. These quench spherulites should not be confused with devritification spherulites. Other varioles are observed as leucocratic mm- to cm-scale globules that weather in positive relief relative to the mafic hosts. These are found within tholeiitic volcanic rocks but are associated with metre-scale rhyolite lobes. The cores of these varioles contain small euhedral crystals of quartz and alkali feldspar that served as nuclei for branching crystals of these minerals (Fig. 4.3.3-2c), but a mingling of basalt and rhyolite is inferred (Fowler et al., 2002). The rhyolite was mechanically disrupted during eruption and entrained within the basalt as variably sized entities. Ropchan et al. (2002) described variolitic rocks of this type within the Holloway Au-Mine (Abitibi greenstone belt).

Spinifex textured komatiites It is easy to say roughly what a komatiite is, but very difficult to come up with a rigorous definition. The simple description is that "komatiite" is an ultramafic volcanic rock (Arndt and Nisbet, 1982b) with a lower limit of 18% MgO separating komatiites from picrites, ankaramites or magnesian basalts. Implicit in the definition of komatiite is the notion, difficult to prove, that komatiites crystallise from liquids that contained > 18% MgO. Complications arise from the existence of other volcanic rocks with more than 18% that either formed through the accumulation of olivine from less magnesian liquids, or crystallised from magmas with chemical characteristics quite unlike those of most komatiites. An example of the first type is a phenocryst-charged basaltic liquid (a picrite according to some definitions); an example of the second is meimechite (Arndt et al., 1995), a rare alkaline lava with unusual major and trace element composition. To distinguish komatiite from other highly magnesian volcanic rocks, it is useful to include spinifex texture in the definition, yet not all komatiite flows have spinifex (Nesbitt, 1971). A workable definition includes the phrase "komatiite is an ultramafic volcanic rock containing spinifex or related to lavas containing this texture". With the last part of this definition allowance is made for the manner in which texture varies within komatiitic units. For example, many komatiite flows have an upper spinifex-textured layer and a lower olivinecumulate layer. Other flows grade along strike from layered spinifex-textured portions to massive olivine-phyric units. With the inclusion of the phrase about spinifex, the lower olivine-cumulate portions of layered flows or the olivine-phyric units can also be described as komatiite. On the other hand, meimechites, picrites and other rock types that contain no spinifex are excluded. For further discussion, consult Le Bas (2000, 2001) and Kerr and Arndt (2001). Within the upper parts of komatiite flows (Fig. 4.3.1-1), the type of spinifex texture varies systematically (Fig. 4.3.3-2d). Beneath a thin (1-5 cm) glassy, commonly por-

4.3. Komatiites

305

phyritic chill zone (A1), a layer of "random" spinifex texture contains isolated randomly oriented crystals or cm-scale "booklets" of parallel plates of olivine, in a matrix of finegrained clinopyroxene and devitrified or altered glass (A2). The underlying layer of platy olivine spinifex has an organised structure wherein arrays of large bladed olivine crystals, possibly metres long, are oriented roughly perpendicular to the flow top (A3), forming sheaf-like structures that fan out from flow tops and serve as top indicators. The lower parts of spinifex komatiite flows are cumulates (B zone) containing settled solid polyhedral olivine crystals (see Pyke et al., 1973; Donaldson, 1982; Shore, 1996). Pyroxene spinifex forms in the upper parts of komatiitic basalt flows. In this texture, needle-like crystals, commonly with pigeonite cores and augite margins, lie in a matrix of augite, altered glass and/or plagioclase and oxides. The pyroxene needles range in length from a few mm to several cm and their orientation is either random or perpendicular to the flow top (see Fleet and MacRae, 1975; Arndt and Fleet, 1979).

Origin of spinifex Viljoen and Viljoen (1969a, b), who first recognised komatiite as a separate rock type, used the term "crystalline quench texture" for spinifex. The formal introduction of the term "spinifex" by Nesbitt (1971) was based on the description and classification of different types of skeletal crystals in komatiites from Australia and Canada. He compared skeletal or dendritic morphologies of olivine and pyroxene crystals in natural spinifex textures with experimental charges and silicate slags, but also outlined what has come to be known as the "spinifex paradox". Spinifex texture is found in komatiite flows, well below the upper chilled crust. In the thickest units, large dendritic crystals may have crystallised at depths ten or more metres below the surface of the flow. Under such circumstances, the loss of heat from the interior of the flow is controlled by conduction through the upper solidified crust. In a 2 m-thick komatiite flow, the cooling rate during crystallisation of the lower part of the spinifex layer is only 1-3~ per hour. In thicker flows the rate is far lower. In contrast, the morphology of olivine or pyroxene crystals in spinifex-textured lavas resembles those produced experimentally at cooling rates never less than about 30~ hr -l (Donaldson, 1976, 1982; Fig. 4.3.3-3a). Simply stated, the spinifex paradox refers to the presence, at depth within a komatiite flow where cooling rates must have been low, of elongate skeletal crystals whose crystal morphologies resemble those formed in experiments at much higher cooling rates. Donaldson (1976) was the first to study experimentally the formation of spinifex textures. By extending an approach used by Lofgren et al. (1974), Donaldson (1976) developed a scheme whereby the morphology of olivine crystals could be related to the experimental conditions, particularly to the rate of cooling and/or the degree of undercooling. The morphology of olivine crystals in spinifex-textured komatiites (Figs. 4.3.3-2a and 4.3.3-3a) is similar to those of olivine crystallised at cooling rates around 40~ hr-l in basaltic melts. The spinifex paradox could be explained by two possible avenues. First, the discrepancy could be due to the high MgO content in komatiites, which led to the development of highly skeletal morphologies at lower cooling rates. Second, the skeletal or dendritic morphology resulted from rapid crystal growth, but not necessarily from rapid

306

Chapter 4." Precambrian Volcanism

g~ r

,4 ~ ,,,~

4.3. Komatiites

307

cooling. Donaldson (1976) did not explain how the growth rate could be decoupled from the cooling rate within the lava flow. Once a thick upper crust develops, the rate of crystal growth at the solid-melt interface of the crust will be controlled by the rate at which heat is lost from the lava flow. The cooling rate and the rate of crystal growth should be controlled by the efficiency with which heat is transmitted through the crust. Turner et al. (1986), and Shore and Fowler (1999) attempted to explain the spinifex paradox by suggesting komatiite cools far more rapidly than predicted by simple conductive cooling models. Turner et al. (1986) suggested that vigorous internal convection in ponded komatiite would greatly enhance the rate of heat loss and proposed cooling rates of 1-100~ hr-1 soon after emplacement. Such rapid cooling would cause the interior lava to become highly supersaturated, leading to the formation of skeletal olivine, but this is only valid if the crust of the flow is very thin. As the crust thickens, interior convection stops and heat loss is controlled by conduction through the crust. In more recent models (e.g., Renner et al., 1994) interior convection is limited to the initial stages of cooling of only the most magnesian (MgO > 28%) komatiites. Shore and Fowler (1999) conducted a detailed petrological and textural study of the classic Pyke Hill outcrop in Munro Township, Canada. They proposed two mechanisms that might cause a flow to cool more rapidly than predicted by conductive cooling models. The first is hydrothermal cooling, whereby as komatiite cools, it contracts, leading to the formation of a network of fractures in the upper part of the crust. Circulation of sea water through these fractures cools the solidified upper portion of the crust. The efficiency of this process is difficult to judgemalthough fractures are present in the upper parts of komatiite flows, they are neither abundantly distributed nor continuous in two dimensions. A second mechanism involves heat transfer by radiative and lattice thermal conductivity through the aligned olivine crystals of spinifex textures. Shore and Fowler (1999) determined that the crystallographic a axis was consistently oriented near perpendicular to the flow top. Experimental work showed that in high magnesian-komatiite, the rate of heat transfer along this axis would be 3-5 times greater than that of conduction. Olivine crystals favourably aligned to the cooling front create a steep thermal gradient in the liquid ahead of their tips, thus supporting self-propagating growth. However, the proposed heat-transfer mechanism is inefficient in less magnesian liquids, and the mechanism does not provide an all encompassing explanation for spinifex. Grove et al. (1994, 1996, 1999), based on field observations and experimental work, proposed a provocative alternative, whereby the spinifex texture in Barberton komatiites could not be explained by normal crystallisation of anhydrous magma. The spinifex para-

Fig. 4.3.3-3. (a) Relationship between olivine morphology and cooling rate as inferred by Donaldson (1974, 1976). (b) Diagram indicating the "spinifex paradox". (c) Stages 1-3 during solidification of a spinifex textured komatiite flow. (d) Synthetic spinifex texture in a fayalite slag illustrating constrained growth of olivine and the mechanism that leads to preferred orientation perpendicular to the cooling surface.

308

Chapter 4: Precambrian Volcanism

dox was used to explain the presence of large water contents. According to these authors, the role of water is two-fold. First, the presence of water in a silicate melt impedes crystal nucleation and increases the diffusion rate in the silicate liquid, leading to rapid growth of large crystals. Second, degassing of hydrous komatiite as it approaches the surface dramatically increases the liquidus temperature, producing a strongly supercooled liquid. The spinifex texture then results from rapid crystal growth in the interior of the supercooled komatiite. However, this mechanism depends on two premises: (1) komatiites are hydrous, and (2) Barberton komatiites crystallised as sills. Both premises have been contested (Arndt et al., 1998; Dann, 2000, 2001; see also sections 4.3.1 and 3.6).

A possible solution to the spinifex paradox A factor mentioned in many papers on spinifex texture that has not received sufficient attention, is the role of constrained crystal growth during solidification of the crust of a komatiite flow. Constrained growth refers to the crystallisation of parallel grains of olivine or pyroxene in the downwards-growing crust of a lava flow. The crystals compete with one another for "nutrients", the atoms of Mg, Fe and Si that are essential crystal framework components. This competition leads to the preferred, near-perpendicular orientation of the olivine crystals in spinifex textures (Figs. 4.3.3-2d and 4.3.3-3b-d). Erupted komatiite contains a small proportion of olivine phenocrysts that grew either during magma ascent or during lava flow. During flow cooling, phenocrysts are trapped in the crust at the top of the flow, and others settle to become part of the cumulate layer. Olivine nucleates in the crystal-free liquid beneath the crust. The crystals growing from these nucleation sites are randomly-oriented and highly skeletal morphologies attributable to the high cooling rate form in the crust. As the cooling proceeds, the olivine grains with near-vertical orientations grow preferentially because their tips extend downwards into unfractionated nutrient-rich liquid, whereby orientations closer to horizontal find nutrient-poor liquid or collide with other crystals, and cease to grow. The crystallisation of olivine produces a residual liquid depleted in Mg and enriched in Si, A1, Ca and Na. It is less dense than the parental liquid. As downwards growth proceeds, this liquid is expelled and accumulates as a layer of low density at the base of the crystal front (Turner et al., 1986). The growing tips of the spinifex crystals are bathed in a liquid depleted in the components they require to grow. Faure (2001) suggested that this situation provides an explanation for the unusual habit of spinifex crystals: a solution to the spinifex paradox? This situation has certain parallels with the accumulation of nutrient-poor zones surrounding rapidly growing crystals in a quenched liquid. Here, the rate of crystal growth exceeds the diffusion rate of the major elements within the silicate liquid, and the elements surround the crystal. The skeletal or dendritic morphology is caused by propagation of fine needles or plates emanating from the crystal and penetrating the nutrient-poor layer (Fig. 4.3.3-3c). The dendritic habit may be a consequence of their growth into the accumulated layer of nutrient-poor liquid. The solution to the spinifex paradox rests in developing a better understanding of constrained crystallisation in high Mg-liquids. The temperature difference between liquidus and solidus contributes to the growth of spinifex, and in komatiites, this difference is

4.3. Komatiites

309

c. 400~ (Fig. 4.3.3-4a), whereas it is < 100~ in basaltic magmas. The crust of a komatiite flow contains a very thick crystal mush zone, in which olivine crystals are enveloped by a silicate liquid. This situation facilitates the expulsion of olivine-depleted liquid and favours the accumulation of olivine. As pointed out by Barnes et al. (1983) and Arndt (1986), spinifex lavas are "coagulation cumulates" containing a higher proportion of liquidus mineral(s) than the liquid from which they crystallised. Growth of spinifex crystals downwards from the crust into nutrient-poor liquid may provide an explanation for several hitherto puzzling aspects of komatiite flows, such as (1) the contrasting mineralogy in the upper and lower parts of komatiitic basalt flows and (2) the precocious pyroxene problem. (1) The mineralogy of the spinifex-textured upper zone differs from that of the lower cumulate zone. In Fred's Flow, a thick layered komatiitic basalt in Munro Township (Arndt, 1977b), the succession of liquidus minerals in the spinifex zone is olivine (+ chromite) --+ pigeonite ~ augite ~ plagioclase. In the lower part of the flow, olivine-chromite cumulates are overlain in turn by orthopyroxene-augite cumulates and orthopyroxeneplagioclase cumulates. The contrasting behaviour might be explained if rapid crystallisation leads to the build-up of olivine-poor, pyroxene-saturated liquid at the base of the growing spinifex layer. Pyroxene spinifex results from crystallisation within this layer, whose composition differs from that of liquid lower in the flow from which the cumulus phases crystallise. (2) In komatiitic basalts, pyroxene crystallises sooner than expected according to equilibrium phase relations. Experimentally, olivine-free spinifex-textured lava with pigeonite as the liquidus phase crystallises olivine first, followed at lower temperatures by augite and plagioclase (Arndt, 1977c; Arndt and Fleet, 1979; Fig. 4.3.3-4b). In the trend of compositions of natural lavas the kink at about 15% MgO indicates the onset of pigeonite crystallisation, whereas in equilibrium melting experiments, the kink at 12% MgO corresponds to the crystallisation of augite immediately followed by plagioclase. This contrast in behaviour is readily explained if the spinifex texture crystallised from a layer of olivine-depleted liquid whose composition was far from that of the equilibrium liquid.

Conclusions

Variolites are rocks containing centimetre-scale leucocratic globular structures in a finegrained mafic rock. The term variole is best used as a descriptive term in the field. Because Archaean rocks are altered and metamorphosed to some degree further work is often required in order to discern their original nature. Upon detailed examination, varioles are spherulites, amygdules, blotchy alteration fronts, magma-mingling textures, and altered phenocrysts. Varioles proven to be the result of liquid silicate immiscibility have yet to be documented in the Abitibi greenstone belt. Many tholeiitic basalts are characterised by large and abundant altered plagioclase spherulites that grew directly from the melt. These basalts are always aphyric demonstrating that they were superheated. Their abundance in numerous Archaean sequences supports the concept that Archaean thermal regimes were different to those of later Eons (see also section 3.6).

Chapter 4: Precambrian Volcanism

310

(a) 1600

T komatiite .., 350~o

~, komatiite liquidus

1500 o

1400

L_

~ 1300

~,~

T basalt

L_

~ , ~ ~

~ 5 0 ~, E

~. 12oo

cpx + plag +ol

1100

~,

basalt liquidus -komatiite 9 and basalt solidus

t I

100

% liquid

(b)

MgO

,70 6O

..... /

/

equilibdum crystalhsation trend

9ee

\

9e~ 9

CaO

25

35

45

55

65%

AI203

Fig. 4.3.3-4. (a) Diagram of percentage liquid versus temperature, illustrating the large difference in the liquidus-solidus gap between komatiite and basalt. Data from Arndt (1976). (b) MgO-CaO-A1203 diagram, with data from Arndt (1977a), showing different trends between the compositions of natural komatiitic basalts and the equilibrium crystallisation trend. The kink away from CaO marks the onset of clinopyroxene crystallisation. The origin of spinifex, a texture restricted to komatiite, remains problematic. It is difficult to explain how a texture that appears to require rapid cooling can form deep below the crust of a komatiite flow, where the cooling rate must have been low. A possible solution is that many of the characteristics of spinifex, particularly the size, preferred orientation, and habit of the crystals are explained by cooling in a thermal gradient, such as exists in

4.4. Archaean a n d Proterozoic Greenstone Belts

311

the upper part of every lava flow. The formation of spinifex is linked to the temperature difference between liquidus and solidus, which is very large in komatiites. This produces a thick partially molten zone in the upper part of the flow where conditions required for spinifex formation are met. The texture highlights the peculiar nature of komatiites, a truly Archaean magma type.

4.4.

ARCHAEAN AND PROTEROZOIC GREENSTONE BELTS: SETTING AND EVOLUTION

EC. THURSTON AND L.D. AYRES Archaean and Proterozoic greenstone belts are typically linear to anastomosing to cusplike, steeply-dipping areas of deformed volcanic, volcaniclastic and sedimentary rocks that are bordered and intruded by voluminous granitoid suites (Fig. 4.4-1; Condie, 1981) (see chapter 2 for extensive discussion of granite emplacement and the evolution of granitegreenstone terranes). The metamorphic grade is generally greenschist with lesser high grade areas (Wilkins, 1997), but in some American and Scandinavian Proterozoic belts, amphibolite grade metamorphism is prevalent (Park, 1991; Condie, 1992b). Granitoid rocks may form the basement to greenstone belts as in the Slave Province (Bleeker et al., 1999; Corcoran and Dostal, 2001; Mueller and Corcoran, 2001) but more commonly are coeval with, or slightly to considerably younger than volcanism (Hirdes et al., 1992; Martin et al., 1993; Swager, 1997; Thurston, 2002). In Archaean and some Proterozoic greenstone belts, early granitoid plutons are the tonalite-trondhjemite-granodiorite suite (TTG) and later plutons are more potassic (Arkani-Hamed and Jolly, 1989; Harris et al., 1993; Wolde et al., 1996b; Carlson, 1997; Doumbia et al., 1998; Whalen et al., 1999), but in other Proterozoic belts early plutons are calc-alkalic granodiorite (Ga~il and Gorbatschev, 1987). Archaean greenstone belts are best studied in the Pilbara and Yilgarn of Australia, the Dharwar of India, the Kaapvaal and Zimbabwe of Africa, and the Superior and Slave cratons of North America as well as the Baltic shield of Europe. Greenstone belts form 5-60% and average 10% of the upper continental crust in these cratons (Condie, 1981; Goodwin, 1996; Barley, 1997; Blenkinsop et al., 1997; Brandl and de Wit, 1997; King and Helmstaedt, 1997; Myers and Swager, 1997; Stott, 1997). In southern Africa and Australia, Neoarchaean cover sequences unconformably overlie older greenstone-granitoid terranes; some are flood basalts (e.g., Clendenin et al., 1988; Marsh et al., 1992; Blake, 2001). Archaean greenstone belts, TTG terranes and high grade terranes typically form craton-wide associations (Brandl and de Wit, 1997; King and Helmstaedt, 1997; Myers and Swager, 1997; Stott, 1997) considered to be mobile belts (Card, 1990). Many volcanic rocks in Archaean greenstone belts are subduction-generated juvenile island arcs and mantle-derived, oceanic crust, including plume-related oceanic plateaus (Goodwin, 1996) (see also section 3.6). Many Archaean cratons, such as the Superior, Yilgarn, and Zimbabwe, have an older sialic basement beneath some greenstone belts (e.g., Jolly and Hallberg, 1990; Martin et al., 1993; Swager, 1997; Thurston, 2002). The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

312

Chapter 4: Precambrian Volcanism

Fig. 4.4-1. Three Archaean greenstone belts at the same scale to show variations in shapes and relationships to other rock units. (a) Cuspate Barberton greenstone belt, Kaapvaal craton, South Africa (simplified from Brandl and de Wit, 1997), (b) curvi-linear North Caribou greenstone belt, Superior craton, Canada (simplified from Stott, 1997), and (c) anastomosing, cuspate greenstone belts of the Pilbara craton, Australia (simplified from Barley, 1997).

Proterozoic greenstone belts generally occur in mobile belts adjacent to older cratonic blocks. These greenstone-granitoid terranes are bordered by sedimentary craton-cover sequences and, in places, also by ophiolites (Lewry and Collerson, 1990; Stern, 1994; Goodwin, 1996; Nironen, 1997) (section 3.7 reviews Precambrian ophiolites). Volcanic rocks in Proterozoic greenstone belts are considered to be juvenile island arc and oceanfloor sequences with minor contributions from older crust (Stern, 1994; Goodwin, 1996; Lucas et al., 1996; Mellqvist et al., 1999). In the Proterozoic, there are many deformed volcanic sequences that have similar lithology, shape, or tectonic setting as those in greenstone belts, but have a higher metamorphic grade (Van Schmus et al., 1993)(see also section 3.9) or a low abundance of granitoid rocks (Sims et al., 1989); these are not considered in this review. Therefore identifying Proterozoic greenstone belts may be problematic. Volcanic cyclicity involves the repetition of rock types or stratigraphic units on large to small scales (Anhaeusser, 1971). Large-scale cyclicity can develop from (1) competition between upwards growth of volcanoes by intermittent or continuous eruptions, and downwards movement resulting from erosional degradation and isostatic loading; (2) interaction between intermittent volcanic processes and sedimentation (Anhaeusser, 1971); and (3) temporal changes in source regions (e.g., Bailes and Galley, 1999). Such cyclicity is common in greenstone belts, but it is not always termed cyclicity because some cy-

4.4. Archaean and Proterozoic Greenstone Belts

313

cles involve repetition of sharply bounded lithofacies rather than progressive changes (e.g., Bailes and Syme, 1989). Furthermore, some large-scale compositional variations termed cyclicity (e.g., Ayres, 1977a; Thurston, 1986) were the result of thrust stacking (e.g., Ayres and Corfu, 1991) (section 3.6). Because of eruption from magma chambers zoned with respect to composition, volatile content, and gas bubble content, many volcanic sequences also have small-scale cyclicity, (Thurston, 1986; Baldwin, 1988; Bailes and Syme, 1989; Allen et al., 1996a). The notion of volcanic cyclicity based on different magma-forming processes, such as plumes (sections 3.2 and 3.3), has added a new wrinkle. The constant interplay between plume- and subduction-generated magma has significant impact on the interpretation of volcanic cyclicity in greenstone belts. Focus is placed on evolutionary trends and cyclicity in Archaean and Proterozoic greenstone belts with emphasis on the stratigraphy and physical volcanology of arc and back-arc as well as oceanic settings. Plume-related volcanism occurs at all localities. The Canadian (Fig. 4.4-2), Australian and Baltic shields are the primary sources of information with numerous insights drawn from other cratons. Archaean Greenstone Belts

Archaean greenstone belts developed as oceanic or continental arcs may be autochthonous as indicated by basal and inter-assemblage nonconformities, isotopic inheritance, contamination of magmas, and stratigraphic relationships (Chauvel et al., 1985; Bickle et al., 1994; Bleeker et al., 1999; Thurston, 2002). They also developed as oceanic terranes or oceanic plume-related sequences (Ohta, 1996; Corcoran, 2000; Tomlinson and Condie, 2001) with some oceanic volcanism preserved at accretionary margins (Devaney and Williams, 1989; Krapez and Eisenlohr, 1998). Plume-generated komatiites have generally been associated with oceanic plateaus (Storey et al., 1991; Kerr et al., 1997), ponded ocean floor lava fields (Squire et al., 1998), continental or oceanic assemblages (Hollings et al., 1999; Tomlinson et al., 1999; Tomlinson and Condie, 2001) but also as an integral component of arc sequences (Dostal and Mueller, 1997; Ayer et al., 2002) (also see discussions in section 3.6). The arc-plume concept, pivotal in understanding greenstone belt evolution, is especially evident in the Abitibi greenstone belt (see also section 2.4) where komatiites are intercalated with subduction-derived andesites (Houl6 et al., 2001) or arc volcanoes (Dostal and Mueller, 1997). The following settings are currently recognised in Archaean greenstone belts: (1) shallow water quartz- and carbonate-rich platforms (see also section 7.3) with minor volcanism unconformably overlying basement rocks or greenstone belts, (2) shallow to deep water komatiites and tholeiitic basalts overlying either platformal sequences or granitoid basement, (3) deep-water oceanic komatiite-tholeiite or tholeiite sequences including mafic plain sequences, (4) deep-water oceanic sequences consisting largely of tholeiitic basalt with minor felsic pyroclastic rocks, chert, and iron-formation (reviewed in section 5.4), (5) shallow water to emergent arc sequences with bimodal volcanic successions, and (6) subaerial pullapart basins with alkalic to calc-alkaline volcanism (Thurston and Chivers, 1990). On the scale of the Superior Province (Fig. 4.4-2) an orderly progression from (1) to (6) is inferred.

314

Chapter 4: Precambrian Volcanism

Fig. 4.4-2. Canadian Superior Province with major tectonic zones and terranes (after Stott, 1997). Greenstone belts in dark grey and indicated as island arc supracrustal belts.

The geodynamic setting of many Archaean greenstone belts has modern counterparts, with a scarcity of oceanic assemblages (Williams et al., 1992) and ophiolites (Sylvester et al., 1997).

4.4. Archaean and Proterozoic Greenstone Belts

315

The bimodal nature of Archaean arcs and lack of orogenic andesites may be explained by high heat flow and spreading rates (Abbott and Hoffman, 1984). Cooling of the hotter Archaean Earth produced a thicker oceanic crust with increased spreading rates and ridge lengths (see, however, discussion in section 3.6). Shallow subduction (section 3.5) is the consequence of the younger, more buoyant oceanic crust. Experiments indicate that the melting of hydrated basalt at shallow depths with higher geothermal gradients will yield lower temperature siliceous melts compared to steeper dipping subduction (Helz, 1976). The shallow subduction of plates favours bimodal volcanism, but also explains the lack of modern style orogenic andesites, which require steeper subduction gradients. Adakites and ubiquitous presence of TTG support the notion of shallow dipping plates (Boily and Dion, 2002; Wyman et al., 2002b). Similarly, boninites are high temperature, shallow level melts generated either by plume impingement on the subduction zone (Kerrich et al., 1998) or by subduction processes (Boily and Dion, 2002). Arc and back-arc volcanism Basins adjacent to unroofed Archaean arcs locally have shoshonitic lava flows (Picard and Piboule, 1986b; Dostal and Mueller, 1992), similar to successor basins of dissected arcs with calc-alkaline to alkaline volcanism (Cooke and Moorhouse, 1969), as well as high Mg-basalts (Gaal and Gorbatshev, 1987; Krapez and Eisenlohr, 1998). Largely Andean style arcs, back-arcs and island arcs constitute the Archaean Superior Province (Thurston et al., 1991; Rogers et al., 2000; Ayer et al., 2002). Inferred arcs are prominent in the Lower and Upper Bulawayan sequences in the Zimbabwe craton (Blenkinsop et al., 1997), the Yilgarn craton (Myers and Swager, 1997), and the Murchison greenstone belt of the Kaapvaal craton (Vearncombe, 1991). In contrast, ensimatic back-arcs appear rare (Krapez, 1993). The volcanology of Archaean arcs varies with the rate of magma effusion and the degree of magma ponding en route to the surface. Greeley (1982) divided basaltic volcanism into high volume flood volcanism, intermediate volume lava plain volcanism, and low-volume Hawaiian volcanism, all of which are integral parts of arcs (Table 4.4-1) and the oceanic floor of back-arcs. In stratigraphic terms, Archaean arc and back-arc sequences consist of a broad basal lava plain succeeded upwards by mafic shield volcanoes (Ayres, 1982), upon which central (bimodal) volcanic complexes developed. The advent of rhyolitic volcanism is commonly associated with mature arc construction, arc rifting and back-arc development (Chown et al., 1992; Corcoran and Dostal, 2001). Volcanology of lava plains and shield volcanoes Lava plains and shield volcanoes (Fig. 4.4-3), generated either by plumes or subduction processes, are dominated by mafic volcanism. The 5-7 km thick subaqueous plains, 100-150 km long, are composed of overlapping shield volcanoes that are > 25 km in diameter (Dimroth and Rocheleau, 1979; Thurston and Chivers, 1990). Lava flows up to 150 m thick and traceable for tens of kilometres, grade from thick, proximal massive gabbroic textured basaltic flows to master tubes with branching megapillows that change at the distal ends to normal-sized pillows with a cross-sectional area averaging 2600 cm 2 (Sanshagrin, 1982). The extent of a representative single glomeroporphyritic flow in the

316

Chapter 4: Precambrian Volcanism

Table 4.4-1. Characteristics of mafic volcanism (after Thurston and Chivers, 1990). References: (1) Dimroth and Rocheleau (1979); (2) Greeley (1982); (3) Hooper (1997) Type

Flow thickness, m Flow area, km2 Flow facies

Slope

References

Lava plain volcanism

1-5

1, 2

101-102

Massive, tube and pillowed flows; local breccia

Low shield with 1/2 ~ slope

Flood basalt 101-102 volcanism

102-103

Massive thick subaerial and subaqueous flows with hyaloclastite

Fissure with 3 2-10 ~ slope

Hawaiian volcanism

101-102

Subaqueous setting: Low shield massive to pillowed volcano with to hyaloclastite 1-3 ~ slope

<5

lava plain is about 156 k m 2 (Dimroth et al., 1985). Such glomeroporphyritic basalts with extremely calcic plagioclase megacrysts serve as marker horizons and generally occur high in the stratigraphy (Phinney et al., 1988; Blackburn et al., 1991). Flow top breccia commonly caps massive flows, occurring less commonly above pillowed flows. Typically, lava plain basalt and/or komatiite flows have minor inter-flow sedimentary units. A relatively deep water environment is suggested by a paucity of oxide facies iron-formation, vesicular lavas, and hyaloclastic units (Dimroth et al., 1985). Subaqueous shield volcanoes with steeper slopes follow up-section (Table 4.4-1), and display higher proportions of pillowed flows and hyaloclastite, as well as a general increase of basalt vesicularity that may indicate shallowing of the sequence. The shoaling upwards aspect of mafic shield volcanoes is indicated by flow-foot breccias that form prograding deltas diagnostic of littoral environments (Dimroth et al., 1985). Sheets of massive lava forming topsets and foresets grading upwards and downwards into pillow lava and pillow breccia are suggestive of a littoral setting. These combined features are common to subaqueous segments of modern oceanic islands (e.g., Staudigel and Schmincke, 1984). Abitibi shield volcanoes are up to 7 km thick and > 30 km in diameter. Similar facies associations of arc-related basalts are seen in the Yilgarn (Brown et al., 2002) and Pilbara cratons of Australia (Kiyokawa and Taira, 1998; Krapez and Eisenlohr, 1998; Pike et al., 2002). These arc-related sequences vary from immature oceanic arcs (Kiyokawa and Taira, 1998) to shield volcanoes on older substrate (Krapez, 1993). Volcanic construction

In Archaean arc sequences the transition from lava plain and shield volcano to central volcanic construction is well preserved in the rock record, and is documented both by the subtle change in geochemistry or the abrupt change in physical volcanology from mafic to felsic volcanism. The upwards transition into the proto-arc stage is marked by the appearance of low Ti-basalts (Wyman et al., 1999b) intercalated with komatiites, and has been inferred at Kidd Creek, which hosts a giant massive sulphide deposit. From this discrete

4.4. Archaean and Proterozoic Greenstone Belts

317

Fig. 4.4-3. Cross-section of the Blake River assemblage in the Abitibi greenstone belt, Superior craton (after Dimroth and Rocheleau, 1979). The basal mafic plain sequence is overlain by two coalescing shield volcanoes, Montsabrais in the west and Reneault in the east. change, the sudden input of felsic volcanism ensues. The subaqueous, felsic-dominated, arc edifices are composite volcanoes with multiple vents (Lafrance et al., 2000), volcanic complexes (Legault et al., 2002), a series of small felsic centres formed along strike (Scott et al., 2002), or calderas (Gibson and Watkinson, 1990; see section 4.6). The complex arc edifices generally remained submerged, yet several centres breached the Archaean ocean, albeit temporarily. Stromatolites (section 6.5) fringing the Joutel volcanic complex (Hofmann and Masson, 1994; Legault et al., 2002) and the Back River volcanic complex (Slave craton; Lambert et al., 1990, 1992), as well as shallow water, wave-reworked volcaniclastic deposits in the Noranda caldera (Lichtblau, 1989) and islands of the Chibougamau arc (Mueller, 1991), attest to shoaling of edifices and local formation of atolls. The striking feature of almost all volcanic edifices is their subsequent flooding by subaqueous mafic to ultramafic lava flows. This mafic overlap is the base of the next volcanic cycle. A similar volcanological evolution of edifices has been recorded in the Yilgarn and Pilbara cratons of Australia (Hallberg, 1986; Brown et al., 2002; Pike et al., 2002). The principal components of subaqueous felsic volcanism are effusive flows and their autoclastic derivatives (e.g., de Rosen-Spence, 1976; de Rosen-Spence et al., 1980; see section 4.6). Inferred subaqueous ash-flow volcanism has been proposed for the Sturgeon Lake caldera (Morton et al., 1991; see section 4.6), the Selbaie caldera (Larson and Hutchinson, 1993), and for the subaerial welded ignimbrites in the western Superior Province (Thurston, 1980). The products of ash-flow volcanism range from welded to non-welded ash-flows (Thurston et al., 1985) to a variety of mass flow and air fall products (Hallberg, 1986; Barley, 1992; Krapez and Eisenlohr, 1998).

Arc unroofing volcanism Arc unroofing is marked by distinct volcano-sedimentary sequences overlying arc assemblages unconformably. These volcano-sedimentary sequences reflect synorogenic basins (see section 7.3) with plutonic detritus and are characterised by calc-alkaline to alkaline volcanism. Examples are found in the Wabigoon subprovince (Ayer and Davis, 1997),

318

Chapter 4: Precambrian Volcanism

Abitibi greenstone belt (Mueller and Dimroth, 1987; Dostal and Mueller, 1992), and Slave craton (Mueller and Corcoran, 2001). In the Abitibi belt (section 2.4), the 1-2 km thick, 2705-2715 Ma Hauy Formation received both volcanic and plutonic detritus from the older sequences but was also the locus of shoshonitic and calc-alkaline volcanism (Mueller and Donaldson, 1992a). The feldspar-phyric absarokites, amphibole-pyroxene-phyric absarokites, shoshonites and banakites or high-K andesites in various segments of the basin constitute up to 60% of the stratigraphy (Dostal and Mueller, 1992). The Hauy Formation flows feature thin chilled margins, locally abundant vesicles, and flow-oriented feldspars. The flows are intercalated with coarse subaerial clastic fans suggestive of eruption from a stratovolcano. In the western Wabigoon subprovince, the clastic deposits of the White Patridge Bay Group (Ayer and Davis, 1997) may be considered part of an arc unroofing phase.

Arc dissection volcanism During the terminal stages, the arc is dissected and tectonically controlled molasse basins develop. These late orogenic (see section 7.3) Archaean successor basins have the following features: (1) a pronounced unconformity with older volcanic and sedimentary rocks, (2) a bounding strike-slip fault, (3) abundant granitoid debris, (4) alluvial-fluvial deposits, and (5) calc-alkaline to alkaline volcanism (Brooks et al., 1982; Krapez and Barley, 1987; Swager et al., 1990; Thurston and Chivers, 1990; Mueller and Corcoran, 1998). The Australian Whim Creek basin is related to tectonic escape during orogeny (Krapez and Eisenlohr, 1998), but developed as a response to strike-slip faulting. Comparable depositional units are found in the Baltic shield (Ga~l and Gorbatschev, 1987) (see also section 3.9), the Zimbabwe craton (Blenkinsop et al., 1997) and the Yilgarn craton (Swager et al., 1990; Myers and Swager, 1997). In Canada contemporaneous volcanism varies considerably in these basins, ranging between 0 and 40% (Swager et al., 1990; Mueller and Corcoran, 1998). The Oxford Lake shoshonites are mafic to felsic flows with minor crystal tufts (Brooks et al., 1982) and the Crowduck Lake Group shoshonites fall into the same category (Ayer and Davis, 1997). The Stormy basin has prominent 2-30 m thick subaerial massive and brecciated tholeiitic mafic flows, local pillowed-shaped flow units as well as subaerial calc-alkaline felsic lobate and brecciated flows. The Kirkland basin displays both ultrapotassic lava flows with a blocky or an aa flow morphology, and 5-40 m thick pyroclastic surge and air fall deposits with accretionary lapilli (Cooke and Moorhouse, 1969; Mueller and Corcoran, 1998). High-level porphyry stocks characterise these Canadian molasse basins. Plume-generated greenstone belts Mantle plumes represent the convective uprise of thermally anomalous mantle and produce large volumes of mafic and ultramafic magmatism (Coffin and Eldholm, 1994) (see sections 3.2 and 3.3 for detailed discussions of plumes). Several criteria characterise plume volcanism (Wilks and Nisbet, 1988; Campbell, 2001) and include: (1) uplift prior to volcanism, leading to enhanced weathering, erosion and deposition of shallow water sedimentary units characterised by pinch outs and erosional features; (2) uplift during volcanism

4.4. Archaean and Proterozoic Greenstone Belts

319

caused by the lesser density of basaltic magma relative to mantle, seen in many oceanic plateaus; (3) subsidence associated with extraction of melt from the mantle, and cooling, leading to development of sedimentary basins; (4) dyke orientation in plume-related sequences; and (5) a predominance of lava plains and shield volcanoes. Archaean plume sequences are oceanic (Desrochers et al., 1993), intracontinental (Tomlinson and Condie, 2001 ) or arc-related (Sproule et al., 2002) in the Superior Province. Similarly, they represent an oceanic or intracontinental setting in the Zimbabwe craton (Kusky and Kidd, 1992; Blenkinsop et al., 1993; Bickle et al., 1994) and are continental in the Slave Province (Bleeker et al., 1999). The Yilgarn and Pilbara cratons have comparable plume settings (Myers and Swager, 1997; Krapez and Eisenlohr, 1998). The Steep Rock sequence (Fig. 4.4-4) in the Wabigoon subprovince Archaean succession displays an uncharacteristic assemblage of pyroclastic komatiites, the Dismal Ashrock (Schaefer and Morton, 1991), which are an integral part of a shallow water sequence, with a palaeoregolith and stromatolites overlying a 2.9 Ga granitoid basement (Wilks and Nisbet, 1988; Kusky and Hudelston, 1999). A quartz arenite-komatiite association is quite widespread in the western Superior Province and commonly overlies granitic crust unconformably (Thurston and Chivers, 1990). The transition to komatiite flows directly overlying quartz arenites has been documented in the c. 2.9-3.0 Ga Keeyask Lake deposits, Sachigo subprovince (Thurston and Chivers, 1990; Donaldson and de Kemp, 1998), as has been the komatiite-basalts-stromatolite association on the Zimbabwe craton (Nisbet et al., 1993b; Bickle et al., 1994). The > 2.8 Ga Slave craton quartz arenites overlie a basement complex unconformably at several localities (Pickett, 2002) and are inferred to be associated with ultramafic rocks (Bleeker et al., 1999). Some komatiite-tholeiite sequences display only isotopic evidence of basement (Chauvel et al., 1985), and Tomlinson and Condie (2001) used Thffa vs La/Yb relations to distinguish Archaean MORB and plume-related lithologies from subduction-related basalts. Furthermore, Fe- or Mg-rich tholeiitic basalts in greenstone belts with MORB geochemistry, a spatial association with komatiites and a lack of major felsic volcanic units may represent plume magmatism, such as the Kinojevis assemblage in the Abitibi greenstone belt. Oceanic volcanism and ophiolites Identification of Archaean oceanic crust is problematic (Thurston, 1994; Sylvester et al., 1997) and has been a contentious issue (Bickle et al., 1994; Kusky and Winsky, 1995) because of the genetic link with ophiolites, but also because of the interpretation of high strain, and of structural contacts with basement as d6collements (Kusky and Kidd, 1992; Kusky and Winsky, 1995) (see detailed discussion of ophiolites by Chiarenzelli and Moores in section 3.7). In Archaean greenstone belts, oceanic crust must be postulated on the basis of stratigraphy, volcano-sedimentary structures and geochemistry (Sylvester et al., 1997). Moores (1982) identified two types: (1) Cordilleran ophiolites, as thrust beneath continental margins, yet over a coeval continental arc, and (2) Tethyan ophiolites as thrust over passive continental margins. Sylvester et al. (1997) reviewed the variability of Phanerozoic ophiolites and the problematic aspects of Archaean ophiolites, including: (1) the scarcity of basal, variably serpentinized ultramafic units at the base of Archaean ophiolites, (2) the

Chapter 4: Precambrian Volcanism

320

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Fig. 4.4-4. Generalised stratigraphic column of an Archaean continental plume-related sequence, showing basal 3003 Ma Marmion tonalite overlain by the Steep Rock Group sequence (Wilks and Nisbet, 1988) and ultramafic pyroclastic rocks of the Dismal Ashrock (Schaefer and Morton, 1991).

variable orientation of sheeted dykes in Phanerozoic ophiolites and their resemblance to massive flows, and (3) the inherent variable vesicularity of basalts at any given water depth. Numerous oceanic tholeiites with a MORB signature occur on the margins of terranes or subprovince boundaries, are highly deformed and interleaved with turbiditic sequences. They show aspects of m61anges favouring their interpretation as oceanic crust. The inferred 2.505 Ga Dongwanzi ophiolite of the North China craton is composed of a 70 m thickness of tectonised ultramafic basal unit, succeeded by up to 5 km of gabbro, grading from a mafic-ultramafic transition zone to layered gabbro with the layering less prominent upwards (Kusky et al., 2001). The gabbro-sheeted dyke transition is grada-

4.4. Archaean and Proterozoic Greenstone Belts

321

tional. The massive flows in the ophiolite suggest mafic plain volcanism. The sheeted dyke complex is c. 5 km long and 2 km thick. Chilled dyke margins in dykes are asymmetric and variable. Pillow lavas and pillow breccia have interflow chert, banded iron-formation and pelite. Pillows commonly have 2-3 cm epidote-rich selvages indicative of hydrothermal circulation and alteration. Seamounts are mafic volcanic edifices rising from the ocean floor that are associated with mid-oceanic, back-arc and arc settings. They are 0.05-10 km thick, attain diameters up to 100 km, and are characterised by central feeders. Corcoran (2000) describes seamounts lying unconformably above a combined sequence of granitoid basement, platformal quartz arenites and iron-formation. Facies mapping revealed an inferred deep water (500-2000 m), moderate deep water (500-200 m) and shallow water (< 200 m) settings. Initial seamount development is composed of thick, pillowed units and a central dykesill feeder system (Fig. 4.4-5a). The moderate water depth facies features massive and pillowed flows plus disorganised pillow breccia and hyaloclastite with minor interstratifled shale, bedded tufts, and an evolved plumbing system (Fig. 4.4-5b) with local peperite formation. The shallow water setting (Fig. 4.4-5c) displays pillowed flows with abundant stratified pillow breccia and hyaloclastite. A change in vesicularity from 0-5% in the deep water facies to 27% in the medial section and finally 21-49% in the shallow water facies is consistent with an overall change in depth, as is the increase in hyaloclastite. The 3.2-3.3 Ga Cleaverville Formation of the Pilbara craton (Ohta et al., 1996; Fig. 4.4-6) exhibits a typical ocean floor stratigraphy. The sequence forms a series of thrustbound slices consisting of basal MORB basalts overlain by banded iron-formation (BIF), a bedded chert unit and clastic deposits. The basalts are massive sheet flows, pillow lava, hyaloclastite, reworked hyaloclastite, and feeder dykes. Pillowed flows have minor amygdules. The presence of massive sheet flows suggests either mafic plain or shield volcanism. The chert units lack terrigenous debris such as quartz, feldspar or granitic fragments suggesting an oceanic provenance. The cherts gradually increase upwards in grain size and are interbedded with fine-grained clastic rocks culminating in conglomeratic units. They are interpreted as mid-ocean ridge basalts, deep sea pelagic cherts and trench-fill turbidites. Kitajima et al. (2001) describe an Archaean oceanic crust with downwards increasing metamorphic grade for the 3.5 Ga Warrawoona Group at North Pole in the Pilbara craton The basaltic rocks are pillow basalts (> 500 m thick) with minor sheeted dykes and overlying bedded cherts (> 30 m) and minor sedimentary barite. The sequence displays three thrust-bound basalt/chert units with abundant silica dykes confined to the upper parts of basaltic units. The metamorphic grade increases stratigraphically downwards from zeolite facies to a greenschist/amphibolite transition facies. Proterozoic greenstone belts The four selected greenstone-granitoid terranes in this study include: (1) 2.27-2.05 Ga Birimian of the African craton, (2) 2.04-1.8 Ga Trans-Hudson orogen of the North American craton, (3) 2.2-1.8 Ga Svecofennian of the Baltic craton, and (4) 0.95-0.45 Ga PanAfrican of the Arabian-Nubian Shield of the African craton. Prior to opening of the Atlantic Ocean, the Birimian was probably contiguous with the Maroni-Itacaiunas mobile

322

Chapter 4: Precambrian Volcanism

Fig. 4.4-5. Evolution of an Archaean seamount (Corcoran, 2000): (a) Initial deep water seamount formation with thick pillowed units of low vesicularity. A central intrusive system with dykes and sills forms. (b) A moderate deep water setting with massive and pillowed flows as well as pillow breccia and hyaloclastite. Shale and tuff turbidites accumulate at the more distal portions of the edifice while the dyke-sill plumbing system continues to develop. (c) A shallow water sequence displays abundant massive or stratified hyaloclastite with fragments having a high vesicularity.

4.4. Archaean and Proterozoic Greenstone Belts

323

Fig. 4.4-6. Map relations of a typical oceanic assemblage based on the Cleaverville Formation (after Ohta et al., 1996). Note thrust-bound panels of pillowed flows, bedded chert and BIF with minor terrigenous sedimentary units.

belt (2.2-1.9 Ga) of the South American craton (Feybesse and Mil6si, 1994; Ledru et al., 1994). Although the tectonic setting is controversial for some areas (cf. Windley, 1992 and Stern, 1994 for the Pan-African), Trans-Hudson and Pan-African volcanoes apparently developed in oceans between older cratons or microcontinents and were deformed into greenstone belts by collision and granitoid plutonism during ocean closure (Green et al., 1985; Patchett and Arndt, 1986; Windley, 1992; Stern, 1994). Birimian and Svecofennian volcanoes, on the other hand, developed in open oceans and were accreted to craton margins (Patchett and Arndt, 1986; Boher et al., 1992; Windley, 1992; Feybesse and Mil6si, 1994; Hirdes and Davis, 2002). Proterozoic greenstone belt volcanism appears episodic with peaks at 2.2-2.1, 1.9 and 1.3 Ga (e.g., Condie, 1994b, 1995), although more continuous volcanism is found on cratons (e.g., Melezhik and Sturt, 1994) (see also sections 3.2-3.4). In Palaeoproterozoic greenstone belts volcanism typically spanned 30-95 My, and orogenic culmination was less than 100-150 My after the first widespread volcanism (Lucas et al., 1996; Nironen, 1997; Hirdes and Davis, 2002). In Neoproterozoic greenstone belts, volcanism spanned

324

Chapter 4: Precambrian Volcanism

200 My and the orogenic culmination was 220-300 My after the initiation of volcanism (Stern, 1994; Stein and Goldstein, 1996; Blasband et al., 2000).

Tectonic setting of greenstone belt volcanism Volcanism in Proterozoic greenstone belts is generally interpreted as subduction-related, mantle-derived, juvenile island arcs and back-arcs with minor spreading centre and plumerelated oceanic environments. In Trans-Hudson and Pan-African belts, arc volcanoes and back-arc basin crust were juxtaposed tectonically by accretion in an oceanic setting (Abdelsalam and Stern, 1996; Lucas et al., 1996; Zwanzig et al., 1999; Blasband et al., 2000). There is no evidence of such accretion in the Svecofennian (see also section 3.9) where Allen et al. (1996a, b) have correlated facies over long distances. Greenstone belts separated by large sedimentary basins probably represent different arcs that were either coeval or developed sequentially. For example, there could have been two coeval arcs in the Svecofennian (e.g., Pharaoh and Brewer, 1990; Windley, 1992; Billstr6m and Weihed, 1996; Kumpulainen et al., 1996; Nironen, 1997) and as many as five arcs of similar or differing ages in the Pan-African (Stoesser and Camp, 1985; Abdelsalam and Stern, 1996) (see also Frimmel, section 5.8, on evolution of the southern African Neoproterozoic terranes). The inferred tectonic setting is comparable to that of modern plate tectonics, but several key ingredients of modern plate tectonics first appear in the Neoproterozoic Pan-African orogen (e.g., Engel et al., 1980; Stern and Abdelsalam, 1998): (1) widespread ophiolites tectonically interspersed with deformed and accreted island arc sequences (Abdelsalam and Stem, 1996), (2) m61anges (Shackleton, 1994), and (3) possible blueschist facies metamorphism (De Souza Filho and Drury, 1998). Ophiolites are rare in Palaeoproterozoic terranes (e.g., Scott et al., 1992; Carlson, 1993; Peltonen et al., 1996; St. Onge et al., 1997) (section 3.7 reviews Precambrian ophiolites). In the Svecofennian and Trans-Hudson terranes, 2.04-1.92 Ga allochthonous, in part ophiolitic basalt occurs on bordering cratons rather than in greenstone belts. Palaeoproterozoic greenstone belts contain only rare high-Mg basalt and komatiitic basalt (Fox and Johnston, 1981; Attoh and Ekwueme, 1997; Leybourne et al., 1997; Lewry and Stauffer, 1997; Zwanzig et al., 1999), but the latter is common on adjacent cratons (Scoates, 1981; Hynes and Francis, 1982; Melezhik and Sturt, 1994; St. Onge et al., 1997). Composition and stratigraphy of island arc volcanoes Proterozoic arc volcanoes are dominantly tholeiitic and calc-alkalic with variable amounts of boninitic, and rare komatiitic, shoshonitic, and alkalic units. Exposed stratigraphic sections of arc volcanoes include: (1) bimodal sequences, in which tholeiitic and calc-alkalic dacite and rhyolite are intercalated with more abundant tholeiitic and calc-alkalic basalt and basaltic andesite; in some volcanoes there is an upwards change from tholeiitic to calc-alkalic lineages; in other volcanoes, the entire section is calc-alkalic; (2) bimodal sequences, in which calc-alkalic dacite and rhyolite mostly overlie more abundant tholeiitic basalt and basaltic andesite, and andesite is sparse; (3) unimodal sequences, many of which change upwards from relatively primitive, tholeiitic basalt and basaltic andesite to more

4.4. Archaean and Proterozoic Greenstone Belts

325

evolved, commonly calc-alkalic andesite, dacite, and rhyolite; and (4) unimodal sequences in which the dominant lithology is calc-alkalic andesite or dacite to rhyolite (Fig. 4.4-7; Roobol et al., 1983; Bentor, 1985; Furnes et al., 1985; Stoesser and Camp, 1985; Vail, 1985; Pallister et al., 1988; Abouchami et al., 1990; Kr6ner et al., 1991; Boher et al., 1992; Sylvester and Attoh, 1992; Schandelmeier et al., 1994; Allen et al., 1996a, b; Bailes and Galley, 1996, 1999; Lucas et al., 1996; Berhe, 1997; Leybourne et al., 1997; Maxeiner et al., 1999; Syme et al., 1999; B6ziat et al., 2000; Blasband et al., 2000). Bimodal sequences dominate whereas andesite varies from rare in bimodal sequences to abundant in some unimodal sequences (e.g., Bentor, 1985; K~ihk6nen, 1987, 1989) and in evolved volcanoes (e.g., Hirdes et al., 1996). Volcanism in many Svecofennian greenstone belts is dominantly rhyolite, dacite, and andesite, and some belts are almost entirely rhyolite (Kahk6nen, 1987, 1989; Allen et al., 1996a, b). Boninite is typically intercalated with tholeiite in the lower mafic part of volcanoes (Leybourne et al., 1997; Bailes and Galley, 1999; Wyman, 1999b), whereas shoshonitic and alkalic units along with iron-rich, MORB-like basalt lava flows were erupted late, in part related to rifting (e.g., Roobol et al., 1983; Bailes and Galley, 1999; Syme et al., 1999). The boundary between lower primitive and upper evolved sequences varies from faulted to sharp and apparently non-faulted, and reflects a change in mantle source and possibly a hiatus in volcanic activity (e.g., Stoesser and Camp, 1985). Upwards compositional changes may reflect an evolution from early immature oceanic arcs to more mature oceanic or continental margin arcs developed on thickened crust (Roobol et al., 1983; Bentor, 1985; Fumes et al., 1985; Stoesser and Camp, 1985; Pallister et al., 1988; Berhe, 1997). Lucas et al. (1996) have proposed that the more evolved sequences were deposited in basins on, or near, older tholeiitic arcs. Locally, there is evidence of possible cyclic eruptions from chemically zoned magma chambers such as intercalated rhyolite and basalt in the Trans-Hudson orogen (Bailes and Syme, 1989), and upwards progressions from rhyolite to andesite in the Svecofennian (Allen et al., 1996a). Morphology of island arc volcanoes Proterozoic arc sequences are characterised by lenticular rock units with rapid lateral and vertical facies changes (Bailes and Syme, 1989; Ayres et al., 1991; Sylvester and Attoh, 1992; Allen et al., 1996a, b; Berhe, 1997; Syme et al., 1999). The following volcanoes, which range in diameter from 2-200 km, have been identified on the basis of lithofacies studies: (1) large, basalt to rhyolite, subaqueous to emergent stratovolcanoes, some with calderas (Fig. 4.4-7) and low slopes more characteristic of shield volcanoes; (2) shallow water to emergent rhyolite--dacite caldera volcanoes, in which the calderas were 5-20 km wide and the lower part of the near-vent area was intruded by numerous subvolcanic rhyolite and dacite plutons; (3) andesite to rhyolite, intrusive cryptodome-tuff complexes with local extrusive lava flows; (4) andesite to basalt cones; and (5) basalt lava shields (Gilbert et al., 1980; Roobol et al., 1983; Baldwin, 1988; Ayres et al., 1991; Allen et al., 1996a, b; Syme et al., 1999). In places, volcanoes of various types are stratigraphically superposed to form volcanic complexes. The construction of an arc edifice commenced as a seamount that grew upwards to become an island, where the subaerial segment is indicative of inter-

326

Chapter 4." Precambrian Volcanism

4.4. Archaean and Proterozoic Greenstone Belts

327

mediate and late stages of arc volcano evolution. Some volcanoes, with sequences as much as 10.5 km thick, were erupted largely in shallow water, but an alternation of shallow and deep water, or of shallow water and subaerial environments is also observed (Fig. 4.4-7; Bailes and Syme, 1989; Ayres et al., 1991; Allen et al., 1996a, b; Leybourne et al., 1997). Island arc basalt lithofacies

Lithofacies information is best obtained from the mafic-dominated Trans-Hudson Flin Flon greenstone belt (Bailes and Syme, 1989; Ayres et al., 1991), and the felsic-dominated Svecofennian Skellefte and Bergslagen areas (Allen et al., 1996a, b). Lithofacies investigations are significant because of the massive sulphide deposits associated with these rocks. The basaltic lithofacies in both subaqueous and subaerial sections of the Flin Flon greenstone belt is amygdaloidal lava flows intercalated with mafic volcaniclastic rocks. Pillowed and sheet flows, in part capped by flow breccia, and pillowed flows grading laterally into pillow or pillow fragment breccia are common in subaqueous environments (e.g., Gilbert et al., 1980; Ferreira, 1984; Baldwin, 1988; Dolozi, 1988; Bailes and Syme, 1989; Lawrie, 1992; Sylvester and Attoh, 1992). The relative proportion of pillowed and sheet flows is variable (Bailes and Syme, 1989), possibly reflecting distance from vents, but overall, pillowed flows predominate slightly over sheet flows (Ferreira, 1984; Dolozi, 1988; Bailes and Syme, 1989). Pillowed units also form the upper part of compound sheet flows, the more distal parts of sheet flows, and rarely the base of sheet flows (Ferreira, 1984; Dolozi, 1988; Bailes and Syme, 1989). The common upwards transition from sheet to pillowed morphology in individual flows could be the subaqueous equivalent of subaerial, inflated,

Opposite: Fig. 4.4-7. Columnar sections of various Palaeoproterozoic arc volcanoes in the Trans-Hudson orogen showing eruptive environment on left, composition and lithofacies in centre, and magma lineage, where known, on right. Wedge-shaped facies represent discontinuous units. Side-by-side patterns represent interlayered facies; the width of the pattern block is proportional to the abundance of the facies. Abbreviations for eruptive environments are: A --- subaerial; CALD = caldera phase; D = water depth > 1 km; E = emergence of island; M = subaqueous, depth unspecified; S = submergence of island; Sh = water depth < 1 km; U = unconformity with regolith; and / = both environments in adjacent sections. Abbreviations for magma lineages are: B = boninite; C -- calc-alkalic; F -- iron-rich tholeiite; P -- picrite; So -- shoshonite; T = tholeiite; Tr -transitional tholeiite to calc-alkalic, and / = interlayered. Column 1 is part of a unimodal sequence that records several periods of emergence and submergence of the Amisk Lake stratovolcano, Flin Flon greenstone belt (modified from Ayres, 1977b, 1981; Ayres et al., 1991; Leybourne et al., 1997); this measured section is overlain by thick sequences of calc-alkalic andesite and dacite (Fox, 1976; Reilly, 1993). Column 2 is the shoaling Bear Lake shield volcano, Flin Flon greenstone belt. On top of the shield is a submarine caldera filled by more felsic units; late rift development is marked by eruption of iron-rich tholeiite and shoshonite (modified from Dolozi, 1988; Bailes and Syme, 1989; Syme et al., 1999). Column 3 is the subaqueous Bakers Narrows bimodal basalt-rhyolite sequence in the Flin Flon greenstone belt (modified from Bailes and Syme, 1989; Stern et al., 1995a). Column 4 is a subaerial, dominantly rhyolite-dacite, caldera-fill sequence in the Rusty Lake greenstone belt (modified from Baldwin, 1988).

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pahoehoe flows (e.g., Self et al., 1998). Cyclicity produced by both upwards variations in proportion of sheet to pillowed flows (Bailes, 1987; Bailes and Syme, 1989), and by an upwards decrease in pillow size and increase in amygdule content (Bailes and Syme, 1989) has been observed. Subaqueous sheet flows range in thickness from 0.1 to > 60 m, and average flow thickness is 4-10 m (Ferreira, 1984; Baldwin, 1988; Bailes and Syme, 1989). Sheet flows with brecciated surfaces are generally thicker than those with smooth surfaces (Bailes and Syme, 1989), possibly indicating a higher viscosity for these flows. Pillowed flows are generally thicker than all sheet flows and range in thickness from 1 to 105 m with average thicknesses of 6 to > 20 m (Ferreira, 1984; Baldwin, 1988; Dolozi, 1988; Bailes and Syme, 1989). Subaerial flows are dominantly pahoehoe with rare aa or block flows; pahoehoe toes are preserved in the upper part of some flows (Fig. 4.4-8a; Ayres, 1977b, 1982; Gilbert et al., 1980; Ferreira, 1984). Flows range in thickness from 0.6 to 10 m and average about 4 m. They differ from subaqueous flows in thinner chilled surfaces, absence of pillows, presence of pahoehoe toes, and high amygdule content with well-developed, basal pipe amygdules (Ayres et al., 1991). Many arc volcanoes have abundant mafic volcaniclastic rocks intercalated with lava flows, forming as much as 85 % of subaerial and 40% of subaqueous sequences (Fig. 4.4-7; Bailes and Syme, 1989; Ayres et al., 1991). Subaerial deposits, as much as 800 m thick, are monolithic to heterolithic, well-bedded, fall, surge, and reworked tuff and lapilli-tuff that may represent tuff cones produced by phreatomagmatic eruptions close to sea level (Fig. 4.4-8b; Ayres et al., 1991). Subaerial and subaqueous sequences are separated by monolithic, flow-foot breccia deposits (5-300 m thick) considered to represent both transgressive and regressive shorelines (Ayres et al., 1991). Subaqueous volcaniclastic rocks include fire-fountain breccias and other fall deposits with bomb sags and abundant scoria (Bailes and Syme, 1989; Allen et al., 1996b), but most are reworked, typically thick-bedded, monolithic to heterolithic tuff to tuff-breccia deposited by turbidity currents and debris flows on volcano flanks (Syme, 1988; Bailes and Syme, 1989; Ayres et al., 1991; Dolozi and Ayres, 1991). Numerous sections have repeated alternations of volcaniclastic units and lava flows. Reworked deposits, some of which are scotia-rich, were derived from subaerial or shallow-water phreatomagmatic or magmatic explosions (Bailes and Syme, 1989; Ayres et al., 1991; Dolozi and Ayres, 1991), as well as upslope collapse of pillowed lava flows (Bailes and Syme, 1989). The large volumes of explosively generated basalt tephra suggest that: (1) there were long periods of shallow water or emergent basaltic volcanism, but areas of in situ emergent basaltic volcanism were sparse because many exposed sections are flank rather than near-vent deposits; (2) once vents reached shallow water, or became emergent, intermittent but voluminous explosive basaltic volcanism was characteristic of the eruptions; (3) the phreatomagmatic nature of many explosive eruptions indicates that vents were close to sea level over long periods of time; and (4) the accumulation of repeated and thick sequences of volcaniclastic debris may indicate that submarine slopes of these ancient volcanoes were relatively gentle. Bailes and Syme (1989), Ayres et al. ( 1991 ), and Dolozi and Ayres ( 1991) have proposed that the basaltic lower part of Proterozoic arc volcanoes had a shield mor-

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p h o l o g y with low slopes, but, because of the high abundance of intercalated volcaniclastic rocks in some volcanoes, they should be termed stratovolcanoes, not shield volcanoes (Ayres et al., 1991).

Fig. 4.4-8. Photographs of various lithofacies from arc volcanoes of the Palaeoproterozoic Flin Flon greenstone belt, Trans-Hudson orogen, Canada. Pencil for scale in A, C and D is 12 cm long; coin for scale in B is 2 cm in diameter. (a) Subaerial pahoehoe basalt toes and thin flow units characterised by variable amygdule abundances, shapes, and sizes. Lower quarter of photograph is the upper part of a 40 cm thick flow unit in which many amygdules are 5-8 mm diameter; flow unit thins to left. The unit above is the bulbous termination of a 25 cm-thick flow unit that contains mm-size amygdules forming layers parallel to flow termination. Unit in right-centre, adjacent to bulbous flow unit termination and containing 5-10 cm amygdules, is part of an 85 cm long and 20 cm thick toe that extends beyond the fight edge of the photograph; upper contact of toe is the non-amygdaloidal zone on which the pencil-scale is lying. Note inclined pipe amygdules next to pencil. (b) Cross-bedding in subaerial basalt tuff of probable surge origin; beds are defined by slight variations in particle size. Discordant white units are late veins. (c) Longitudinal section of columnar joints in subaqueous rhyolite lava flow. Dark colouration at margins of columnar joints and some of cross-fractures is the result of alteration induced by ingress of sea water. Incipient brecciation has occurred adjacent to joints in the lower left of the photograph. A faint flow foliation is defined by sparse, flattened chlorite amygdules; this foliation is approximately parallel to the cross-fractures. (d) Part of a several-metres-long, lobate pillow in subaqueous, plagioclase-phyric, andesite lava flow. Areas between lobes and between this pillow and adjacent pillows are hyaloclastite composed of rounded to angular blocks in a tuffaceous matrix.

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Chapter 4: Precambrian Volcanism

Island arc rhyolite to andesite lithofacies Rhyolitic to dacitic subaqueous and subaerial rocks form domes, intrusive cryptodomes and sills, relatively short lava flows, and volcaniclastic debris (Baldwin, 1988; Bailes and Syme, 1989; Syme et al., 1999; Ayres and Peloquin, 2000). Extrusive domes and flows are common in Trans-Hudson greenstone belts (Baldwin, 1988; Bailes and Syme, 1989) but are rare in the Svecofennian where most felsic rocks are volcaniclastic or cryptodomes (K~ihk6nen, 1987, 1989; Allen et al., 1996a, b). Subaqueous rhyolite domes and flows range in thickness from 1.6 to 150 m, but many are 30-100 m thick with lateral extent uncertain because of faulting. Domes and flows form complexes up to 700 m thick with associated monolithic to locally heterolithic, hyaloclastic and pyroclastic rhyolitic units (Bailes, 1986; Bailes and Syme, 1989; Ayres and Peloquin, 2000). Domes and flows, some of which have flow layering, columnar jointing, and rare pillows, commonly have a brecciated upper zone, a massive interior, and a massive to brecciated lower zone (Fig. 4.4-8c; Bailes and Syme, 1989; Ayres and Peloquin, 2000). Locally, rhyolite lava lobes, 0.5-5 m thick and 0.5 to > 50 m long, occur within hyaloclastic and pyroclastic tuff and lapilli-tuff of the same composition. This lobe facies, which is characteristic of subaqueous, relatively shallow water rhyolite (de Rosen-Spence et al., 1980; Furnes et al., 1980; Bailes and Syme, 1989), occurs in tuff cones flanking dome and flow complexes, as a marginal facies, and as discrete units (Bailes, 1986; Bailes and Syme, 1989; Ayres and Peloquin, 2000). Cryptodomes and associated sills, up to 5 km long and several hundred metres thick, typically have brecciated margins and are distinguished from extrusive domes by the intrusive nature of the upper breccia (Allen et al., 1996b). Subaerial rhyolite and dacite flows have been described from a 7.5 km thick, calderafill sequence in the Trans-Hudson, Rusty Lake greenstone belt (Fig. 4.4-7; Baldwin, 1988). The flows are mostly 30-100 m thick, with a maximum thickness of 250 m, and have faultbounded, lateral extents of 2-4 km; they form sequences as much as 950 m thick. Most flows have 0.5-1.5 m thick, upper and lower breccia zones, but some flows are entirely breccia. Flow sequences are intercalated with rhyolitic and dacitic pyroclastic flow, fall, and surge deposits and various reworked volcaniclastic units; flow abundance decreases and abundance of reworked deposits increases upwards (Fig. 4.4-7). Regoliths associated with nonconformities have been recognised in the caldera-fill sequence by Baldwin (1988). Felsic volcaniclastic rocks are both subaqueous and subaerial and include pyroclastic fall, flow, and rare surge deposits as well as reworked pyroclastic equivalents (Roobol et al., 1983; Baldwin, 1988; Stern and Kr6ner, 1993; Allen et al., 1996a, b; De Souza Filho and Drury, 1998). Explosive eruptions, most of which were probably subaerial, were the result of magmatic and phreatomagmatic explosions (Baldwin, 1988; Bailes and Syme, 1989; Allen et al., 1996a, b). In the Svecofennian, pyroclastic eruptions produced proximal, shallow water to subaerial, unwelded to locally welded pyroclastic flow and fall deposits characterised by abundant pumice and bubble-wall shards (Allen et al., 1996a, b). Intracaldera pyroclastic flow units are as much as 1 km thick, and bomb sags were identified in some fall deposits (Allen et al., 1996a). On volcano flanks, these grade laterally into shallow to moderately deep water, medial to distal sequences of unwelded pyroclastic flow sheets and

4.4. Archaean and Proterozoic Greenstone Belts

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mass flow deposits of reworked pyroclastic debris. Reworked deposits are well bedded to poorly bedded and are composed of monolithic to heterolithic particles that range in size from ash to blocks, and include recognisable shards and pumice (Bailes, 1986; Bailes and Syme, 1989; Bailes and Galley, 1999; Syme et al., 1999). Finer units are mostly turbidity current deposits whereas coarser units are debris flow deposits (Bailes and Syme, 1989; Allen et al., 1996a). Reworked deposits vary fi'om contemporaneous with volcanism, including some double-graded units that were probably derived from a pumiceous eruption column (Allen et al., 1996b), to post-volcanism and formed by erosion, transport, and redeposition of subaerial and shallow water, unconsolidated fall and flow deposits (Bailes and Syme, 1989; Allen et al., 1996a). Andesite to dacite lithofacies include lava flows and volcaniclastic rocks erupted in subaqueous and subaerial environments (Fig. 4.4-8d; Ayres, 1977b, 1981; Maxeiner et al., 1999). A volumetrically important volcaniclastic component is poorly to moderately bedded, medium to very thick bedded, heterolithic tuff-breccia containing rounded to angular clasts as much as 2 m long. They are inferred to be debris flow deposits on volcano flanks (Ayres, 1977b, 1981; Reilly, 1993). These coarse, reworked deposits may have been derived from vulcanian pyroclastic units deposited in higher parts of the volcanoes and moved downslope by gravity sliding or slumping, possibly in stages, because of edifice instability (Roobol et al., 1983; Car and Ayres, 1991).

Subsidence of arc volcanoes and lithosphere strength The cyclic emergence and submergence of volcanoes, alternation of eruptive depths, and the great thickness of shallow water deposits indicate that (1) volcanoes rapidly subsided during upwards construction because of isostatic loading of a relatively weak and young oceanic lithosphere, (2) rates of both upwards volcano construction and downwards subsidence were variable over time, and, during periods of waning volcanism, subsidence was greater than growth, but (3) overall, eruption rates were relatively rapid such that upwards growth was generally greater than subsidence (Bailes and Syme, 1989; Ayres et al., 1991). Proterozoic arc volcanoes have a higher proportion of volcaniclastic rocks and more shallow water and subaerial volcanism than Archaean arc volcanoes. To explain this difference, Condie (1994b) proposed that Archaean volcanism was generally in deeper water than Proterozoic volcanism. Such deeper water volcanism in the Archaean could be related to higher rates of isostatic subsidence engendered by volcano loading. This implies that the lithosphere, on which arc volcanoes erupted, was initially relatively thin and weak, and the lithosphere thickened and strengthened over time; lithosphere thickness and strength is related to a number of factors, but an important factor is decreasing thermal conditions in the mantle between the Archaean and present (Richter, 1985) (see also section 3.6). Ocean-floor volcanism and oceanic plateaus Ocean-floor volcanism includes ophiolites (section 3.7) and basaltic sequences that lack sheeted dykes and an ultramafic component. Ophiolites are well developed in Pan-African greenstone belts (Fig. 4.4-9; Kr6ner, 1985b; Kr6ner et al., 1987; Pallister et al., 1988; Berhe, 1990; Quick, 1991; Shackleton, 1994; Abdelsalam and Stem, 1996) and an ophio-

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Chapter 4: Precambrian Volcanism

12

Pillowed lava flows overlain by pillow breccia

10

8

Km

Sheeted gabbro sills

Layered gabbro

m

J

i m

J

J

Dominantly dunite with some wehrilite Iherzplite and possibly harzburgite; layering; no tectonic fabric

Fig. 4.4-9. Generalised columnar section of the Neoproterozoic Sol Hamid ophiolite, Sudan (modified from Fitches et al., 1983). About 100 km to the southwest, the Onib complex, which appears to be part of the same ophiolite, contains a thin unit of isotropic gabbro above the layered gabbro, sheeted dykes rather than sheeted sills, and more pillowed lava flows with intercalated chert (KrOner, 1985). This is one of the few ophiolites in the Arabian-Nubian shield where units are not tectonically dismembered.

lite allochthon is reported from the Birimian, Nangodi greenstone belt (Carlson, 1993). In contrast, in the Trans-Hudson and Svecofennian, ophiolites are found only thrust onto adjacent Archaean cratons (Kontinen, 1987; Scott et al., 1992; Peltonen et al., 1996; St. Onge

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333

et al., 1997). Pan-African ophiolites are mostly nappes that form distinct belts between arc sequences, and between arc sequences and older cratons and microcontinents (Abdelsalam and Stern, 1996). The ophiolites, more deformed than adjacent arc terranes, are commonly dismembered and locally are ophiolitic m61ange; they are inferred to be suture zones between accreted terranes (Kr6ner, 1985b; Kr6ner et al., 1987; Pallister et al., 1988; Berhe, 1990; Quick, 1991; Shackleton, 1994; Abdelsalam and Stern, 1996). The ophiolites are not all the same age, and both Berhe (1990)and Shackleton (1994) have proposed progressive changes in age of ophiolites across the Arabian-Nubian Shield. The volcanic component of Pan-African ophiolites appears to be mostly pillowed tholeiitic basalt flows with variable chemical characteristics that include similarities to MORB, island arc tholeiites, and oceanic islands (De Souza Filho and Drury, 1998; Pallister et al., 1988; Zimmer et al., 1995). The chemical characteristics, combined with similarity in age to adjacent and overlying arc terranes, have led most workers to suggest that the ophiolites represent oceanic crust produced in supra-subduction zone settings, particularly back-arc basins, although local boninites may indicate a forearc setting for some ophiolites (Kr6ner et al., 1987; Pallister et al., 1988; Berhe, 1990, 1997; Schandelmeier et al., 1994; Stern, 1994; Wolde et al., 1996a; De Souza Filho and Drury, 1998). However, Zimmer et al. (1995) and Blasband et al. (2000) have proposed that at least some of the ophiolites are remnants of ocean floor produced at mid-ocean ridges, and Stein and Goldstein (1996) have suggested that they represent an oceanic plateau produced by a plume head. In the Trans-Hudson orogen, geochemical characteristics of some tholeiitic basalt sequences that lack sheeted dykes and ultramafic components have been used to infer an ocean-floor origin, probably in back-arc basins. These 0.3-3 kin-thick sequences are pillowed, and sheet lava flows apparently erupted in moderate to deep water (Stern et al., 1995b; Syme et al., 1999). Flows with N-MORB chemistry are dominantly pillowed whereas those with E-MORB chemistry are dominantly 1.5 to > 30 m thick sheet flows that have a higher amygdule content than N-MORB flows (Stern et al., 1995b), possibly reflecting a higher original volatile content. Layered mafic-ultramafic plutonic complexes are a common component of these sequences, but volcaniclastic rocks are rare (Syme et al., 1999). The basalt sequences are tectonically juxtaposed with arc volcanoes, but the paucity of volcaniclastic rocks in these sequences indicates that they either formed a considerable distance away from the arc volcanoes, which have a high volcaniclastic component, or that they predated explosive eruptions in arc volcanoes. Rare ocean-island basalts have also been reported (Stern et al., 1995b; Zwanzig et al., 1999). Possible oceanic plateau, basaltic assemblages have been identified from geochemical characteristics in Birimian (Abouchami et al., 1990; Boher at al., 1992), Trans-Hudson (Stern et al., 1995b), and Pan-African terranes (Stein and Goldstein, 1996). In the Birimian, the possible plateau assemblages are those identified by other workers as the lower part of primitive arc volcanoes, and in the Pan-African as supra-subduction zone oceanic crust now found in ophiolites. In the Trans-Hudson, one possible plateau has been identified; it is a 2.5-3 km thick, subaqueous sequence of tholeiitic basalt, pillowed and sheet lava flows with E-MORB characteristics (Stern et al., 1995b).

C h a p t e r 4: P r e c a m b r i a n Volcanism

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4.5.

EXPLOSIVE SUBAQUEOUS VOLCANISM

J.D.L. WHITE Under some circumstances, and despite high hydrostatic confining pressures, explosive eruptions occur in water depths exceeding a kilometre and produce substantial pyroclastic debris. These eruptions are scarcely addressed in volcanological texts, yet because of the preservation bias in favour of subwave base marine deposits in the geological record, subaqueous explosive eruptions probably exceed subaerial counterparts in volume. Important sites of modern explosive submarine volcanism, such as the coastal exposures of Honshu, Japan and offshore New Zealand, erupt mainly rhyolitic magmas. Other subaqueous settings include the deep sea where, in back-arc basins and on seamounts, gas-rich magmas are able to vesiculate strongly, even under high confining pressures (Gill et al., 1990; Fouquet et al., 1998; Hekinian et al., 2000). A critical but still poorly understood aspect of deep marine eruptions is the virtual absence of deposits unequivocally formed by explosive processes. Both experiments and thermodynamic analysis suggest that confining pressures in oceans are insufficient even below 2 km depths to prohibit explosive magma-water interaction (Wohletz, 2003). Information concerning processes of subaqueous eruptions must be gleaned from their deposits. Deposits from such eruptions form by settling from aqueous suspension or are emplaced by density currents. Subaqueous density currents bearing unmodified eruptionformed fragments may originate either directly from volcanic eruptions (i.e., pyroclastic flows and eruption-fed density currents) or indirectly by remobilisation and redeposition of material initially emplaced by a different process (Fisher and Schmincke, 1984; Cas and Wright, 1987; McPhie et al., 1993). For subaerial settings there is agreement that primary pyroclastic deposits constitute those formed as a result of eruptive fragmentation followed by single-stage transport through the ambient atmosphere. For subaqueous settings, however, even deposits formed by fragmentation followed by single-stage transport through the ambient water column have, despite the absence of any "unreworked" initial deposit, commonly been considered "reworked" (e.g., Cas and Wright, 1987; McPhie et al., 1993). Subaqueous eruption-fed deposits generally involve water-supported transport, but the transport and depositional processes are controlled by the nature of the eruption and its interaction with water. It is important that eruption-fed deposits should be distinguished from reworked deposits which may postdate an eruption by years or centuries and do not provide information about the subaqueous eruptive process. This section deals with identifying subaqueous pyroclastic rocks, many of which are Proterozoic and Archaean, explains the fragmentation process and elucidates the transport mechanisms and depositional processes.

Classification of Subaqueous Density Currents White (2000) grouped subaqueous density currents arising from eruptions into three broad categories: (1) subaqueous pyroclastic flows, (2) eruption-fed turbidity currents, and The Precambrian Earth: Tempos and Events FAired by EG. Eriksson, W. Allermann, D.R. Nelson, W.U. Mueiler and O. Catuneanu

4.5. Explosive Subaqueous Volcanism

335

Fig. 4.5-1. Summary of general categories of density-current dispersal from subaqueous eruptions (White, 2000). (3) lava-fed density currents (Fig 4.5-1). Only the former two are discussed here. Deposits formed by lateral transport of a gas-particle dispersion directly from the gas-thrust region of a subaqueous eruption are considered subaqueous pyroclastic flows in sensu stricto. Such eruptive currents require a very high particle content to maintain a density greater than water. Pumice, which in gas-supported flows has a density less than water, cannot be the sole

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Chapter 4: Precambrian Volcanism

constituent of subaqueous pyroclastic flows. Only high proportions of crystals and/or fine glass shards without vesicles increase the density sufficiently to permit subaqueous flow. Eruption-fed turbidity currents are dilute to high-concentration particulate gravity flows having water as the continuous intergranular phase and involving turbulence as a particle support mechanism. The key here is that they have been generated directly by an explosion. This group is inferred to encompass the majority of eruption-fed aqueous density currents, though vertical and/or lateral segregation may result in some parts of the currents being dominated by intergranular collisions and hindered settling. Individual clasts within such flows may remain hot during transport and deposition, but sufficient gas to exclude water from the bulk of the flowing mass is neither generated nor entrained during column collapse.

The Classic Study Area The "doubly-graded sequences" of the Tokiwa Formation, Japan (Fiske and Matsuda, 1964) are excellent examples of deposits emplaced by eruption-fed turbidity currents. They comprise extensive non-welded lapilli tufts and tuff breccias in beds from a few metres to several tens of metres thick, and are over- and underlain by fossiliferous subwave base marine deposits. Tuff and lapilli tuff consists of glassy juvenile clasts ranging in vesicularity from dense to pumiceous, with plagioclase and quartz crystals. Fiske and Matsuda (1964) identified an association of thick, internally unstratified beds formed of larger clasts (graded by fall velocity) together with interstitial ash, overlain by thin, turbiditic beds consisting entirely of ash. These thin beds show weak or no size grading, but are marked by a strong density grading with an upwards enrichment in pumice. Each thinning- and finingupwards set of thin beds coupled with an underlying thick and unstratified bed, represents a series of density currents fed from a single eruption. Each eruption was viewed as progressing from gradual expansion of an eruption column into which water was ingested, subsequent descent and then outwards flow of material in a water-supported "pyroclastic flow" to form the thick bed (Fiske, 1963; Fiske and Matsuda, 1964, p. 84). Subsequent decay of dilute regions of the aqueous convective column formed a series of increasingly fine-grained and low volume turbidity current deposits. Such water-supported "pyroclastic flows" are terminological hybrids, and their properties are best captured by the term "eruption-fed density current". Accordingly, the lower bed reflects deposition from a highconcentration turbidity current or cohesionless grain flow (Lowe, 1982; Postma et al., 1988; Manville and White, 2003). The eruptions envisaged by Fiske and Matsuda (1964) involved strong magmatic fragmentation at the vent, but were insufficiently vigorous and sustained to produce pyroclastic flows in which gas formed the continuous interparticle phase.

Subaqueous Pyroclastic Flows and Rhyolitic Pumice The styles of subaqueous eruptions are reflected in the various types of subaqueous pyroclastic deposits. The breadth of deposits is a function of the magma composition, volume

4.5. Explosive Subaqueous Volcanism

337

and magnitude of explosion. Rhyolite volcanoes are not widely reported from the modern sea floor. At various times in the past, sea levels were higher (e.g., section 7.1 and chapter 8), and products of subaqueous eruptions of evolved magma were common in island arc sequences, particularly those formed along youthful subduction systems lacking thick crust, and in deposits of deeply rifted, mature arcs. Lava pillows of dacitic or rhyolitic composition have not yet been found in modern deep sea environments, but Archaean pillows of evolved compositions have been described from several localities (de Rosen-Spence et al., 1980; Yamagishi and Dimroth, 1985). Silicic lava domes and flows are quite rare at mid-ocean ridges but appear to be more common in back-arc basins and along submerged arcs. Vesicularity of silicic rocks varies greatly, and identification of pumice at an increasing number of sites kilometres below sea level (Binns, 2003; Kano, 2003; Yuasa, 2003) suggests that most non-vesicular silicic rocks reflect significant degassing rather than an overriding hydrostatic pressure control. Explosive subaqueous rhyolite eruptions have been inferred based on geological evidence (Busby-Spera, 1984; Cashman and Fiske, 1991; Mueller et al., 1994a; Kano et al., 1996), and recent sea floor studies have interpreted caldera-related pumice deposits formed to depths of 1Y2 km (Fiske et al., 2001). In the Myojin Knoll seafloor caldera welded ignimbrite was not described (Fiske et al., 2001), and it is inferred that the water column was instead "choked" with pumice and ash during the eruption climax, which then settled under the modifying influence of simultaneously active lateral outflow from the eruption centre (Fig. 4.5-2a). The deposits are thereby produced from eruption-fed density currents, but modified by aqueous tephra fall. The latter would be expected to mute the stratification associated with dilute density current deposition (cf. Lowe, 1988; Arnott and Hand, 1989). Specific to subaqueous eruptions of low-density or low dry density tephra is for portions of the erupted tephra population to reach the water surface and float, whereas other pyroclasts sink as a result of water ingestion (Whitham and Sparks, 1986; Manville et al., 1998). The climatic Myojin Knoll eruption is only one style of explosive subaqueous rhyolite eruption on the present-day sea floor. Studies of older rock sequences have shown that proximal deposits of subaqueous rhyolitic eruptions can be deposited hot (Kano et al., 1996), and may include welded facies (Kokelaar and Busby, 1992; Schneider et al., 1992). Some other deposits, such as the Devonian Taylor Formation, California (Figs. 4.5-2b, c; Brooks, 2000), have transitional textures, with large pumice blocks of irregular shape, locally pressed conformably into one another, in an apparently unwelded matrix. Sound Archaean examples have not yet been identified.

Subaqueous Felsic Fire-Fountains Magmatic gas and steam will condense upon cooling and change the gas-supported subaqueous pyroclastic flow to aqueous eruption-fed density currents. Mueller and White (1992) described a continuous c. 80 m thick felsic Archaean sequence ranging from: (1) a massive lower division, containing large juvenile and vesicular rhyolite clasts that show soft-state deformation upon deposition, (2) a stratified division with thin beds of blocky pyroclasts, and (3) a graded bedded division composed of low-concentration tur-

3 38

Chapter 4: Precambrian Volcanism

Fig. 4.5-2. (a) Diagram indicating scale of eruption inferred to have formed the Myojin Knoll caldera on the modern sea floor near Japan (after Fiske et al., 2001 ). The eruption is inferred to have breached the surface, with tephra transported from the vent by a combination of eruption-fed density currents, aqueous suspension transport, and subaerial dispersion. The pyroclastic currents and suspension fall from the plume adds material to the top of the aqueous suspension zone. Tephra settles from suspension synchronously with density-current runout, and interacts with the currents. (b) and (c) Outcrop photos of inferred subaqueous pyroclastic flow deposit, Taylor Formation (Devonian), California. Note large size of pumice blocks, and their irregular, fluidal form. Where large clasts adjoin one another, they are mutually conformable, indicating soft-state emplacement. bidity deposits (Fig. 4.5-3a). A fountaining eruption is inferred on the basis of fluidallyshaped clasts and high vesicularity (Mueller and White, 1992). The ability of the clasts to retain heat during transport to the depositional site indicates isolation from surrounding water. Because fountaining eruptions are typified by efficient separation of particles from expanding volatiles at the vent, the exclusion of water from part of the eruption-fed current is ascribed to the generation of steam where the outer part of the hot current transfers heat to water. The non-welded matrix in the upper part of the lower division indicates that the current changed by incorporating more water at the head of the flow with time. Initially the steam, generated as heat transferred from the closely spaced particles, caused water to be almost entirely displaced from lower part of the depositing current. With time and upwards in the current, water became more prominent although larger clasts or clast pockets remained insulated by self-generated steam jackets, persisting because of the greater heat content of the large clasts. The eruption evolved from a dense magmatic fountain to a phreatomagmatic eruption, in which particle fragmentation was in part driven by water en-

4.5. Explosive Subaqueous Volcanism

339

Fig. 4.5-3. Transitions from hot to cold eruption-fed density currents. (a) Column on left shows simplified stratigraphy of Archaean pyroclastic deposits produced by a subaqueous fountaining eruption (after Mueller and White, 1992). (1) Massive deposit with hot-emplaced amoeboid clasts. (2) Diffuse contact. (3) Matrix-rich and matrix-poor layers with predominantly blocky, hydroclastically fragmented grains. (4) Sharp contact. (5) Deposits of dilute turbidity currents with low angle scouring and truncation. (6) Post-eruption suspension deposit (aqueous fall of material elutriated from the currents) overlain by iron-formation. (b) A fluid-form clast, inferred to be hot-emplaced like the amoeboid clasts from the opening phases of a subaqueous phonolitic eruption (see Martin and White, 2001). Field of view is 1.5 m wide.

tering the vent, and the pyroclasts were carried upwards convectively, entrained in an aqueous column. This sort of gradation, through a sequence from basal beds with hot-emplaced clasts to overlying strata that are better bedded and lack clasts that were emplaced while hot, is common (Fig. 4.5-3b; Martin and White, 2001). Similar gradations in emplacement temperature, recognised by changes in the palaeomagnetic character of clasts and matrix, have also been recognised in younger, less altered rocks (Kano et al., 1994).

340

Chapter 4: Precambrian Volcanism

Surtseyan Explosions: Eruption-Fed Turbidity Currents Deposits of Pahvant Butte volcano, which grew within Lake Bonneville during the late Pleistocene, illustrate the importance of eruption style in determining the nature of explosions and resulting sediment-gravity flows (White, 1996). Several tens of metres thickness of shallow-dipping (< 5 ~ beds of sideromelane tephra accumulated as Pahvant Butte grew from the lake floor, with the low dip reflecting efficient outwards transport of debris in eruption-fed aqueous currents. Most of these beds are relatively thin and show a variety of tractional current structures such as scours and cross-bedding (Fig. 4.5-4a), and reflect deposition from numerous dilute gravity currents with tractional flow-boundary zones, each reflecting a discrete eruptive pulse (White, 1996). The pulses are equivalent to the intermittent tephra jets observed subaerially during the eruption of Surtsey (Thorarinsson, 1967), but when occurring subaqueously the steam in the jets rapidly condenses, leaving concentrations of pyroclasts from the jets suspended in the water column. The pyroclast suspensions transform to vertical gravity currents that impinge on the lake floor and flow both back towards the vent and away from it as eruption-fed turbidity currents (Fig. 4.5-4b). The result is a mound of shallow-dipping, bedded tephra. Subsequent pulses initiated before deposition from a preceding current is complete, will pass shock waves through the moving currents, and temporarily inhibit ventwards flow. A variant of this process becomes active as the mound shoals to near lake surface level, and involves interaction of the density currents with surface waves. Resulting combined-flow deposits are characterised by low, broad dune forms, which increase in development and abundance upwards in the mound sequence (Fig. 4.5-4c). Although the Pahvant Butte volcano formed in a lacustrine setting, similar styles of bedding characterise submarine and englacial deposits of Surtseyan volcanism (Mueller et al., 2002a), with flat to low-angle scour cross-bedded deposits (Figs. 4.5-4d, e). Wavy to planar bedded, eruption-fed density current deposits have been inferred for the Palaeoproterozoic Ketilidian mobile belt of Greenland (Mueller et al., 2000) and the Neoproterozioc Gariep belt of Namibia (Mueller, 2003) (see also section 5.8). Similarly, the subaqueous komatiitic tuff and lapilli tufts of the c. 2.9 Ga Dismal Ashrock (Fig. 4.4-4), with vesicular pyroclasts and armoured lapilli, are probably products of Archaean eruption-fed density currents (Schaefer and Morton, 1991).

Deep-Water Subaqueous Mafic Explosive Deposits As water depth increases, and/or volatile content of erupting magma decreases, explosive eruptions become less intense. Interestingly, glassy basaltic tephra at water depths of c. 2 km have been identified in both seamount and spreading-ridge environments. Wright and Gamble (1999) reported pyroclastic rocks associated with basaltic caldera volcanoes along the Kermadec arc. Spreading ridge deposits consisting of bedded tephra with vesicular pyroclasts, and locally associated with sulphide mineralisation, appear to be in many ways analogous to deposits of much shallower Surtseyan-style eruptions (Fouquet et al., 1998; Hekinian et al., 2000; Eissen et al., 2003). Such eruptions clearly involve magma

4.5. Explosive Subaqueous Volcanism

341

Fig. 4.5-4. Model of a subaqueous Surtseyan volcano at shallow water depths (modified from White, 1996) and photographs of eruption-fed density current deposits. (a) Incipient subaqueous eruption with stratified and graded deposits. (b) Column-margin fall deposits resulting from particulate density flows. (c) Eruption breeching the surface causing the formation of combined flow deposits.

342

Chapter 4: Precambrian Volcanism

Fig. 4.5-4 (continued). (d) Stratified lapilli tufts with low angle scours. Oamaru, New Zealand. Scale, pen = 13 cm. (e) Graded lapilli tuff-tuff bed, Oamaru, New Zealand. driven into the water column by exsolving magmatic gas, and the pyroclast shapes are compatible with modification of the eruption style by interaction with ambient seawater. The extent and precise nature of this modification at depths where the water-steam transition produces expansion one to two orders of magnitude less than at Earth's surface remains to be determined. A subaqueously modified "hawaiian" eruption style (Fig. 4.5-5; Head and Wilson, 2003), perhaps with "strombolian" interludes, in which magma delivery is discontinuous and marked by discrete bursts, may adequately explain the major features of such deposits, although to date the clastogenic lava flows and welded spatter proposed by Head and Wilson (2003) remain to be identified. Doucet et al. (1994) proposed a basaltic fountain at depth for an Archaean volcaniclastic sequence, based on the eruption mechanisms suggested by Smith and Batiza (1989), and Gill et al (1990). Less known are "sheet hyaloclastites", which are layered deposits (Clague et al., 2000; Maicher et al., 2000). In addition to evidence of current-formed layering, these deposits are characterised by a mixture of non-vesicular sideromelane clasts of blocky polyhedral form with curious curved plates of glass (Fig. 4.5-6a), termed "limu" fragments (Hon et al., 1998). The polyhedral sideromelane shards are formed by cooling-contraction granulation and dynamo-thermal spalling, which require only contact of magma with water and hence are insensitive to water depth. Interpretation of the eruption mechanism for these deposits hinges, however, on explaining the nature and distribution of the limu fragments. Maicher et al. (2000) reported results from hyaloclastites from Seamount Six. Instead of a model for deep-marine "hawaiian" style lava fountaining eruption (Smith and Batiza,

4.5. Explosive Subaqueous Volcanism

343

Fig. 4.5-5. Eruption-ted density currents inferred to originate from subaqueous hawaiian-style eruption (modified from Head and Wilson, 2003). Note the prediction of clastogenic lava and agglomerate. 1989), a new model for limu and sheet-hyaloclastite formation was developed that involved modest (c. 10 fold) expansion of water droplets as they flashed to steam (Fig. 4.5-6b). Ingestion of the droplets in thin, fluid lava flows was favoured where water-saturated sediment was crossed (Maicher and White, 2001). This explanation drew specific analogies with the blowing of magma "bubbles" in littoral settings (Hon et al., 1998). The study showed that at c. 2 km water depth, the hydrostatic pressure would result in bubbles of centimetre diameter, rather than the up to 2 m diameter bubbles observed subaerially. A recent sea floor study has identified widespread limu fragments, with some accompanying polyhedral shards, dispersed in sedimentary deposits along the Gorda Ridge. These deposits are at depths below the critical depth for water, and Clague et al. (2003) infer that the limu fragments formed by bursting of basalt-glass bubbles enclosing magmatic volatiles (CO2). The bubbles are believed to form by accumulation of exsolved magmatic volatiles in shallow magma chambers where tiny bubbles join to form larger ones. Bouyant passage of these larger bubbles of supercritical CO2 fluid through the surface of lava ponded in a vent is inferred to drive the "bubble blowing" process to form limu (Fig. 4.5-6c). These types of deposits have yet to be identified in ancient sequences. An Assessment

Explosive subaqueous eruptions are both more common and more varied than hitherto perceived. Rhyolite pumice forms both during effusion and during explosive eruptions at

344

Chapter 4: Precambrian Volcanism

Fig. 4.5-6. Limu fragment from Seamount Six (a) and models for limu formation by water interaction with thin lavas (b) versus passage of magmatic bubbles through vent-ponded lava (c). In each case, "bubbles" pass through the magma/lava, stretching the surface to form thin glass bubbles, only a couple of centimetres in diameter, that break to form the curved and wrinkled limu fragments. In (b), the numbers indicate (1) rise of a bubble through a crack in thin lava crust to form a limu bubble; (2) capture of water beneath lava, boiling, and buoyant rise of steam bubble to the lava surface at "1"; (3) at rapidly advancing lava front there is no crust at all and bubble buoyancy is sufficient to deform lava skin to form limu bubble; (4) heat and resulting turbulent thermal plume from lava flow carries glass fragments from interactions into water column, aiding dispersal.

depth, and volcanological facies analysis is needed to determine the origin of any specific pumice deposit. Vesicular to non-vesicular basalt is also erupted both effusively and explosively at depths exceeding a kilometre, with explosively formed deposits known from spreading ridges and subaqueous arc calderas. Dispersion of tephra from subaqueous explosive eruption sites is quite different from that of subaerial eruptions. For instance, the height of a convective eruption plume formed subaqueously is determined by water depth rather than eruption intensity. Pumice can be buoyant for months or years after eruption, with subglobal dispersion possible even from eruptions of modest intensity (Coombs and Landis, 1966). Deposition of smaller particles initially suspended in the water column appears to be non-Stokesian, with particles travelling in negatively buoyant plumes rather than settling individually (e.g., Weisner et al., 1995; Carey, 1997). Adjacent to the eruption site, a range of density-current distribution

4.6. A r c h a e a n C a l d e r a s

345

styles is possible, varying from gas-phase pyroclastic flows, through eruption-fed dense or strongly stratified aqueous density currents carrying dispersed large and hot clasts, to dilute eruption-fed turbidity currents in which all particles are fully cooled by the time of deposition. Interpretation of ancient volcaniclastic deposits as subaqueous is best approached in steps. The first question would address the bounding facies: is the setting subaqueous? The next question is whether a given deposit or a specific bed is reworked tephra, or of a primary, eruption-fed origin. Only eruption-fed deposits provide direct information concerning the style and tempo of an eruption because reworked deposits provide little direct evidence of eruption dynamics.

4.6.

ARCHAEAN CALDERAS

W.U. MUELLER, J. STIX, J.D.L. WHITE AND G.J. HUDAK Calderas are large volcanic collapse structures with a central depression that commonly contains a resurgent dome. Smith and Bailey (1968) defined the classic piston caldera, but several other varieties exist and are a function of the collapse mechanism, including trapdoor, piecemeal, down-sag, and funnel calderas (Lipman, 1997, 2000; Roche et al., 2000). Mapping of the Valles caldera, New Mexico (Smith and Bailey, 1968) and Long Valley caldera, California (Bailey, 1989) showed that these large-scale circular to ellipsoidal structures resulted from paroxysmal eruptions, which partially evacuated shallow magma chambers. The large-scale eruptions caused collapse of the central volcanic edifice and the recurrence of explosions c. 105 years later suggest replenishment of the magma chamber. Intracontinental ash-flow calderas, with a 1-3 km intracaldera thickness, generally span < 2 My. At some convergent margin settings, best exemplified by the Taupo Volcanic Zone (Wilson, 1993; C.J.N. Wilson et al., 1995), caldera volcanoes are shortlived, erupt more frequently, and are closely spaced within extension zones (Houghton et al., 1995). Far less devastating mafic calderas are present on shield volcanoes such as Hawaii (Tilling and Dvorak, 1993) or Ambrym Island, New Hebrides arc (Robin et al., 1993). The mafic counterparts collapse due to drainage of the shallow magma chamber via rift faults and/or into secondary magma reservoirs. However, some mafic calderas, such as the Masaya caldera, Nicaragua are explosive (Williams, 1983a, b). Archaean calderas are well described because of the close association with massive sulphide deposits (Chartrand and Cattalini, 1990). They compare favourably to the mineralised deposits in the active submarine Myojin Knoll (Fiske et al., 2001) and in the Rumble volcanoes II-V (Wright et al., 1998; Wright and Gamble, 1999). The striking feature of Abitibi greenstone belt calderas is evidence for the predominance of events such as magma fountaining or high-volume effusive volcanism (de Rosen-Spence, 1976; Mueller and White, 1992; Mueller and Mortensen, 2002), rather than evidence for voluminous magma-draining explosions inherent in subaerial ash-flow calderas. In fact, de Rosen-Spence (1976) suggested that only 1% of the subaqueous volcanic rocks in the The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Allermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

346

Chapter 4: Precambrian Volcanism

Central Noranda caldera were of explosive origin. Because the strata in Archaean centres are subvertical, a cross-section of the edifice is provided. Textures are generally well preserved due to low grade greenschist facies metamorphism. Focus is placed on Archaean subaqueous examples: ( 1) the 2734-2728 Ma Hunter Mine (Mueller and Mortensen, 2002), and (2) the 2703-2698 Ma Noranda caldera complexes (Lichtblau, 1989; Gibson, 1989; Galley, 1994) of the Abitibi greenstone belt, as well as (3) the c. 2735 Ma Sturgeon Lake caldera complex of the Wabigoon subprovince (Morton et al., 1991 ; Hudak et al., 2002a, b). Hunter Mine Caldera

The felsic Hunter Mine caldera of calc-alkaline composition (Dostal and Mueller, 1996), located in the Northern Volcanic Zone of the Abitibi belt (Figs. 4.3.1-2 and 4.3.1-3a), is traceable for 35 km and spans c. 6 My (Mueller and Mortensen, 2002). The 6 km thick sequence is conformably overlain by the 2724 Ma, SRG basalt-komatiite sequence (Fig. 4.3.1-3b) which reflects the important change from subduction to mantle plume magmatism. The caldera is divided into three formational stages based on U-Pb age determinations and lithogical units: (1) a basal 2734-2730 Ma unit, (2) a middle unit, the 2732 Ma tholeiitic sill, and (3) a 2730-2728 Ma upper unit (Table 4.6-1). A model of the Hunter Mine caldera displays the formational stages and subsequent flooding of the edifice (Figs. 4.6-1a, d). The early stage is composed of lobate felsic lava flows (Fig. 4.6-2a), and dome-flow complexes with interstratified banded iron-formation (BIF), pyroclastic deposits generated by lava fountaining (Fig. 4.6-2b; Mueller and White, 1992), and graded bedded tuff and lapilli tuff deposits that collectively suggest a subaqueous setting (Fig. 4.6-1a). An extensive felsic dyke swarm, indicating the heart of the caldera, intrudes felsic flows and volcaniclastic deposits (Fig. 4.6-2c; Mueller and Donaldson, 1992b; Dostal and Mueller, 1996). Dykes locally billowed into lobate structures and small domes (Mueller and Mortensen, 2002). An up to 1 km thick, Fe-rich, gabbro sill defines the middle stage that contributed to edifice inflation (Fig. 4.6-lb). Sill emplacement occurred during incipient rifting of this arc edifice. The upper formational stage is more diverse (Fig. 4.6-1c). Numerous mafic sills/dykes, and massive-pillowedbrecciated lava flows of tholeiitic affinity are interstratified with felsic flows, abundant tuff and lapilli tuff deposits (Fig. 4.6-ld) and hydrothermal BIF and carbonate iron-formation (Figs. 4.6-1 e, f; Chown et al., 2000). The graded bedded tuff and lapilli tuff have been replaced selectively by hydrothermal fluids, and the lateral and vertical change in alteration intensity is observed at the outcrop scale (Figs. 4.6-1 d, e, f). In the heart of this intensely altered carbonate iron-formation is a low tcmpcrature massive sulphide deposit. Pumice (Fig. 4.6-1 g), euhedral and broken crystals, and minor glass shards in the tuff suggest a pyroclastic origin, but the sequence of 2-30 cm thick classical Bouma cycled turbidite beds is suggestive of remobilisation. In addition to the diverse lithological units, abundant synvolcanic faulting producing a horst and graben structure helped define the caldera structure (Figs. 4.6-1a, b, c). The hydrothermal fluids used the caldera-forming faults which commonly show silica and iron-formation emplacement. The final phase displays shale capping the succession and indicates that the edifice remained subaqueous. The upper part of the

Table 4.6-1. Characteristics of the Hunter Mine caldera, Abitibi greenstone belt Hunter Mine Group, caldera (thickness and ages) U~?l~rrfbr~notionci/ stage (0.5-2 !ir?~-rlzick) Ages: 2724.6f

::$ ;:

2727.6 i

Characteristics and petrographic Seaturcs

Interpretation, process and locus

-

3

rC

2

-... 2 (c) Sills and dykcs

(c) Aphanitic to porphyritic mafic to felsic dykcs and sills contemporaneous with edifice construction

(c) Intrusion of dykes-sills during edifice construction; feldspar-quartz-phyric dykes arc associated with late plutonic suitcs. Extension and crustal-thinning processes. Loc~fs:central part of volcanic edifice

Ma (lava flow)

(b) Flcaniclastic and ~ron-lbrmation lithofacies

(b) Volcaniclastic lithol'acies: 2-20 m thick with 2-50 cm-thick tuff and lapilli tuff turbidites (Ta or S3 bcds and Tab, Tabc, Tad, Tabcde beds); beds composed of shards. wispy vitric and angular lithic volcanic fragments, pumice, and broken and euhedral crystals: cm-thick black mudstonc capping tuft' turbidites is probably tine-grained fclsic vitric tuff. Metre-thick mudstone (shale) bcds (locally silicified) cap iron-formation. Iron-formation lithofacies: 0.2-5 m-thick units of chert-jasper, chert-magnetite and jaspermagnetite in thin beds and as m-thick folded rafts; 1-30 m-thick chert-iron carbonate (a) Felsic lithofhcies: 2-30 m-thick coherent to brecciated felsic flows with lobate terminations; extensivc flow banding; breccia units contain angular to subangular clasts with flow-banding; presence of m-thick massive to stratified lapilli tuff brcccias and lapilli tugs containing suhordinatc pumice Mafic lithofacies: tholeiitic and Mg-rich, 5080 m-thick basalt flows composed of massive, pillowed, pillow breccia and pillow fragment breccia; 2-20 m-thick massive columnar-jointed flows grade into lobate-pillowed flows and upsection into pillows and pillow hrcccia. Hetcrolithic breccia dominated by mafic clasts with minor chert, BIF and feisic clasts. Chaotic breccia 30-40 m-thick composed of large rafts of BIF, volcaniclastic blocks, and segments fclsic and mafic flows

(b) Components in tuff and lapilli tuff beds indicate pyroclastic origin; either primary or syneruptive with limited reworking. Transport mechanism: high and low-concentration density currents. Mudstonc represents calm suspension deposition, as pelagic background sedimentation and as vitric fines scttlcd through water column

Ma (fclsic dyke)

Fclsic flows arc contemporancous with Selsic dykes form upper and lower formational stage

b

?

Ma (lava flow)

::;

2728.3 f

Lithofacies and units

a

6

(a) Felsic and mafic volcanic lithofacies

p 3

Iron-formation formed by pervasive percolation of hydrothermal fluids Locus: subaqueous intracaldera setting (a) Subaqueous massive to brecciated lava flows indicative of autoclastic and hydroclastic fragmentation processes; massive to stratified lapilli tuff suggestive of synvolcanic resedimentation by subaqucous density currents Lateral and vertical flow changes in basalt flows characteristic of subaqueous flow processes; brecciation causcd by autoclastic processes and thermal granulation Hetcrolithic breccia is a mass flow deposit, and chaotic breccia reprcscnts a talus scrce deposit. Locus: subaqucous intracaldera setting adjacent to caldcra faults

w P 4

Table 4.6- 1 (continued). Hunter Mine Group, caldera

Lithofacies and units

Characteristics and petrographic features

Interpretation, process and locus

Intrusive gabbro-diorite sill (Roquemaure Sill)

E-trending gabbro-quanz diorite with a subophitic ( f ophitic) hypidiomorphic granular texture; pods with cm-scale hornblende, pyroxene, and plagioclase with interstitial quartz (sampled for U-Pb age determinations); dark green to brown weathered intrusive body; geochemically a tholeiitic ferrogabbro with FeO contents ranging from 17-21 %; early felsic dykes locally intruded unconsolidated phases of gabbro; late feldspar-quartz-phyric dykes cut sill

Early mafic intrusive phase of HMG within limits of central dyke swarm. Thick mafic sill intruding central volcanic complex; geochemical signature suggests taping of mantle-derived magma formed during arc extension-crustal thinning phase Locus: central part of edifice

(c) Felsic-dominated dyke swarm ( < 10% mafic dykes)

(c) Abundant, N-trending aphanitic to porphyritic felsic dykes occurring in eastern and western portions of the Hunter Mine Group. Dyke generations occur as multiple magma injections/pulses. Western part well documented (ca. 2.8 km thick; traceable 2.5 km up-section); eastern segment poorly exposed; combined thickness of 5-7 km for high-density dyke population

(c) High dyke density indicates dyke swarm and an extensive volcanic plumbing system. A 5 km extent of swarm suggests rifting and supports caldera formation

z d d l e fortnational stage (up to I knl-thick) Age: 273 1.8 f

i:; Ma (sill)

Lower j'orn~ationalstage ( 3 4 kaz-thick) Ages: 2729.6 f 1.4 Ma (dyke)

Dykes locally balloon into endogenous lobes and domes. QFP-dykes feed flows in upper formational stage. Locus: central intracaldera part of edifice

P

P 4

2S

8

Table 4.6- 1 (continued). Hunter Mine Group, caldera (thickness and ages) 2728.9f 0.8 Ma (lobe of dyke)

Lithofacies and units

Characteristics and petrographic fcatures

Interpretation, process and locus

(b) Volcaniclastic and iron-formation lithofacies

(b) Volcaniclastic lithofacies: ( I ) pyroclastic lithofacies (2) reworked pyroclastic and autoclastic lithofacies. Pyroclastic lithofacies divided into (i) a basal 7-20 m-thick, massive lapilli tuff breccia (ii) a middle up to 51 mthick, stratified lapilli tuff and (iii) an upper 2 m-thick turbiditic tuff-lapilli tuff. Volcaniclastic lithofacies wcll preserved in screens of dyke swarm. Pyroclastic deposits feature massive beds with irregular amoeboid-shaped clasts, stratified beds with blocky clasts and graded tuff beds; pyroclasts contain ca. 3&60 5% quuart-filled vesicles. Reworked pyroclastic and autoclastic debris is 1-5 m-thick tuff and lapilli tuff deposited in massive, graded and laminated beds Iron-formation lithofacies: 5-100 cm-thick units of magnetite and jasper-magnetite in mm-cm-scale beds and as large rip-up clasts. BIF: banded iron-formation (a) Coherent and brecciated lithofacies: Aphanitic to porphyritic rhyodacites and rhyolitcs (Si02 contents of 68-78%); coherent to brecciated flows or segments of exogenous domes (5-50 m-thick, possibly thicker); endogcnous lobes and lobate dyke terminations intrude disorganised and massive lapilli tuff breccias; m-scale lobes may display an arcuate columnar joint array near the margins; chilled lobe and flow-banded margins prominent and perlitic cracks and spherulites common

(b) Pyroclastic lithofacies originating directly from a magmatic fire fountain that was insulated from the ambient medium, watcr, by a steam carapace. Collapse of fountain caused water ingestion and hydroclastic fragmentation processes. Transport by highand low-concentration density currents. Reworked debris resulted from slope failure or tremors; transport by turbidity currents or as pelagic rain during volcanic quiescence

Dykes located in lower formational stage crosscut flows and pyroclastic deposits Age: Aphanitic fclsic flows and pyroclastic debris not favourable for U-Pb age dctcrminations, but are > 2730 Ma based on field relationships

(a) Coherent and brecciated felsic lithofacies

Iron-formation formed by pervasive percolation of hydrothermal fluids (subsurface) and diagenesis. Locus: deep-water deposits of incipient intracaldera setting (a) A complex association of lava flows and domes in situ and autoclastic breccia deposits. Lapilli tuff breccias may represent carapace and hyaloclastite breccias with their reworked counterparts; abundant interaction with water Effusive flows and domes in a composite volcanic structure. Dykes balloon into lobes and domes. Locus: central part of subaqueous volcanic edifice

2

6

5 $

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Chapter 4: Precambrian Volcanism

Fig. 4.6-1. Subaqueous Hunter Mine caldera with edifice-forming stages. (a) Early caldera-forming phase with central horst and graben structure, and pyroclastic activity. (b) Effusive lava flows with intrusion of thick mafic sill. Hunter Mine caldera is so well preserved because it remained submerged during its evolution, and because the basalts and komatiites of the Stoughton-Roquemare Group flooded the felsic edifice (Fig. 4.6-2d). Noranda CaMera

The Noranda caldera of the Southern Volcanic Zone (Fig. 4.3.1-2) is renowned for the giant 53.7 x 106 tonne Horne Mine (1927-1989; Kerr and Mason, 1990), as well as for numerous other VMS deposits (Figs. 4.6-3a, b; Kerr and Gibson, 1993). The Central Noranda camp exhibits a 20 km diameter caldera structure that is dissected by numerous NE-ENE trending synvolcanic faults affecting edifice geometry and volcanism (de RosenSpence, 1976; Lichtblau, 1989; Gibson, 1989; Gibson and Watkinson, 1990). As initially recognised by de Rosen-Spence (1976), mafic-andesitic to felsic volcanic rocks of tholeiitic to calc-alkaline composition were emplaced in a central subsidence structure, which formed due to voluminous effusive volcanism (de Rosen-Spence, 1976; Lichtblau, 1989). The Waite Rhyolite flow 8, Hdre Creek Rhyolite flow 1 and Don Rhyolite flow 5, are

4.6. Archaean Calderas

351

Fig. 4.6.-1 (continued). (c) Second caldera-forming phase with dome-flow complexes, hydrothermal iron-formation, and massive sulphide activity. (d) Flooding of the caldera by plume-generated komatiite and tholeiitic basalt volcanism of the Stoughton-Roquemaure Group. See text for details. Modified from Mueller and Mortensen (2002).

30-400 m thick and traceable for more than 10 km; they are examples of thick effusive lava flows derived from the high-level Flavrian and Powell plutons (de Rosen-Spence, 1976; de Rosen-Spence et al., 1980). The interpretation as a caldera, with displacement of lava flows along synvolcanic faults, was suggested by Lichtblau and Dimroth (1980). The setting of the caldera, based on detailed volcano-sedimentary facies analysis, ranges from below storm wave base, 200-500 m deep, to shallow water with local breaching of the edifice. Lichtblau (1989, pp. 60-72) mapped tuff beds with cross-bedding occurring in isolated sets or cosets as well as scour structures in the Powell Tuff of the Powell Formation, that are consistent with wave-induced current action and migrating wave-generated dunes. The present deep water Myojin Knoll caldera with a caldera floor at 1400 m depth, at one stage breached the surface (Fiske et al., 2001) showing that a significant depth change for subaqueous silicic calderas is not uncommon. The Blake River Group (volcanic cycle 3), hosting the Noranda caldera, has a complex stratigraphy that was resolved using a combined geochemical (Gdlinas et al., 1977b)

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Chapter 4: Precambrian Volcanism

Fig. 4.6-2. Salient features of the Hunter Mine caldera. Scale in photographs: pen = 15 cm. (a) Lobate quartz-feldspar-phyric flow displaying a chilled, glassy margin with hyaloclastite (near pen). (b) Lapilli tuff breccia deposits formed from a subaqueous fountain eruption. The highly vesicular clasts are similar to magma spatter with rounded edges and plastic deformed shapes formed during eruption and transport. (c) Columnar-jointed, feldspar-phyric dyke of felsic dyke swarm. (d) Graded (Bouma Ta), laminated (Tb), rippled to cross-bedded (To) and finely laminated (Td) very coarse to fine-grained tuff. Note incipient load casts at tip of pen. (e) Hydrothermally altered tufts. Silicified and carbonate-altered fine- to coarse-grained tuff. The fine-grained tuff is silicified (compaction and diagenesis of vitric-rich fine-grained tuff?), displays Bouma Tb-laminations and has the appearance of cherty tuff. Coarse-grained tuff is altered to Fe-rich carbonate. (f) Intensely altered tufts. These deposits are commonly referred to as chert-iron carbonate. (g) Pumice in unaltered graded bedded tuff. F, feldspar; V, quartz-filled vesicle.

4.6. Archaean Calderas

353

Fig. 4.6-3. (a) General geology of the Central Noranda Camp with numerous synvolcanic faults dividing the caldera structure into blocks (modified from Riverin et al., 1990). (b) Cross-section of Noranda and Despina calderas (modified from Gibson and Watkinson, 1990). Notice massive sulphide deposits concentrated within the subaqueous intracaldera sequence and the importance of synvolcanic faults. and volcanic facies approach (de Rosen-Spence, 1976; Lichtblau and Dimroth, 1980). The basal tholeiitic Pelletier Subgroup is the pre-caldera subaqueous basalt plain, upon which calc-alkaline assemblages developed. The calc-alkaline assemblages were volcanic complexes with the central Noranda sequence being prominent. The tholeiitic Dufresnoy Subgroup is thought to be the youngest part of the sequence but lies outside the principal caldera sequence. The Noranda caldera was divided into five intra-edifice cycles by Gibson and Watkinson (1990), based on the "Rhyolite Zones" of de Rosen-Spence (1976). Each cycle shows a differentiation trend from a basaltic andesite base to an andesite-rhyolite top (Gibson and Watkinson, 1990). P61oquin et al. (1990) recognised that the thickness of volcanic cycles varied greatly between volcanic blocks, as individual flow units could be traced across synvolcanic faults. Lichtblau (1989), Gibson (1989) and P61oquin et al.

354

Chapter 4: Precambrian Volcanism

(1990) showed that faulting was contemporaneous with volcanic construction and caldera development. The Noranda caldera displays a piecemeal confguration that is consistent with incremental collapse (e.g., Skilling, 1993). In the case of the Noranda caldera, subsidence was caused by large volume flows rather than ignimbrite eruptions. The intracaldera sequence (Fig. 4.6-3b), referred to as the Mine Sequence, representing cycle 3 of Gibson and Watkinson (1990), straddles the Flavrian and Powell blocks and overlaps the neighbouring blocks to the NW (Hunter Block) and SE (Horne Block), respectively. The c. 5 km thick caldera moat sequence is composed of the two caldera-forming successions, with a central feeder plumbing system, the Old Waite dyke swarm. Caldera subsidence ranges between 0.5 and 1.2 km, depending on bounding faults and satellite calderas (e.g., Despina) formed adjacent to the principal Noranda structure. Based on the synthesis of Gibson (1989), cycles 1 and 2 are the pre-caldera construction phases, whereas cycles 4 and 5 post-date caldera evolution (Gibson and Watkinson, 1990). The difference in displacement along caldera margin faults may suggest that this piecemeal caldera structure had a strong trap door style caldera component during the latter stages of evolution.

Sturgeon Lake Caldera The Sturgeon Lake caldera (Fig. 4.6-4) of the Wabigoon subprovince displays a bimodal, 4-5 km thick, volcanic sequence showing a pre-caldera basalt volcanic base, represented by a shield volcano, and a prominent upper rhyolitic caldera phase (Groves et al., 1988; Morton et al., 1991; Hudak et al., 2002a, b). The caldera succession compares favourably to Abitibi greenstone belt analogues, with six VMS deposits producing 18.4 • 106 tonnes of combined ore, at a grade of 8.5% Zn, 1.06% Cu, 0.91% Pb, and 3.73 ounces/tonne Ag; the VMS deposits are linked to the early and late phases of caldera evolution. Hudak et al. (2002a, b) and Morton et al. (1991) defined four sequences: (1) a 200-2100 m thick, precaldera sequence composed of mafic to felsic volcanic rocks; (2) an early, 650-1300 m thick, caldera stage composed of pyroclastic deposits with the Mattabi VMS deposit; (3) a 500-1500 m thick, late caldera stage dominated by effusive andesite-dacite flows and endogenic domes with BIF, and with volcaniclastic debris largely of pyroclastic origin; and (4) the poorly correlated Lyon Lake Fault sequence composed of basaltic-andesitic flows and volcaniclastic rocks. The intracaldera deposits are in a 25 km diameter subsidence structure. Similarly to the Noranda caldera (Lichtblau, 1989) and Joutel volcanic complex (Legault et al., 2002), the Sturgeon Lake caldera locally breached the Archaean ocean surface (Hudak et al., 2002b). The pre-caldera sequence (Fig. 4.6-5), initially interpreted by Groves et al. (1988) as an emergent mafic sequence, seems more consistent with subaqueous volcanism featuring massive to locally pillowed flows and abundant massive to poorly stratified volcaniclastic debris. The 3 km thick, intracaldera sequence has two distinct phases as in the Noranda caldera, but the Sturgeon Lake caldera displays a prevalence of pyroclastic debris, with subordinate andesite to rhyolite lava flows. Of the numerous intracaldera deposits recognised, three thick pyroclastic units stand out because of their close association with the mineralisation. The High Level Lake, Mattabi and Middle L tufts have calculated volumes

4.6. Archaean Calderas

355

Fig. 4.6-4. Geology of the Sturgeon Lake area (western part of Wabigoon subprovince) with the caldera sequence showing a piecemeal organisation as indicated by synthetic and antithetic synvolcanic faults cross-cutting the succession. Massive sulphide deposits are located in an intracaldera setting (from Hudak et al., 2002a, b).

of 16 km 3, 27 km 3, and 7 km 3, respectively (Hudak et al., 2002a, b) and thicknesses of these composite units ranges between 15 and 650 m (Fig. 4.6-5). The subaqueous Mattabi tuff pyroclastic units exhibit a distinct vertical depositional fining-upwards sequence composed of (a) a basal 10-155 m thick massive lapilli tuff, (b) a medial 6-48 m thick massive to graded lapilli tuff, and (c) a normal to inversely graded medium bedded to laminated tuff, up to 13 m thick (Hudak et al., 2002a, b); this vertical sequence permits comparison with subaqueous pyroclastic deposits (e.g., Cashman and Fiske, 1991; Cousineau, 1994). The prevalence of abundant pyroclastic debris traceable for 25 km along strike, as well as vertical facies architecture, argue that these eruptions probably occurred at a shallow depth and possibly breached the surface. The Sturgeon Lake caldera is a piecemeal struc-

Chapter 4: Precambrian Volcanism

356

ture that resembles classical subaerial, ash-flow calderas (i.e., Valles caldera; Smith and Bailey, 1968). It is significantly different from the Abitibi greenstone belt counterparts.

4.7.

COMMENTARY

W.U. MUELLER Volcanic effusive or explosive surface-forming processes that occurred in air or water have remained the same through geological time, whereas specific magma types, such as komatiites, are constrained to early Earth. The formation of arcs requires subduction processes and arcs are inferred to have been present throughout Earth's history. Archaean subduction processes must have been somewhat different from modern processes because of the inferred higher heat flow and mantle temperatures (e.g., section 3.6). The recognition of boninites in Palaeo- to Neoarchaean sequences supports the notion of flat-plate subduction (see also section 3.5), as do the prominence of TTG suites and the presence of adakites. Plume-generated magmatism (sections 3.2 to 3.4) certainly had a profound impact on volcanic assemblages and geodynamics as indicated by the ubiquity of komatiites during the early stages of Earth's history. Komatiites formed subaqueous oceanic plateaus or islands, but also penetrated stable > 2.8 Ga continental crust, as documented by the quartz arenite-komatiite association (North Caribou greenstone belt), and also affected oceanic arc sequences, as indicated by interaction with arc volcanism (e.g., Abitibi greenstone belt). Archaean and Proterozoic greenstone belts are thought to represent cross-sections of oceanic, arc and back-arc crust, as well as continental arcs and intracontinental settings. Arc-type greenstone belts are highly favourable sites for volcanic massive sulphide deposits, for which subaqueous caldera structures are important loci. Surprisingly, numerous subaqueous felsic calderas display a protracted effusive history and small-magnitude explosive eruptions with major subsidence, rather than magma-draining eruptions of subaerial ash-flow calderas. Whilst ophiolites (see section 3.7) have been a contentious issue, indications of their presence in Archaean greenstone belts is becoming more evident. Proterozoic ophiolites are not contested and are commonly associated with supra-subduction zones. Although modern settings facilitate the interpretation of volcanic processes, ancient volcanic rocks have increased our knowledge of subaqueous lava flow morphology, explosive subaqueous volcanism, and volcanic textures. Significant new results are related to komatiite flow morphology and textures. The recognition of flow inflation features, vesicle-rich komatiites, distinct flow fields composed of sheet flow and tube-shaped flows, as well as large massive flows are consistent with a low viscosity and high effusion temperatures. The association of such komatiite flow morphologies defines compound flows that are surprisingly similar to the geometry of modern pahoehoe flows. Precambrian volcanic textures such as variolites and spinifex in mafic and ultramafic flows were associated with The Precambrian Earth: Temposand Events Fxiited by EG. Eriksson, W. AItermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

4. 7. Commentary 357

Fig. 4.6-5. Selected diamond drill core sections across the caldera with distinct fining-upwards trends. The principal ash-flow tuff units display variable thickness (Hudak et al., 2002a, b) and locally host massive sulphide deposits.

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Chapter 4: Precambrian Volcanism

the cooling of superheated magmas. Both grow directly from a melt, but spinifex requires superheated magma that has a large temperature difference between liquidus and solidus. Subaqueous explosive volcanism has been neglected as an important process, generally because of limited access to modern subaqueous sites, but with the identification of Phanerozoic and Precambrian subaqueous fountaining and Surtseyan-type eruptions, a new awareness has arisen. The fact that primary subaqueous eruption-fed density currents can be recognised and apparently formed at depth, shows that constraining hydrostatic pressures, although important, does not make explosive eruptions impossible. It stands to reason, thus, that many inferred reworked subaqueous pyroclastic deposits could be of primary origin.

The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) 9 2004 Elsevier B.V. All rights reserved

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Chapter 5

EVOLUTION OF THE HYDROSPHERE AND ATMOSPHERE

5.1.

INTRODUCTION

EG. ERIKSSON AND W. ALTERMANN The previous four chapters in this book have examined the cosmic beginnings of Earth and the solar system (chapter 1), the formation of continental crust (chapter 2), and the interplay of plate tectonics and mantle superplumes during the Precambrian (chapter 3). In chapter 4, volcanism was treated as an independent variable, to study the Precambrian geological evolution of the Earth. Ultimately, the mantle and its outgassing are the source of Earth's atmosphere and hydrosphere, which form the subject of this chapter. Intimately bound to these, is biological evolution, touched upon here, but examined in more detail in chapter 6. The action of both atmosphere and hydrosphere on the surface of the Earth provide sediments (chapter 7), and eustasy (and, concomitantly, sequence stratigraphy; chapter 8) reflects the combined influences of palaeoclimate, tectonics and mantle thermal anomalies (e.g., Eriksson et al., 2001a, b). Divergence of scientific opinion is endemic to all disciplines including geology, and, in this book, we have striven to accommodate a variety of views, many of them incompatible to varying degrees. Mutually exclusive views become more apparent in study of Earth's palaeo-atmosphere and -hydrosphere. This becomes very evident in Ohmoto's overview of Archaean atmospheric and hydrospheric evolution (section 5.2), where the two major groups of theories currently endemic to this field are discussed. A more popular model (the " C - W - H - K " model) supported by a large group of proponents (e.g., Cloud, 1968; Walker, 1977; Holland, 2002; Kasting and Siefert, 2002) proposes early reducing conditions, biogenic methane as the main greenhouse gas to counteract the "faint young Sun", and the onset of an oxic atmosphere at c. 2.2 Ga. In contrast, the "D-O" model (e.g., Dimroth and Kimberley, 1976; Lasaga and Ohmoto, 2002) supports an early, single rise in atmospheric oxygen, and CO2 as the primary greenhouse gas. As pointed out by Ohmoto (section 5.2), both models use the same basic geological, palaeontological and bio-geochemical data sets to support their significantly contrasting arguments. Due to the intimate association between life, ocean and atmosphere chemistry on Earth (Ohmoto, section 5.2), biogeochemical signatures preserved within the sedimentary record enable study of the early atmosphere. As atmospheric oxygen influences the geochemical cycles of sulphur, carbon and many other elements, isotopes of carbon and sulphur are often used as proxies for inferred Precambrian palaeoredox. Thus, Lindsay and Brasier (section 5.3) examine the global carbon isotope record, based largely on Australian car-

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

bonate data, and conclude that it is far more release of Earth's endogenic planetary energy through hydrothermal and plate tectonic processes that controls such proxies, rather than atmospheric compositional changes p e r se. Similarly, in section 5.4, Trendall and Blockley provide a state-of-the-art review of iron-formation, an essentially Precambrian sedimentary rock type, and find that it provides a poor proxy for atmospheric redox, and owes its temporal control more to depository influences and the relative levels of basin floor and oceanic pycnocline. This reasoning contrasts strongly with the original "Cloud hypothesis" (Cloud, 1968), in which the deposition of major early Proterozoic iron-formations was explained as the consequence of increasing atmospheric and hydrospheric O2 content. Trendall and Blockley's analysis (section 5.4) also provides some support for Ohmoto's hypothesis (section 5.2) of an oxidised atmosphere in the Archaean already, in contradiction to the viewpoint of most other researchers. On the other hand, Lyons et al. (section 5.5), in a detailed summary of the most recent research in the field of sulphur isotopes, stress the efficacy of this method as a proxy for tracking Precambrian palaeo-atmospheric evolution and essentially support the mainstream " C - W - H - K " model. Equally difficult to explain with great confidence are Earth's two major Proterozoic glaciations, those at c. 2.4-2.2 Ga and in the Neoproterozoic. Frimmel (section 5.8) discusses in detail an example of the latter three events from southern Africa, but questions the use of carbon isotopes as palaeo-atmospheric proxies in many successions, due to their deposition within partly or wholly restricted basins. According to the "snowball Earth hypothesis" (SEH) (Kirschvink, 1992; Hoffman et al., 1998b), the Earth experienced periods of total glaciation and complete freezing of the oceans, due to lower luminosity of the Precambrian Sun and to equatorial position of continental masses, increasing the total albedo. The global ice-cover only melted when CO2 levels were considerably enhanced (greenhouse effect) by volcanic degassing, and by reduced silicate weathering and photosynthetic CO2 consumption during glaciation. The reactivation of hydrospheric circulation and exchange with the atmosphere led to a rapid deposition of the so-called cap carbonates via increased alkalinity in the oceans, because the CO2-charged atmosphere enhanced rock weathering. In this way, carbonates were deposited rapidly on glaciomarine sediments. Equally, iron, enriched in ocean waters from hydrothermal activity during glaciation periods, was precipitated as iron-formation when oxygen was introduced to the oceans after melting of the ice-cover. Thus, iron-formation is thought to be loosely associated with Proterozoic glacial deposits in the SEH. Application of the elegant "snowball Earth hypothesis" to Proterozoic glaciations is debated by Frimmel, as well as in sections 5.6 (Young) and 5.7 (Williams), with the latter two authors expressing strong criticism of various versions of the idea, supported by a large number of well-argued lines of data. Climatic and especially palaeoclimatic modelling (e.g., Rautenbach, 2001) is based partly on Earth's inferred palaeorotation, a subject examined by Williams in section 5.9. In the latter section, Williams explains the relationship between the history of the Earth's rotation and the Moon's orbit, and cyclic rhythmites of tidal origin. Palaeoclimatic reconstruction can also be based on ancient weathering profiles (palaeosols) and on sedimentary rock compositions (e.g., sections 5.10 and 5.11). Nesbitt and Young (section 5.10) stress

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

361

the importance of the interaction amongst eustasy, plate tectonics and chemical weathering, while Corcoran and Mueller (section 5.11) add consideration of depositional palaeoenvironment and lack of terrestrial vegetation to the equation of early Precambrian weathering. While discussing palaeotidal data from Precambrian rocks and their application to the orbits of the Earth and Moon, Williams (section 5.9) stresses the importance of periodic tipping of the rotational axis of the former, under the influence of lunar and solar torques and their resonances with the fluid core of Earth. These resonances may have been tied to superplume ascents from the core-mantle boundary, in their turn directly associated with the supercontinent cycle (e.g., Condie, section 3.2). Williams' message is thus the same as that of this chapter, and of the entire book, namely that geological evolution in the Precambrian (and thereafter) is tied to first-order plate tectonic-mantle processes, interacting with second-order eustatic, palaeo-atmospheric and biological influences. Lindsay and Brasier (section 5.3) emphasise the importance of the plate tectonic paradigm and of the concomitant supercontinent cycle in controlling periods of relative stasis, interspersed with shorter periods when biogeochemical proxies exhibit large changes, commonly associated with global glaciation.

5.2.

THE ARCHAEAN ATMOSPHERE, HYDROSPHERE AND BIOSPHERE

H. OHMOTO Introduction

On Earth, life, atmospheric and oceanic chemistry, and climate have been inextricably intertwined. Organisms have influenced atmospheric oxygen through photosynthesis, and essentially all of the free 02 in the atmosphere has been produced biologically. However, oxygenic photosynthetic organisms (cyanobacteria, algae and plants) are aerobes, meaning they require free 02 to produce more 02. The 02 content of the atmosphere has influenced the geochemical cycles of carbon (section 5.3), sulphur (section 5.5), and many other elements, thus influencing the composition of the oceans (e.g., contents of O2, SO]-, and Fe z+ ) and the activity of aerobic and anaerobic organisms. Earth's biota have also influenced the concentrations of atmospheric CO2 and CH4 via the production and decomposition of organic matter, and via weathering and formation of silicates and carbonates. Since CO2 and CH4 are primary greenhouse gases, life has influenced climate and the fate of organisms has been influenced by climatic conditions. A fundamental problem in earth science has been the determination of the exact links between the evolution of organisms, atmospheric-oceanic chemistry, and climate. The most important questions concern the evolution of atmospheric O2, especially the timing of and the causes of the rise of O2 and the controlling mechanisms for the atmospheric 02 level. This section will review major models of the evolution of the atmosphere, hydrosphere, and biosphere, and will evaluate critically the lines of geologic evidence used in each model. The Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Models The present oxygen budget The present atmospheric level (PAL) of 02 is 0.21 atm (= 0.21 • 106 ppm), corresponding to a total amount of 3.8 • 1019 moles of 02. This 02 has been generated by oxygenic photosynthetic organisms utilising CO2 and H20 through the following simplified biochemical reaction: CO2 + H20 = CH2Oorg + 02,

(1)

where CH20 refers to organic matter. The current production rate of 02 (and of organic matter) by photosynthesis is c. 14 x 1015 moles yr -1, c. 7 x 1015 each by terrestrial and marine organisms (Sundquist, 1985). Most of the organic matter produced by reaction (1) decomposes back to CO2 and H20 upon exposure to the atmosphere and surface water through a variety of pathways, such as the generation of organic acids and consumption by heterotrophic organisms. Important pathways are through fermentation by fermentation microbes, methane (CH4) generation by methanogenic microbes, and oxidation of CH4 by methanotrophic bacteria; these reactions may be simplified as: 2CH20 --+ CO2 -~- CH4

(2)

CH4 + 202 ~ CO2 + 2H20.

(3)

and

The current CH4 production flux is c. 3 • 1013 moles yr -1 (Logan et al., 1981), indicating that about 0.2% of newly formed organic matter is decomposing to CO2 and H20 via CH4. Note that the combination of reactions (2) and (3) yields the following overall reaction, which is the reverse of reaction (1): CH20 + 02 ~ C02 -+- H20.

(1')

This example illustrates the fact that, regardless of the decomposition pathways of the primary organic matter produced by reaction (1), one mole of O2 is required to decompose one mole of the primary organic matter. Essentially all the organic matter produced by reaction (1) is decomposed by the reverse reaction (1'), resulting in the short-term O2 consumption rate being also c. 14 x 1015 moles yr -1. At this rate, all atmospheric O2 is renewed every c. 3,000 years (= 3.8 x 1019 moles/14 x 1015 moles yr -1) by biological processes. An accumulation of atmospheric O2 molecules on a scale of more than 3000 years (i.e., the long-term O2 production) occurs because a small fraction of organic matter produced

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

363

by reaction (1) is buried in sediments, thus escaping the reverse reaction (1'). Since the burial of one mole of organic carbon generates one mole of 02, the long-term production flux of 02 equals the burial flux of organic carbon in sediments. Today, only about 0.14% of the organic matter produced in the oceans is buried in marine sediments, producing values of 1013 moles yr-l for the burial flux of organic C and the long-term 02 production flux (Holland, 1978; Lasaga and Ohmoto, 2002). This value implies that the long-term residence time of atmospheric 02 is about 4 My (= 3.8 • 1019 moles/1013 moles yr-l). The organic matter generated on land does not contribute to the long-term production of atmospheric 02 because the amounts of organic matter buried in soils and terrestrial sediments are insignificant when compared to those buried in marine sediments. Furthermore, the fluctuation in the production flux of biogenic methane does not affect the longterm 02 budget because, regardless of the pathways, most organic matter decomposes to CO2 and H20 in less than 3000 years, when exposed to 02. The long-term consumption of atmospheric 02 is mostly attributed to" (a) the oxidation of reduced volcanic gases (H2, H2S, SO2, CH4, CO), c. 0.25 • 1013 moles yr-l; and (b) the oxidation of fossil carbon (i.e., kerogen) in sedimentary rocks during soil formation, c. 0.75 • 1013 moles yr -l (Holland, 1978; Lasaga and Ohmoto, 2002). Therefore, the total long-term 02 consumption flux, c. 1013 moles yr - l , is essentially the same as the longterm 02 production flux, indicating that the present-day atmospheric 02 level is a steadystate value. When the two flux values are not balanced, the atmospheric pO2 continues to decrease or increase, resulting in an O2-free atmosphere or a runaway build-up of 02. The important issues in atmospheric evolution, therefore, include the changes through geologic time in various O2-flux values and the negative feedback mechanisms for controlling the atmospheric 02 level. These aspects are discussed in the last part of this section.

Evolution of atmospheric oxygen There are two major theories with regard to the environments and mechanisms controlling the emergence of life on Earth: (i) biosynthesis in shallow ponds near the ocean shores utilising solar energy (e.g., Miller and Urey, 1959); and (ii) biosynthesis in deep submarine hydrothermal environments utilising geothermal energy (e.g., W~.chtersh~user, 1988, 1990; Russell and Hall, 1997) (see also sections 6.2 and 6.6). The first theory implies that life originated after the formation of oceans and islands; in the second model, life may have originated before the formation of islands and continents. For oxygenic photosynthetic organisms to have become the dominant primary producers in the oceans and to generate 02 globally, the presence of a reasonable surface area of continents, perhaps larger than c. 10% of the present area, may have been necessary in order to provide the necessary amounts of nutrients (e.g., phosphate) through weathering of rocks. However, a possible role of submarine hydrothermal fluids in supplying the necessary nutrients to the primitive organisms must also be evaluated. Before the emergence of the first oxygenic photosynthetic organisms, the only mechanism for the generation of atmospheric 02 was the photo-dissociation of O-bearing gaseous molecules (H20, CO2, CO, and SO2) from volcanic gas. The estimated pO2 value for this stage, however, varies greatly among researchers, ranging from c. 10 -13 atm (Kasting,

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1987) to c. 10 -3 atm (Berkner and Marshall, 1965), depending on the values estimated for the volcanic fluxes of various gases, the hydrogen escape flux to outer space, and other parameters. A current popular theory (e.g., Kasting, 2001) favours a very low pO2 value (c. 10 -13 atm) for the early atmosphere. However, based on the minimum oxygen requirements for some key enzymes in cyanobacteria, which are aerobic organisms, Towe (1978, 1994) suggests that some level of free oxygen molecules (pO2 ~ 0.02 PAL ,~ 4,000 ppm) must have existed before the emergence of oxygenic photosynthetic organisms. A counter argument has been raised by suggesting that the earliest cyanobacteria may have been anaerobes (e.g., Kasting and Siefert, 2002). While many different models have been proposed for the evolutionary history of atmospheric oxygen, they can be classified into two contrasting groups. The first group, termed by Ohmoto (1997) the "Cloud-Walker-Holland-Kasting (C-W-H-K) model", has been advocated by these four principal researchers and their co-workers (Cloud, 1968; Walker, 1977; Walker and Brimblecombe, 1985; Holland, 1964, 1966, 1984, 1994, 2002; Rye and Holland, 1998; Kasting, 1987, 2001; Kasting and Brown, 1998; Kasting and Siefert, 2002; Pavlov et al., 2001a, b), and many other investigators (e.g., Prasad and Roscoe, 1996; Habicht et al., 2002). This model is based on the fundamental assumptions that life originated under a reducing atmosphere (e.g., Miller and Urey, 1959) and on a Darwinian concept that the formation of an oxic atmosphere triggered the emergence of eukarya (e.g., Knoll, 1992; Holland, 1994). It proposes a major change from an essentially anoxic (e.g., pO2 << 1,000 ppm) to oxic (e.g., pO2 > 0.1 PAL ~ 0.02 atm) atmosphere, at around 2 Ga. However, the proposed age of this oxygenation event has been shifted from c. 1.0 Ga (Holland, 1964) to 1.8 Ga (Cloud, 1968), between 1.9 and 2.2 Ga (Cloud, 1972; Holland, 1994), between 2.05 and 2.3 Ga (Holland, 1999), and to c. 2.3 Ga (Holland, 2002; Kasting and Siefert, 2002) (see Fig. 5.2-1a). These shifts in the proposed age of 02 rise were imposed by continuing discoveries of new geologic evidence in older rocks (discussed below). Although this group of investigators has come to share a common hypothesis regarding the rise of 02 around 2 Ga, large differences exist in the proposed pO2 values for the pre-rise stage. For example, the proposed pO2 value at c. 2.5 Ga varies from c. 10-13 atm (Kasting, 1987), to c. 10 -8 atm (c. 0.01 ppm; Pavlov and Kasting, 2002), to c. 10 ppm (Holland, 2002), and to c. 10 -3 PAL (c. 200 ppm 02; Rye and Holland, 1998). Recent models by Holland's and Kasting's groups propose an earlier major rise of atmospheric 02 from < 10 ppm to c. 10 ppm at c. 3.0 Ga (Rye and Holland, 2000) or from < 0.01 ppm to c. 0.01 ppm at c. 2.8 Ga (Kasting and Siefert, 2002), as well as a third major rise from c. 0.5 PAL (c. 0.1 atm) to c. 1 PAL at around 600 Ma (e.g., Canfield and Raiswell, 1999) (Fig. 5.2-1 a). The second group of models, termed by Ohmoto (1997) the "Dimroth-Ohmoto (D-O) model", was first proposed by Dimroth and his associates (e.g., Dimroth and Kimberley, 1976; Dimroth and Lichtblau, 1978). It has been supported strongly by Clemmey and Badham (1982) and refined by Ohmoto and his co-workers (e.g., Ohmoto, 1992, 1996b, 1997, 1999; Ohmoto and Felder, 1987; Ohmoto et al., 1993; Watanabe et al., 1997, 2000; Lasaga and Ohmoto, 2002). This model postulates the emergence of oxygenic photo-

365

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

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Age (Ga) Fig. 5.2-1. The Cloud-Walker-Holland-Kasting model for the evolution of atmospheric chemistry (a) and Canfield's model for the evolution of ocean chemistry (b). Fig. 5.2-la: the grey area (a) represents the range of pO 2 suggested by Holland (1966). Curve (b) is the pO 2 evolution curve by Rye and Holland (1998), and (c) is the pO2 evolution curve by Kasting (1987, 2001). Curves (d), (e) and (f), respectively, represent the pCO2, pH2 and pCH4 values suggested by Kasting (2001) and Pavlov et al. (2001a). Fig. 5.2-1b: the evolution curves for the sulphate, sulphide and Fe 2+ contents of the oceans as suggested by Canfield and Raiswell (1999), Bjerrum and Canfield (2002) and Habicht et al. (2002).

Chapter 5: Evolution of the Hydrosphere and Atmosphere

366

synthetic organisms (cyanobacteria or their precursors) shortly after the formation of the oceans and continents (c. 4 Ga) (see also sections 2.8, 3.2 and 3.6), one major rise of pO2, from < 10 -3 PAL to c. 1 PAL shortly after, and an essentially constant pO2 level (within +50% of PAL) since then (Fig. 5.2-2a). An important implication of this model is that the necessary pO2 condition for the emergence of eukaryotes was already created by c. 4 Ga.

Evolution of atmospheric C02 and CH4 Sagan and Mullen (1972) expressed the fundamental problem in climate evolution on billion-years time scales: Earth's hydrosphere must have been completely frozen prior to 2.0 Ga, if the Sun's luminosity evolved as predicted by the solar physics theory (e.g., Newman and Rood, 1977), and if the atmosphere had maintained it's present balance of greenhouse gases during Earth's history. Yet, the geologic record indicates the presence of flowing water on the land surface and of a fluid ocean capable of continuously supporting life since at least 3.5 Ga. Their (Sagan and Mullen, 1972) solution to this "Faint Young Sun Paradox" (FYSP) was to invoke large abundances of the greenhouse gases methane and ammonia within the early atmosphere. Others (e.g., Henderson-Sellers, 1979), suggested that the early Earth might have had a very low albedo, leading to higher surface temperatures. Atmospheric modelling by Kuhn and Atreya (1979) and Kasting (1987) demonstrated that the chemical mix proposed by Sagan and Mullen (1972) was unstable and unlikely to persist. Moreover, the albedo required by the Henderson-Sellers (1979) model was probably unrealistically low. Owen et al. (1979) and Kasting (1987) argued that an atmosphere strongly enriched in CO2 was the likely solution to the FYSE According to Kasting's (1987) climatic model, atmospheric pCO2 gradually decreased from c. 1 atm (c. 3,000 PAL) at 4.5 Ga to 0.1 atm (c. 300 PAL) at 2.5 Ga, and to c. 10 -2 atm (c. 10,000 ppm: c. 30 PAL) at 1.0 Ga (Fig. 5.2-2a). Kasting, Holland, and their associates (e.g., Rye et al., 1995; Rye and Holland, 2000; Pavlov et al., 2001b; Kasting and Siefert, 2002) are currently advocating that biogenic methane was the primary greenhouse gas prior to the rise of O2 at c. 2.2 Ga. They suggest values around 1000 ppm CH4, c. 3,000 ppm CO2 (i.e., only c. 10 PAL, instead of c. 300 PAL), and much less than c. 0.1 ppm O2 for the atmosphere at c. 3.0 Ga (Fig. 5.2-1 a). This subject will be discussed further below. The atmosphere cannot contain high amounts of both CH4 and 02, because they will react to form CO2 and H20 by photochemical reaction: CH4 + 202 ~ CO2 4- 2H20.

(4)

However, when the biologic production rate of CH4 and/or 02 is higher than the reaction rate of equation (4), the atmosphere can contain appreciable amounts (c. 1 ppm) of CH4 with high pO2, such as in today's atmosphere, or c. 10 ppm 02 with high CH4 contents as proposed by the Kasting and Holland research groups for the Archaean. In contrast, the Dimroth-Ohmoto model implies that atmospheric CH4 content has been low (c. 1 ppm) and CO2 has been the major greenhouse gas since c. 4.0 Ga (Fig. 5.2-2a).

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

367

Fig. 5.2-2. The Dimroth-Ohmoto model for the evolution of atmospheric chemistry (a) and ocean chemistry (b).

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Evolution o f ocean chemistry

As deceased biomass sinks through the water column, it decomposes by reactions with dissolved 02 in the water. This process causes (semi-) closed water bodies to become stratified with respect to 02: i.e., an oxygenated surface layer is underlain by an anoxic water body, as in the Black Sea. Modern ocean profiles typically exhibit oxygen minima at depths around 500 m (Drever, 1997), but the oceans basically remain oxygenated because of the deep circulation of O2-rich surface waters from the polar regions. The 02 content of the bottom water, therefore, depends on the organic productivity in the surface oceans and on the 02 content of high latitude surface water, which in turn depends on atmospheric pO2 and the temperature and salinity of surface water (Sarmiento, 1992; Lasaga and Ohmoto, 2002). In general, when atmospheric pO2 exceeds about 0.5 PAL, the bottom ocean water will remain oxygenated; below c. 0.5 PAL, the bottom water will become anoxic. Conversely, any geochemical data indicating that some Archaean sediments were deposited under deep (> c. 500 m) oxygenated water conditions suggest that the atmospheric pO2 was greater than c. 0.5 PAL (e.g., Lasaga and Ohmoto, 2002). The atmospheric 02 evolution models of Kasting and Holland (Fig. 5.2-la), therefore, imply that Earth's entire oceans, except for the photic zone (< c. 100 m), remained anoxic until c. 600 My ago. In contrast, the Dimroth-Ohmoto model (Fig. 5.2-2a) suggests that Earth's entire oceans, except for semi-closed local basins (e.g., the modem Black Sea), have remained basically oxygenated since c. 4.0 Ga. The sulphur chemistry of the ocean is closely linked to the evolution of sulphatereducing bacteria (SRB) and the atmospheric pO2 level (see also section 5.5). Under an anoxic atmosphere, the SO ] - content of the ocean water is expected to be much lower than the present value of 28 mM (900 ppm S) (section 5.5). Walker and Brimblecombe (1985) have estimated that the SO42- in the Archaean oceans was generated only by photochemical reactions of volcanic SO2, with concentrations less than c. 1/30 of the present value, i.e., less than c. 1 mM. A recent proposal by Canfield and his group (Canfield and Teske, 1996; Canfield and Raiswell, 1999; Canfield et al., 2000; Bjerrum and Canfield, 2002; Habicht et al., 2002) suggests that the oceanic SO 2- content remained at c. 200 laM, except in local evaporitic basins, until about 2.2 Ga, then gradually increased to c. 10 mM approximately at 800 Ma, when the second step-wise increase to the present level occurred. In contrast, the Dimroth-Ohmoto model proposes essentially the same SO ] - content in the oceans since c. 4.0 Ga (note discussion of an essentially opposite point of view in Lyons et al., section 5.5). At temperatures below c. 250~ aqueous solutions cannot contain high concentrations of both Fe 2+ and H2S due to the formation of iron sulphides (e.g., Walker and Brimblecomb, 1985; Ohmoto and Goldhaber, 1997). Walker and Brimblecomb (1985) and Bjerrum and Canfield (2002) suggest that the oceans prior to c. 1.8 Ga contained high concentrations of Fe2+(c. 0.1 to c. 1 mM) but low H2S (< 0.1 mM). Bjerrum and Canfield (2002) also proposed a major decrease in Fe 2+ accompanied by an increase in H2S since c. 1.8 Ga, to create H2S-rich and Fe2+-poor anoxic global oceans between 1.8 Ga and c. 800 Ma. The second rise of 02 to > 0.5 PAL at about 0.6 Ga changed the global oceans to SO]--rich and H2S-poor (Fig. 5.2-1 b) (see also section 5.5). In contrast, the Dimroth-

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

369

Ohmoto model proposes that the contents of both Fe 2+ and H2S in the normal oceans have remained low (< 0.1 mM) since c. 4.0 Ga, except in local anoxic basins, where submarine hydrothermal activity supplied high concentrations of Fe z+ or where high concentrations of HzS were produced by SRB (Fig. 5.2-2b) (see also, discussion of the Fe-stratified Archaean ocean model; Trendall and Blockley, section 5.4).

The Geological Evidence Different models for the evolution of atmospheric and oceanic chemistry and the evolution of organisms have developed from differences in the interpretation of the geological, palaeontological, and (bio-)geochemical records. These different interpretations consider: (1) the fossil record and the evolution of major organisms (see also sections 6.2 and 6.3); (2) the significance of unstable minerals (uraninite, pyrite, and siderite) in fluvial sedimentary rocks; (3) the behaviour of Fe in subaerial environments (palaeosols, laterites, and red beds); (4) the behaviour of Fe in marine environments (banded iron formations, volcanogenic massive sulphide deposits, and pillow lavas) (section 5.4); (5) the evolution of sulphur-utilising bacteria (SRB, sulphide-oxidising bacteria, sulphur-disproportionating bacteria) and the geochemical cycle of sulphur (6348, and 633S records of sulphides and sulphates) (section 5.5); and (6) the geochemical cycle of carbon (613C records of organicand carbonate carbons in sedimentary rocks) (section 5.3).

The fossil record and the evolution of major organisms Oxygenic photosynthetic organisms. Among oxygenic organisms living today, one of the oldest lineages in the tree of life is cyanobacteria. Major scientific questions concerning cyanobacteria include the time they first appeared on Earth, and whether or not they were the first oxygenic photosynthetic organisms. For many years, most geologists have accepted the "oldest microfossils" reported by Schopf (1993) from the 3.45 Ga Apex chert in the Pilbara district of Australia to represent remnants of cyanobacteria. However, Brasier et al. (2002) have raised questions as to whether these "microfossils" may represent the products of non-biological chemical reactions in submarine hydrothermal environments (cf. Schopf, section 6.2, for a different viewpoint). Probably all geologists accept the microfossils in the c. 2.6 Ga Campbellrand Subgroup, South Africa (Altermann and Schopf, 1995; Kazmierczak and Altermann, 2002) as cyanobacteria, and those in the 2.1 Ga Negaunee Iron Formation (Michigan, U.S.A.) as eukaryotic algae (Hahn and Runnegar, 1992). Holland's and Kasting's groups have used this fossil evidence to support their models for a major rise of pO2 at c. 2.2 Ga (e.g., Holland, 1994; Kasting, 2001). Biomarkers (molecular fossils) of eukaryotes and cyanobacteria have been found in the 2.7 Ga Jeerinah Formation of the Hamersley basin, Australia (Brocks et al., 1999). Most modern eukaryotes are aerobic. Some anaerobic eukaryotes exist but they evolved apparently quite recently through the loss of mitochondria (Williams et al., 2002). According to an experimental study by Jahnke and Klein (1979), aerobic eukaryotes require more than 0.01 atm pO2 (i.e., > 5% PAL). The biomarkers in the Jeerinah Formation are interpreted

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

as remnants of planktonic cyanobacteria and eukaryotic algae that lived in the photic zone at locations more than 50 km away from the contemporaneous shorelines, and which were deposited in the oceans at depths greater than 400 m, rather than in a so-called "oxygen oasis" on the shores (R. Summons, 2000, pers. com.). Therefore, it may be reasonable to assume that the atmospheric pO2 level 2.7 Ga ago was already greater than 0.01 atm. However, a major concern of organic geochemists and palaeobiologists is whether these biomarkers in Archaean sedimentary rocks represent the remnants of organisms that were buried with the host sediments or those, introduced much later as petroleum contaminants (e.g., Brocks et al., 1999; see also section 6.2). The carbon isotopic fractionation factor between CO2 and oxygenic photosynthetic organisms (ACO2-org - - 6 1 3 C c o 2 - - 613Corg) is typically around 18%c, resulting in about a 25%o difference between the 613C values of marine carbonates and marine organic matter. The 613C relationships between carbonates and organic matter of many Archaean sedimentary rocks are very similar to those of Phanerozoic age (Fig. 5.2-3), although the variations are quite large. Rosing (1999) reports 613C values of c. -25%0 for kerogen from > 3.7 Ga organic carbon-rich shales in the Isua district, Greenland. This is the oldest evidence of organisms in the oceans. However, it is uncertain whether the organic matter in these sediments reflects remnants of cyanobacteria or other organisms. This is because many other organisms, including some methanogens, utilise the Calvin cycle to fix carbon and possess similar 613C values (Schidlowski and Aharon, 1992). In order to solve the ambiguities concerning the nature of microfossils and biochemical fossils in Archaean sedimentary rocks, future research should be directed towards finding biomarkers in sedimentary rocks older than 2.7 Ga and developing new methods to study the biochemistry/metabolism of Archaean organisms.

Methanogens and methanotrophs. Biogenic methane, produced through reaction (2) in anoxic environments by methanogenic microbes (anaerobes) utilising the remnants of other organisms (e.g., cyanobacteria), is characteristically very depleted in 13C; the 613C values are typically 40-80%o lighter than the organic matter (Schidlowski and Aharon, 1992; Hayes, 1994). Recycling of biogenic methane back to CO2 may be carried out through reaction (3) by methanotrophs (aerobes) in an oxic water zone above an anoxic water body (e.g., Hayes, 1994). Cyanobacteria and other primary producers that utilised the recycled CO2 would also exhibit very negative 313C values. Therefore, the presence of organic matter with 613C values less than about -35%0 in marine sediments of about 2.8 Ga (the Tumbiana Formation in Australia) was suggested by Hayes (1994) as evidence in favour of the appearance of methanotrophs and of the development of stratified oceans (see also section 5.4), where methanogens were actively producing methane in the lower anoxic water body and cyanobacteria and methanotrophs were active in the overlying oxic zone. Kasting's group suggests that the dominant primary producers in the oceans before 2.8 Ga were methanogens directly utilising the atmospheric CO2 and H2 for biosynthesis and methane production through the following reactions:

CO2 -~- 2H2 = CH2Oorg + H20

(5)

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

371

and CO2 --[-4H2 = CH4 + 2H20.

(6)

Organic matter produced by reaction (5) may be expected to possess 613C values around -25%0 (cf. Schidlowski and Aharon, 1992). At about 2.8 Ga, cyanobacteria became the dominant primary producers, allowing methanotrophs to emerge and recycle CH4 to produce organic matter with 613C values < -35%o; the biogenic methane production rate also increased to produce a methane-rich atmosphere (Pavlov et al., 2001 a; Kasting and Siefert, 2002). However, younger sediments with similarly low 613Corg values, such as the 2.0 Ga Franceville Formation in Gabon (Gauthier-Lafaye et al., 1996) and the modern Black Sea sediments (Michaelis et al., 2002), demonstrate that such low 613C values do not necessarily require a methane-rich and O2-poor atmosphere. In the absence of O2, recycling of methane to CO2 may be carried out by sulphatereducing microbes (Bacteria and Archaea) by utilising SO 2-, as are seen in modern euxinic basins (e.g., Hinrichs et al., 1999; Orphans et al., 2001): CH4 -k- SO 2- -+- 2H + --CO2 -+- H2S -+- 2H20.

(7)

Hinrichs (2002) suggests that the low 613C -organic carbon at 2.8 Ga (see also section 5.3) represents the time of appearance of sulphate-reducing microbes. However, based on the sulphur isotope record (see also section 5.5), some researchers propose that the appearance of sulphate-reducing microbes dates back to at least 3.4 Ga (Ohmoto et al., 1993; Shen et al., 2001). Reaction (7) would not have been important in the Archaean oceans, if the seawater contained only c. 1/100 of the sulphate of the modern oceans, as suggested by Habicht et al. (2002). Watanabe et al. (1997) have recognised that the number of organic carbon samples with 613C > -35%0 by far exceeds that with 613C < -35%o in geologic units of all ages (Fig. 5.2-3). They proposed that most organic matter with 313C > c. -35%0 was deposited in oxic oceans, while the < -35%0 matter was deposited in local, restricted anoxic basins; their suggestion is compatible with the D-O model for atmospheric evolution (Fig. 5.2-4).

The significance of unstable minerals (uraninite, pyrite and siderite) in fluvial sedimentary rocks Uraninite (UO2) is economically the most important uranium-bearing mineral. Uraniniterich ore deposits are hosted in: (1) pegmatites and quartz veins, (2) quartz-pebble conglomerates, and (3) sandstones; this group is often divided into the roll-front type (3-1) and the unconformity (3-2) type. Uraninite crystals in group (1) formed by hydrothermal fluids at T > c. 200~ Those in the roll-front type most likely formed by reactions at nearsurface temperatures between (i) shallow oxygenated groundwater that leached uranium from tufts and sandstones and (ii) reductants (e.g., petroleum, hydrocarbon gases, and kerogen). Uraninites of the unconformity type are thought by many geologists to have formed

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Fig. 5.2-3. Carbon isotope records of carbonates (a) (Shield and Veizer, 2002) and organic matter (b) (Pavlov et al., 2001b; Yamaguchi, 2002).

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

Fig. 5.2-4. Schematic illustrations of the Archaean w o r l d according Cloud-Walker-Holland-Kasting model (a) and the Dimroth-Ohmoto model (b).

373

to

the

by the mixing of reducing hydrothermal fluids ( T > 200~ and shallow U-enriched oxic groundwater (e.g., Guilbert and Park, 1985). The Oklo uranium deposits in Gabon, famous for natural fission reactors, formed at about 2.0 Ga (Gauthier-Lafaye et al., 1996). A current popular model postulates that they are the oldest sandstone-type U deposits (i.e., group 3-1); thus they are considered as evidence for the rise of 02 at around 2.2 Ga (Holland, 1994). They are also compatible with the D-O model. An important question for future research is whether or not similar types of deposits formed prior to c. 2.2 Ga.

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

Since Holland's (1964) paper, group (2) uranium deposits (i.e., those hosted in alluvial quartz-pebble conglomerates) have been used as strong evidence for an anoxic atmosphere prior to c. 2.2 Ga, because almost all the major deposits are > 2.2 Ga in age and contain uraninite and pyrite, that are not stable under an oxic atmosphere. The quartz pebbles are poorly sorted with their size varying from a few mm to c. 10 cm in diameter (Fig. 5.2-5a). They also contain variable-sized pyrite pebbles (mm to cm scale) and pyrite crystals (c. 10 lain to several mm). The uraninite grains are typically < c. 100 lam in size. Many previous researchers (e.g., Holland, 1984; Minter, 1999; Rasmussen and Buick, 1999; England et al., 2001, 2002) have suggested that these uraninite and pyrite grains are detrital in origin, derived from the chemical weathering of uraninite- and pyrite-rich granite pegmatites (i.e., derived from group (1) deposits above). The abundance of Th-rich uraninite in these deposits has been important evidence for the detrital model, because Th has generally been regarded as immobile in low temperature conditions. Other researchers have argued, on the contrary, that: (a) most, if not all, of the uraninite and pyrite grains were formed by groundwater and hydrothermal fluids during or after the deposition of the host sediments (Barnicoat et al., 1997); (b) thorium-rich uraninite can form from organic acid-rich groundwater (Dimroth and Kimberley, 1976); and (c) even if there are some detrital grains of uraninite and pyrite in these sediments, they do not necessarily reflect the atmospheric oxygen level. For example, Phillips et al. (2001) have suggested that the "round pyrite pebbles" were products of the sulphidisation of detrital pebbles of iron pisolites after sediment deposition. Iron pisolites, composed initially of concentric layers of ferric (hydr)oxides, are common in Phanerozoic iron formations; they probably formed by wave action under an oxygenated atmosphere. Therefore, Phillips et al. (2001) also suggest that pre-2.2 Ga quartz pebble conglomerates actually hold evidence for an oxic, rather than an anoxic atmosphere. Although, some Witwatersrand pyritic conglomerates exhibit all the sedimentological properties of pyrite detritus, including imbrication of elongated, rounded clasts, of up to a few cm across, embedded in dark shale matrix (Fig. 5.2-5b), and geochemically heterogeneous sulphur composition, a very serious problem with their interpretation as "pyrite clasts" arises. Experiments (Ohmoto, unpublished data) to form large (> 1 cm) round pyrite pebbles in a tumbler, were not successful, because pyrite crystals are too brittle. Large pyrites in hydrothermal ore deposits are typically aggregates of smaller crystals, and thus easily disintegrate into smaller pieces during fluvial transportation. Furthermore, examination of many "round pyrite pebbles" using reflected and transmitted light microscopy and electron microscopy reveals that replacement textures are the norm. The shapes of "pyrite pebbles" may simply represent the shapes of the precursor pisolites, cherts, and BIF fragments. Independently from the ongoing discussion, detrital grains of uraninite and pyrite have also been found in Phanerozoic conglomerates at several localities (Dimroth and Kimberley, 1976); the most famous occurrences are modern fluvial sediments of the Indus River in Pakistan where the accumulations of detrital grains of uraninite have reached subeconomic sizes (Simpson and Bowles, 1981). A very important aspect of the occurrences of detrital grains of uraninite, pyrite, and siderite, as first suggested by Dimroth and Kimberley (1976) and later supported by

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

375

Fig. 5.2-5. (a) Hand specimen photo of a uraniferous quartz-pebble conglomerate from the Stanley Mine, Elliot Lake area, Ontario, Canada. (b) Polished slab of a "pyrite clasts in shale matrix" conglomerate from the Witwatersrand Supergroup, South Africa. Clemmey and Badham (1982) and Ohmoto (1997), is that they are mostly restricted to poorly sorted fluvial sediments of all geologic ages (but mature compositionally), yet they are absent in other sediments (alluvial and marine) and in soils of all geologic ages (Fig. 5.2-6). For example, in search of detrital pyrite in many Archaean cherts and shales, Kojima et al. (1998) were able to find only a few samples containing "possible detrital" pyrite grains. Ohmoto (1999) suggests that the surviving grains of unstable heavy minerals in fluvial and marine sediments were initially hosted in less-common rocks (e.g., quartz veins, massive sulphide ore bodies), rather than in normal feldspathic igneous rocks, and were protected from chemical weathering; these grains were liberated by fragmentation and abrasion of the host rocks during flood transportation caused by storms or by glacial ablation, and they were quickly covered by subsequent flood sediments (see Fig. 5.2-6). Therefore, the surviving grains of these unstable heavy minerals in unusual sedimentary rocks may not be connected to the atmospheric 02 level. On the other hand, the absence of these detrital minerals in most sedimentary rocks (excluding texturally immature fluvial deposits) may be evidence in favour of an oxic atmosphere since at least 3.5 Ga ago (Fig. 5.2-6).

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Fig. 5.2-6. Schematic illustrations explaining the presence or absence of pyrite and uraninite crystals in fluvial sediments and soils.

The uranium deposits in quartz-pebble conglomerates of Archaean age are generally much larger than those of younger age. Archaean placer-gold deposits are also generally larger than Phanerozoic deposits. An intriguing question is whether the size of these ore deposits was controlled by: (a) atmospheric chemistry, or (b) tectonic and igneous environments (e.g., geothermal gradients) for the formation of parental bodies (e.g., veins) that hosted the gold, uraninite, and pyrite crystals. The answer is probably (b), as will be discussed in the section on banded iron-formations, below. The behaviour of Fe in subaerial environments (palaeosols, laterites and red beds) In the absence of free O2 molecules, ferrous iron (Fe z+) in silicate minerals and glass is expected to dissolve, like magnesium. The dissolution rates depend mostly on pH, temperature, and the ratio of surface area of solid phase to the mass of water (e.g., Lasaga, 1998), applicable in the following reaction: Fe2SiO4 + 4 H 2+ --+ 2Fe 2+ + H4SiO4 . (olivine) (silicic acid)

(8)

However, in the presence of free 02 molecules, dissolved ferrous iron is converted to ferric (hydr)oxides (goethite and hematite) that are very insoluble: Fe 2+

+ 102 -k- 5H20 ~ Fe(OH)3 + 2H + (goethite)

(9)

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

377

or

2Fe 2+ + 89 + 2H20--+ Fe203 + 4H + .

(10)

(hematite)

Therefore, soils that formed under an anoxic atmosphere may have lost their iron during soil formation; the soil colour may become grey to white. In contrast, soils that formed under an O2-rich atmosphere are expected to retain y~'Fe and to show an increase in the Fe3+/Ti and Fe3+/Fe 2+ ratios and a decrease in the Fe2+/Ti ratio (note Ti is essentially immobile). The soil colour in this case, may become yellowish-orange due to goethite, or red because of hem~.tite. Red shales are typically accumulations of red soils that were transported to the depositional sites (oceans and lakes) by fluvial processes. Goethite/hematite in terrestrial red sandstones typically form by reactions between terrestrial sandstones and shallow, oxygenated groundwater during an early diagenetic stage. Therefore, the loss or retention of Fe in palaeosols, and the presence or absence of red shales/sandstones have been used as good indicators for the oxygen level of an ancient atmosphere. Holland and his associates (e.g., Holland, 1994; Rye and Holland, 1998) have suggested that all palaeosols older than c. 2.0 Ga have lost Fe while younger palaeosols have retained Fe. As Holland (1999, p. 22) states, "the chemistry of palaeosols is the single most important reason for suggesting a dramatic rise in pO2 between c. 2.3 and c. 2.05 Ga" (see also section 5.11 and especially section 5.10 for discussion of palaeosols). However, the loss, as well as the enrichment of Fe is very common in soils of all ages, both pre- and post 2.2 Ga (e.g., Driese et al., 1992; Retallack and German-Heins, 1994, 1995; Ohmoto, 1996b; Beukes et al., 2002). Leaching of Fe from soils is carried out mostly during rainy seasons by organic acids generated from the decay of vegetation on and in soils; organic acids are excellent complexing agents for both ferrous and ferric irons (e.g., Stumm and Morgan, 1996). The Fe dissolved in soil water re-precipitates as ferric (hydr)oxides primarily during dry seasons by reactions with 02 molecules that diffuse through the soil zone. Repeated Fe dissolution during rainy seasons and reprecipitation of ferric (hydr)oxide during dry seasons are the primary processes leading to the formation of laterites (soils highly enriched in goethite) in tropical regions. Beukes et al. (2002) have recognised that laterites and red beds of the same age (c. 2.3 Ga) occur over a very large region in South Africa, suggesting an extensive development of terrestrial biomats and of an oxygenated atmosphere by c. 2.3 Ga. This suggestion is consistent with the discovery by Watanabe et al. (2000) of remnants of microbial mats that developed on and inside soils at c. 2.6 Ga in the Schagen area (Mpumalanga province, South Africa). The oldest red beds known to Cloud (1968) were 2.0-1.8 Ga; thus Cloud suggested the rise of 02 between 2.0 and 1.8 Ga. Since then, the age of the oldest red beds has been extended to between 2.45 and 2.2 Ga (the Jatulian red beds in Finland: Holland, 1994; the Gowganda red beds in Ontario, Canada: Kirkham and Roscoe, 1993) (see also section 5.10). The Pronto and Denison palaeosols in the Elliot Lake district, Ontario, Canada, are probably 2.45 Ga in age (Kirkham and Roscoe, 1993). It has been argued whether these palaeosols formed under a reducing or oxic atmosphere. Because of the losses of Fe from

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some parts of these soil profiles, some (e.g., Kirkham and Roscoe, 1993; Rye and Holland, 1998) infer that these palaeosols formed under an anoxic atmosphere. However, because these palaeosols show trends of increasing Fe3+/Ti with decreasing FeZ+/Ti ratios upwards in the profiles, Ohmoto (1996b) suggests their formation under an oxic atmosphere. The 2.7 Ga red beds at Schebandowan, Ontario, Canada, add another line of evidence for the early rise of atmospheric O2 (e.g., Shegelski, 1980), but a question remains as to whether they are true reds beds or hematitised volcanic rocks. The theory of a methane-rich Archaean atmosphere was advocated by Rye et al. (1995) based on the assumption that siderite (FeCO3) was not stable in surface environments (e.g., soils) prior to c. 2.2 Ga. Using thermodynamic data on the pertinent mineral reactions, they have calculated that the pCO2 value necessary to stabilise greenalite (ferrous silicate mineral), but not siderite, is less than the pCO2 value estimated from a climatic model, where CO2 was the only greenhouse gas (e.g., Kasting, 1987). Rye et al. (1995) have, thus, concluded that an additional greenhouse gas, probably methane, was necessary to maintain the surface temperature above 0~ However, their assumption is not consistent with the well-known fact that siderite and hematite are the two common minerals in chert-jasperbanded iron-formation (BIF) sequences deposited in shallow oceans (e.g., Lake Superiortype BIFs) and in deep oceans (e.g., Algoma-type BIFs) (Kimberley, 1989; Gross, 1991) (cf. section 5.4). The pCO2 and pCH4 values calculated by Rye et al. (1995) and Pavlov et al. (2001a, b) for the Archaean atmosphere need to be re-evaluated.

The behaviour of Fe in marine environments (banded iron-formations, volcanogenic massive sulphide deposits and pillow lavas) Banded iron-formations are sedimentary rock formations with alternating silica-rich layers and iron-rich layers that are typically composed of iron oxides (hematite and magnetite), iron-rich carbonates (siderite and ankerite), and/or iron-rich silicates (e.g., minnesotaite and greenalite). They are typically several metres to several hundred metres in thickness, and extend from a few kilometres to several hundred kilometres (cf. section 5.4). Since Cloud (1968), BIFs have been used as very important evidence for an anoxic atmosphere prior to about 1.8 Ga. Cloud (1968, 1983) suggested that the Fe in BIFs was supplied by weathering of rocks on continents, transported as Fe 2+ in river water to the oceans, where the Fe 2+ was fixed as ferric oxides by the O2 molecules locally generated by cyanobacteria. To transport a sufficient amount of Fe 2+ in Archaean surface water, the atmosphere must have been free of O2. According to Cloud (1968, 1983) the atmosphere remained free of O2 and the oceans FeZ+-rich because the 02 production flux was less than the Fe 2+ flux to the oceans before c. 1.8 Ga ago. The increased 02 production flux c. 1.8 Ga ago caused the atmosphere to become oxic, river and ocean waters to become FeZ+-poor, the BIFs to disappear, and the red beds to appear in the sedimentary record (Cloud, 1968). Holland (1984, 1994) modified Cloud's model, suggesting that the Fe in BIFs originated as Fe z+ in submarine hydrothermal fluids which discharged on mid-ocean ridges and migrated to shallow continental shelf areas by upwelling deep ocean water (section 5.4). The

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379

precipitation of ferric (hydr)oxides (goethite and hematite) was suggested to have occurred by photo-oxidation of Fe 2+ by UV light under an anoxic atmosphere. However, recognising that major B IF formation continued for at least c. 500 My after the "Great Oxidation Event" at c. 2.3 Ga, Holland (1999, p. 20) stated that "The formation of B IFs tells us more about the oxidation of the deeper parts of the oceans than about the atmosphere. The cessation of BIF deposition at c. 1.8 Ga may be a signal that the deep ocean basins became oxygenated, (i.e., mO2 > 0 mol/kg) at that time, and that during the following 1 Ga the hydrothermal flux of iron was oxidised and precipitated close to the vents, as they are today". This statement would imply that the atmospheric pO2 was greater than c. 0.5 PAL to create an oxic deep ocean since c. 1.8 Ga. Compared to the large BIFs like those in the Hamersley basin of Western Australia (c. 2.6-2.4 Ga in age), their time-equivalents in the Transvaal basin, South Africa, and in the Lake Superior region of the U.S.A. and Canada (c. 2.0 Ga), very little attention has been paid to Algoma-type BIFs that formed throughout geologic history, including the c. 3.8 Ga BIFs in Isua (sections 2.2 and 2.3), the c. 2.7 Ga BIFs in the Abitibi Greenstone belt, Canada (section 2.4), the Ordovician BIFs in the Bathurst district, Canada, and the modern Red Sea metalliferous sediments (e.g., Kimberley, 1989; Gross, 1991; Peter, 2001). Algoma-type BIFs formed in submarine volcanic terrains, often together with volcanogenic massive sulphide deposits, at ocean depths > 1 km (Ohmoto, 2002). No difference exists in the mineralogy, ore textures, and geochemistry between B IFs older and younger than 2.3 Ga and between the Hamersley-Superior-type BIFs and Algoma-type BIFs (e.g., Dimroth and Kimberley, 1976; Kimberley, 1989; Gross, 1991) (see also section 5.4, where Trendall and Blockley recommend rejection of the distinction between Algoma and Superior BIF designations). This implies that the main mechanism for precipitation of iron oxides has been the same for all BIFs: mixing of Fe2+-bearing hydrothermal solutions with O2-rich seawater (Ohmoto, 1993, 2002). The suggestion of the oxygenated deep oceans (see, however, section 5.4), and thus the atmospheric pO2 > 0.5 PAL at c. 2.7 Ga is also supported by: (a) the occurrences of 2.7 Ga pillow lavas with oxygenated rims (i.e., increased ferric contents) in the Abitibi district of Canada (Dimroth and Lichtblau, 1978); and (b) by the rare earth element (REE) data on the 2.9 Ga BIFs from India, that display distinct negative Ce anomalies (Kato et al., 2002) (Fig. 5.2-7). Oxygenated, deep ocean water becomes depleted in Ce relative to other REEs because, during deep circulation, Ce 3+ in the ocean water is continuously removed as Ce4+-oxides together with Fe 3+- and Mn4+-hydroxides. BIFs are, therefore, important evidence for, not against, an oxygenated Archaean atmosphere. A major question concerning BIFs remains: why BIFs older than c. 2.0 Ga are generally larger than those younger in age? This is basically the same type of question as asked for the uraniferous quartz-pebble conglomerates (see above). It seems unlikely that the size of these ore deposits was related to atmospheric chemistry (as also found by Trendall and Blockley, section 5.4; they also raise depository issues). The most important factor controlling the size of these deposits was probably the extent of igneous and hydrothermal processes. The extent of igneous and hydrothermal activity is largely controlled by the thermal regime of the continental and oceanic crusts. Since the mantle of the early

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Fig. 5.2-7. Rare earth element chemistry of 2.9-2.7 Ga banded iron-formations in India (Kato et al., 2002). Note the presence of distinct positive Eu anomalies and negative Ce anomalies in most samples, which suggest the formation of the BIFs by the mixing of hydrothermal fluids and an oxygenated bottom ocean water.

Earth was most likely hotter and the geothermal gradients in the crusts higher than today (sections 2.8, 3.2, 3.3, 3.4 and 3.6), it would be reasonable to conclude that the Archaean igneous and hydrothermal systems were generally larger than younger ones (e.g., Barley et al., 1997; Isley, 1995; Condie, 2001a; Ohmoto, 2002). Therefore, the size and age distributions of the uraniferous quartz-pebble conglomerates, gold deposits, banded ironformations, and volcanogenic massive sulphide deposits are similar (Fig. 5.2-8), and were related to the thermal history of the Earth's interior (Fig. 5.2-9). A relationship between ocean-atmosphere composition and bio-geological evolution, on the one hand, and the supercontinent cycle (or plate tectonic paradigm), on the other, is emphasised by Lindsay and Brasier (section 5.3). The evolution of sulphur-utilising microbes and the geochemical cycles of sulphur The modern geochemical cycle of sulphur in the atmosphere-ocean-crust system has been closely linked to atmospheric O2 through: (a) the formation of biogenic pyrite by

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

381

Fig. 5.2-8. Age distributions of Rapitan/Clinton iron-formations, Algoma and Superior type banded iron-formations, volcanogenic massive sulphide deposits, vein-type Au-U deposits, and quartz-pebble conglomerate-type Au-U deposits. Data from Meyer (1985), Kimberley (1989), Barley and Groves (1992), and Ohmoto (1996a).

Fig. 5.2-9. Schematic illustration of the inferred geological environments for genesis of banded ironformations and volcanogenic massive sulphide deposits.

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

utilising oceanic SO 2- and remnants of organisms from the overlying photic zone (reactions 11 and 12); and (b) the oxidation of pyrite during weathering on land (reaction 13): 2R-CH20 + 2H + + SO 2- --+ 2CO2 + 2H20 + H2S + 2R,

(11)

(organics)

where R represents non-metabolisable organic matter that remains in sediments as residual organic matter. The H2S generated from sulphate reduction may react with ferric oxides (goethite) in water columns and in sediments to form diagenetic pyrite through reactions such as: Fe(OH) 3 + 2H2S --+ FeS2 + 3H20 + 89 (goethite)

(12)

(pyrite)

Because of reactions (11 ) and (12), the amounts of pyrite and organic carbon (R) typically show positive correlations (e.g., Ohmoto and Goldhaber, 1997). Large fractionation of sulphur isotopes occurs during reaction (11). The ASO4-H2S (= 634Ss04 - 634SH2s) value ranges from c. 5%0 to c. 60%0, depending on the sulphate content of water and the rate of sulphate reduction, which in turn depends on the substrate and temperature (e.g., Ohmoto and Felder, 1987). Biogenic pyrite crystals are, therefore, characterised by generally negative and variable 334S values. The formation of biogenic pyrite results in a shift of the 634S value of sea water SO 2- towards a positive value (see also Lyons et al., section 5.5, for detailed discussion of the sulphur isotope record and its relation to atmospheric oxygen). Certain quantities of SO 2- in the oceans, perhaps as much as one-tenth of the present concentration of 28 mM, could have been formed by photochemical reactions of volcanic SO2 in an anoxic atmosphere (e.g., Walker and Brimblecombe, 1985). However, an oceanic content of SO 2- greater than c. 3 mM probably requires the production of SO42- by the oxidation of biogenic pyrite, using atmospheric 02: 4FeS2 + 1502 + 14H20 --+ 4Fe(OH)3 + 16H + + 8SO24-.

(13)

(pyrite)

Higher concentrations of SO 2- in the oceans may result in the increased formation of pyrite with more negative 634S values. Therefore, researchers have used the contents and ~34S values of pyrite and sulphates, the Aso4-reS2 values, and also the S/C/Fe ratios, in sedimentary rocks, to estimate whether sulphate reducing bacteria (SRB) were active or not and how the sulphate content in the oceans changed through geologic time (e.g., Ohmoto, 1992; Strauss, 2002) (section 5.5). Earlier researchers (e.g., Hattori et al., 1983; Lambert and Donnelly, 1992) recognised that the 334S values of both sulphides and sulphates of > 2.2 Ga age mostly fell within a range of 0 4- 5%0, but deviated significantly from 0%0 after c. 2.2 Ga. This suggests the emergence of SRB and an increase of seawater SO 2- content at c. 2.2 Ga (suggested also by Lyons et al., section 5.5). However, more detailed investigations have revealed a much

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

383

larger variation in pre-2.2 Ga sediments, including those c. 3.4 Ga in age (Ohmoto and Felder, 1987; Ohmoto et al., 1993; Shen et al., 2001) and those of c. 2.7 Ga (Grassineau et al., 2001) (Fig. 5.2-10). Ohmoto et al. (1993) have suggested that by 3.5 Ga, SRB evolved and the oceans became sulphate-rich. Although Shen et al. (2001) agree on the early emergence of SRB, Habicht et al. (2002) propose that the oceanic SO ] - content remained less than 1/100 of the present ocean until c. 2.2 Ga, except in local evaporite basins (Lyons et al., section 5.5). Canfield and Raiswell (1999), Bjerrum and Canfield (2002) and Habicht et al. (2002) further suggest that the increased SO ] - content since c. 2.2 Ga allowed for a higher production of biogenic H2S and changed the oceans to H2S-rich (except in the photic zone); this caused the removal of Fe 2+ as sulphides, rather than as iron oxides, resuiting in the cessation of BIF formation (see Fig. 5.2-1b). To support this model, they argue that the sulphide contents of Archaean shales are much lower than those of younger shales. However, many investigators (e.g., Dimroth and Kimberley, 1976; Clemmey and Badham, 1982; Holland, 1984; Yamaguchi, 2002) have recognised that the sulphide contents of Archaean shales are essentially the same as those of younger shales. The lack of a trend in the sulphide content of shales with geologic time, and the presence of positive correlations between organic carbon and sulphide contents in shales of all ages, support the Dimroth-Ohmoto model of atmospheric evolution that postulates essentially the same SO ] - content of the oceans through geologic time (Fig. 5.2-2b). Based mostly on new data from molecular clocks, Canfield and Teske (1996) suggested that the emergence of sulphide-oxidising and sulphur-disproportionating bacteria around 700 Ma caused a significant increase in ASO4_H2 S values, to form pyrites with very negative ~345 values, and to further increase the ~345 value of SO ] - in seawater. They attribute the appearance of these aerobic organisms to the last major rise of atmospheric 02. Logan et al. (2001) recently found biomarkers of these sulphur-utilising bacteria in marine sedimentary rocks c. 1.9 Ga in age, which contradicts Canfield's ocean evolution model. The magnitudes of fractionation among different isotopes of the same element that occur during a normal (bio-)chemical reaction are generally proportional to the differences in the isotope mass. Therefore, among most geologic samples, variations in the 335/32S ratio have been found to be about one-half of those in the 345/325 ratio, resulting in the following relationship, known as the "terrestrial fractionation line (TFL)" or "mass dependent fractionation": $33S = 0.514S34S.

(14)

An important discovery was made by Farquhar et al. (2000a), indicating that some samples of sulphides and sulphates from > 2.0 Ga sedimentary rocks exhibited ~335i - $345i relationships that did not plot on the TFL; they have termed such isotopic relationships as "mass independent fractionation (MIF)" (Fig. 5.2-11 a). The magnitude of MIF is characterised by the 33A value: 33 A = ~33S - 0.514634S.

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

Fig. 5.2-10. Sulphur isotope records of sulphides (mostly pyrite) and sulphates (gypsum/anhydrite and barite) through geological time, based on compilations by Canfield and Raiswell (1999) and Yamaguchi (2002). Also added are data from 2.7 Ga sedimentary rocks by Grassineau et al. (2001).

Farquhar et al. (2000a, 2001) advocate that the MIF of sulphur isotopes occurs only through atmospheric photochemical reactions by UV light involving SO2 gas (Fig. 5.2-11 b). Since UV photochemical reactions are inhibited by the presence of an ozone shield and an ozone shield may form when the atmospheric pO2 is greater than c. 0.1% PAL, many researchers (Farquhar et al., 2001; Kasting, 2001; Pavlov and Kasting, 2002) interpreted the existence of MIF in sediments prior to c. 2.0 Ga as the best evidence for the " C - W - H - K " model. However, the isotopic relationships found in natural samples, especially in sulphides, greatly differ from those found in laboratory experiments: natural sulphides with MIF mostly have positive values of 33A, $335 and $345, yet no sulphur-bearing compounds produced by laboratory photochemical reactions possess the same isotopic characteristics (see Fig. 5.2-1 l a). Furthermore, it has not yet conclusively been demonstrated that MIF of S isotopes is absent in rocks younger than 2.3 Ga. Further work is required to establish the connections between sulphur MIF in rocks and atmospheric evolution.

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

385

Fig. 5.2-11. (a) Comparison of the ~34S and ~33S values of pyrite crystals in Archaean sedimentary rocks (Farquhar et al., 2000a; Rumble et al., 2002) and the S0, SO 2- and SO2 produced by UV radiation of SO2 (Farquhar et al., 2001). (b) Schematic illustration of the formational mechanisms of pyrite and barite in the Archaean oceans as proposed by Farquhar et al. (2000a).

The Geochemical Cycles of Carbon and Oxygen and the Mechanisms Controlling the Atmospheric 02 Level The oxygen geochemical cycle is linked to the carbon geochemical cycle, largely through reaction (1). For the carbon geochemical cycle, the other important processes include the

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

precipitation and dissolution of carbonates (Ca 2+ 4- 2HCO 3 = CaCO3 4- H20 4- CO2), volcanic inputs of CO2 and CH4, and the subduction of organic matter and carbonates to the mantle (Lasaga and Ohmoto, 2002). If we assume that the 313C value of the atmosphere-ocean-crust system has remained constant at -5%0, then the nearly uniform 313C values of about 0%c for marine carbonate and about -25%0 for the organic C in Phanerozoic marine sediments (see Fig. 5.2-3) suggest that the burial flux of carbonate carbon (Fb,carb) has been about four times the burial flux of organic carbon (Fb,org), i.e., about 4 • 1013 moles yr -1, during the Phanerozoic period (Lasaga and Ohmoto, 2002). The mass balance equation used for this estimate is:

r 13C~ C n r 13 Corg(Fb, org / (Fb, org 4- Fb , carb)) + S 13Ccarb(Fb,carb/(Fb,org + Fb,carb)).

(16)

Karhu and Holland (1996) recognised a large positive excursion of 613Ccarb values (up to c. +10%0) in carbonates 2.2-1.9 Ga in age (see Fig. 5.2-3a) (see also section 5.3). Applying the above mass balance equation (16) to this ~13Ccarb excursion, they suggested a significant increase in the burial flux of organic carbon (= increased flux of 02 production) for this period, as a cause of the "Great Oxidation Event". However, a serious problem with this interpretation concerns the age relationship: according to the Holland-Kasting model (Fig. 5.2-1a), the "Great Oxidation Event" occurred more than c. 100 My before the deposition of these carbonates. Therefore, the increased burial flux of organic carbon during the 2.2-1.9 Ga period, if it was real, was an unlikely cause for their proposed rise of 02 at c. 2.3 Ga. Another problem in applying the carbon isotopic mass balance approach in constraining the burial flux of organic carbon (= the 02 production flux) is that this approach may provide information on the ratio of burial flux of organic carbon (Fb,org) to burial flux of carbonate carbon (Fb,carb), but not the Fb,org value itself. The burial fluxes of organic carbon must be constrained from contents of organic carbon in sedimentary rocks. As recognised by Dimroth and Kimberley (1976), Holland (1984), and Watanabe et al. (1997), there is no fundamental difference in the organic carbon contents of shales through geologic time. The organic carbon contents of Archaean shales range from < 0.1 wt.% to as high as c. 15 wt.%, with an average of c. 0.6 wt.%; these values are essentially identical to those of Phanerozoic shales. Such data suggest that the 02 production flux has been basically the same, within an order of magnitude, through geological time. Lasaga and Ohmoto (2002) have shown that the burial flux of organic carbon (Fb,org) is largely controlled by the primary productivity (PP) and the burial efficiency (se). The PP is in turn controlled by the nutrient flux, the ocean circulations, and other parameters. The value depends primarily on the sedimentation rate of clastic sediments and the 02 content of deep oceans; the latter is linked to the atmospheric pO2. Lasaga and Ohmoto (2002) further suggest that the 02 production flux would increase by about 7 times if the atmospheric pO2 drops from 1 PAL to less than c. 0.5 PAL, if the phosphate flux to the oceans and the average sedimentation rate remained the same (Fig. 5.2-12). This dependence of the burial flux of organic carbon on pO2 offers a major negative-feedback mechanism controlling

5.2. Archaean Atmosphere, Hydrosphere and Biosphere

387

Fig. 5.2-12. The relationship between the global O2-production (Fprod,O2) and O2-consumption fluxes (Fv,O2 q- Fs,O2) as a function of the atmospheric pO2 level (Lasaga and Ohmoto, 2002). the atmospheric pO2 level. For example, a drop in the atmospheric pO2 would increase the burial flux of organic carbon and the 02 production flux to restore the atmospheric pO2 level. Lasaga and Ohmoto (2002) have also computed the O2 consumption flux by soil oxidation (Fs,o2) as a function of the important parameters (e.g., pO2, soil thickness, soil depth, land area, organic C content). An important conclusion from their study is that the Fs,O2 value decreases with decreasing pO2 (Fig. 5.2-12), offering another major negativefeedback mechanism for controlling the atmospheric pO2 level. According to Holland (1994), Rye and Holland (1998), Kasting and Brown (1998), and Kump et al. (2001), an anoxic atmosphere was maintained prior to c. 2.3 Ga because the volcanic flux of reducing gas (Fv,o2) was about 3 times greater than today (i.e., > 0.75 x 1013 moles yr-l), resulting in the total 02 consumption flux (Fcons,O2 = Fs,o2 + Fv,02) being greater than the O2 production flux (Fprod,O2). However, they did not take into consideration the dependencies of Fprod,O2 and Fs,02 values on pO2 (see Fig. 5.2-12). When these dependencies are taken into account, we can recognise that, in order for Fcons,02 > Fprod,O2 to occur, the Fv,o2 value must be greater than c. 7 x 1013 moles yr -I (i.e., more than 7 times the present O2 production flux), which is more than c. 30 times the present volcanic flux of reducing gas. Therefore, it was very unlikely that the atmosphere remained anoxic once oxygenic photosynthetic organisms became the dominant primary producers and the burial flux of organic carbon reached values similar to those at present. Lasaga and Ohmoto (2002) suggest the

Chapter 5: Evolution of the Hydrosphere and Atmosphere

388

atmospheric pO2 would have been maintained between 0.5 and 2 PAL by the coupling of the two negative-feedback mechanisms (Fig. 5.2-12) since the appearance of oxygenic photosynthetic organisms at more than 2.8 Ga, possibly > 3.7 Ga ago (Rosing, 1999). Conclusions

There are at least two contrasting models for the evolution of atmospheric oxygen: the C - W - H - K (Cloud-Walker-Holland-Kasting) model that postulates the "Great Oxidation Event" at c. 2.3 Ga, and the D-O (Dimroth-Ohmoto) model that postulates essentially a constant pO2 since c. 4 Ga. The C - W - H - K is certainly the more popular, but not necessarily a better, model. Many uncertainties and problems still exist in both models. In addition to those data presented and discussed above, there are other types of geochemical data that have been used to support one or the other model, including: (a) the abundance of sulphate minerals (anhydrite, barite) in Archaean sedimentary sequences; (b) the concentration ratios of some redox-sensitive elements (Mo, U, V, etc.), organic carbon, and sulphide-sulphur in shales; (c) the nitrogen isotope ratios (615N values) of organic matter; and (d) the rare earth element ratios (especially Ce anomalies) in palaeosols. But the published data on these topics are meagre and often too ambiguous. An important message that I wish to convey here is that we need to conduct a lot more geochemical investigations, especially on a variety of rocks older than c. 2.7 Ga, in order to understand better how the environment and life evolved together on the early Earth.

5.3.

THE EVOLUTION OF THE PRECAMBRIAN ATMOSPHERE: CARBON ISOTOPIC EVIDENCE FROM THE AUSTRALIAN CONTINENT

J.F. LINDSAY AND M.D. BRASIER Introduction

The analysis of biogeochemical signatures preserved in the sedimentary record provides one of the most promising means of following the evolution of Earth's early atmosphere and biosphere (Knoll and Canfield, 1998). The analysis of the stable isotopes of carbon (lec and 13C---expressed as 613C in %0 PDB), which are fractionated during autotrophic fixation of COa, provide useful insights into the carbon cycle, the growth of the crustal carbon reservoir and the nature of the atmosphere (see also discussions in section 5.2). The record of these events is well documented on the Australian continent where a succession of predominantly intracratonic Precambrian basins has been preserved. Here we attempt to outline the extensive Australian Precambrian carbon isotopic record preserved in platform carbonate rocks and place it in a global tectonic framework. ]'he PrecambrianEarth: Temposand Events l~ited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

389

5.3. Evolution of the Precambrian Atmosphere

Evolution o f the Australian Craton

The Australian isotopic record is to be found in a series of well preserved largely intracratonic basins that rest upon the ancient craton. The "Australian craton" (used here as reflecting its present-day nature) is complex and consists of a mosaic of crustal fragments, or "mega-elements" (a term defined by Shaw et al., 1996), with a broad range in age and degree of deformation. The continent consists of eight crustal mega-elements (Shaw et al., 1996) four of which underlie the intracratonic basins; the Southern Australian, Western Australian, Central Australian and Northern Australian mega-elements (Fig. 5.3-1). The development of these mega-elements was initiated early in the Archaean but aggregation occurred largely in the Palaeoproterozoic. The final amalgamation of the mega-elements to form the present Australian craton extended into the Mesoproterozoic. The earliest evidence of crustal formation in Australia comes from zircons preserved in sandstones on the Pilbara block in Western Australia (Froude et al., 1983; Compston and Pigeon, 1986; Kober et al., 1989). However, crust appears to have developed at first in the form of microcontinents, which did not assemble into continents until c. 3.0 Ga (Condie, 1998). The first clear evidence for crustal processes in Australia is found at the core of

~p

0 ,

1000 km

o

,

Fig. 5.3-1. Crustal mega-elements of the Australian craton (SA, CA, WA and NA are the Southern, Central, Western and Northern Australian mega-elements, respectively) (after Shaw et al., 1996). Cross-hatched areas indicate younger Palaeozoic successions.

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the Pilbara block where a well preserved stratigraphic succession records the initiation of the protocontinent beginning at some time before c. 3.5 Ga (Buick et al., 1995a; Van Kranendonk, 2000; Lindsay, 2002) (see also section 3.6). Rogers (1993, 1996) has argued that, since only five of the Earth's cratons contain laterally extensive, shallow-water, supracrustal suites (i.e., evidence of intracratonic basinal successions) as old as c. 3.0 Ga, and since they all occur in east Gondwana, they must have formed on an original continent. Cheney (1996) also observed similarities between the Pilbara in Western Australia and the Kaapvaal craton of southern Africa and concluded that they formed part of a larger continent. While the details of these reconstructions differ, there is little doubt that by c. 3.0 Ga modern-style stable cratons had come into existence and begun to accumulate laterally extensive intracratonic successions (cf. Rogers, 1996) (see also sections 3.2, 3.6 and 3.9). The Pilbara block was thus well established as a stable continental nucleus when the Hamersley basin began to subside and accumulate sediments at c. 2.8 Ga, following an episode of crustal extension (Fig. 5.3-2) (Trendall, 1983b, 1990; Blake and Barley, 1992). The basin-fill is complex and polyphase with major erosional surfaces separating areally extensive megasequences (Trendall, 1990; Krapez, 1996), much as seen in the Palaeoproterozoic and Neoproterozoic basins of central and northern Australia (Lindsay and Korsch, 1989; Lindsay and Leven, 1996; Lindsay and Brasier, 2000). The Hamersley succession is largely marine and includes major intervals of banded iron-formation (B IF) (Trendall and Blockley, 1970) (section 5.4) and some of the Earth's earliest platform carbonates (c. 2.7-2.5 Ga: Simonson et al., 1993a, b). The Hamersley basin is thus the first clearly identifiable basinal setting preserved on the Australian craton in which marine sediments, and in particular, platform carbonate rocks have been preserved in response to broad, regional, crustal subsidence. The sedimentary succession also records the earliest evidence of eustasy in the form of large scale upwardsshallowing depositional sequences. The basin was formed on a major crustal block which may be regarded as the earliest continent or supercontinent which then broke up in the Late Archaean to begin the supercontinent cycle. There then appears to be a long hiatus during which this earliest supercontinent dispersed. The Western Australian mega-element (Fig. 5.3-1) consists of two well-defined Archaean blocks, the Yilgarn and Pilbara cratons, which were sutured along the Capricorn orogen at the same time as the Northern Australian mega-element was evolving. Ocean closure was underway by c. 2.3 Ga and the Pilbara and Yilgarn cratonic margins became active, with ocean floor possibly being subducted beneath the Yilgarn Craton, leading to suturing of the two cratons between 2.0 and 1.8 Ga (Tyler and Thorne, 1990; Thorne and Seymour, 1991; Pirajno et al., 1998; Occhipinti et al., 1998). A series of basins (Ashburton, 2.2 to 1.8 Ga; Yerrida, c. 2.2-1.9 Ga; Bryah, c. 2.0 Ga; Padbury, c. 2.0 Ga; Earaheedy, c.1.9-1.65 Ga) formed along the cratonic suture, recording the convergence and collision of the two cratons. Because of their active margin settings, the fill of these small basins is dominated by clastic sediments. However, platform carbonate units are preserved in the Ashburton and Yerrida basins and a significant thickness of carbonate rocks also occurs in

5.3. Evolution of the Precambrian Atmosphere

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the Earaheedy basin (Tyler and Thorne, 1990; Thorne and Seymour, 1991; Gee and Grey, 1993; Occhipinti et al., 1997, 1998; Pirajno et al., 1998; Pirajno and Adamides, 2000), allowing for isotopic coverage of this critical time interval.

Fig. 5.3-2. Distribution of Australian Precambrian intracratonic basins through time. The Hamersley basin is the earliest intracratonic depository in Australia. Major periods of basin formation occur in the Palaeoproterozoic and Neoproterozoic, associated with supercontinent formation (see Lindsay et al., 1987; Lindsay and Brasier, 2000).

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There is a considerable body of evidence to suggest that by 2.0 Ga a supercontinent similar in significance to Pangaea had come into existence (Hoffman, 1988, 1989a, b, 1991). Regional data thus suggest that the supercontinent began to assemble at some time close to 2.0 Ga and then began to disperse again at approximately 1.8 Ga, probably as a result of mantle instability (cf. Gurnis, 1988) (see also sections 3.2, 3.3 and 3.9). Recently, Rogers and Santosh (2002) have attempted to define the overall structure of this Palaeoproterozoic to Mesoproterozoic supercontinent, which they refer to as Columbia (see also section 3.11). The northern and western Australian mega-elements evolved in parallel during the Barramundi and Capricorn orogenies, respectively, as part of the assembly of Columbia. The Barramundi orogeny, a significant event across much of northern Australia (Page and Williams, 1988; Plumb et al., 1990; Needham and De Ross, 1990; Le Messurier et al., 1990; O'Dea et al., 1997), appears to be associated with the final phase of the assembly of the supercontinent. Crustal shortening, voluminous igneous activity (Wyborn, 1988) and low pressure metamorphism (Etheridge et al., 1987) during this event produced the basement rocks underlying much of northern Australia (Plumb et al., 1980). The crust, which probably evolved on earlier Archaean continental crust, is thick (43-53 km; Collins, 1983), and may well have been thicker in the past. Beginning at approximately 1.8 Ga, large areas of the Northern Australian megaelement began to subside, possibly as a response to mantle instability and the intrusion of anorogenic granites at the time of the breakup of Columbia (Rogers and Santosh, 2002; cf. Gurnis, 1988; Wyborn, 1988; Idnurm and Giddings, 1988a; Pysklywec and Mitrovica, 1998). Upwelling of the plume and granite intrusion initially caused domal uplifting of the continental lithosphere and regional peneplanation which in turn led to thermal relaxation and widespread subsidence following the output of flood basalts and the cessation of plume activity (Lindsay, 1999, 2002) (see also sections 3.2 and 3.3). Subsidence in response to these events led to the development of a series of Palaeoproterozoic to Mesoproterozoic intracratonic basins including the McArthur, Mount Isa, Victoria and Kimberley basins (possibly also the Birrindudu basin) (Fig. 5.3-2), which blanket the Northern Australian mega-element (Lindsay, 1998,2001). The Bangemall basin developed on the newly formed Western Australian mega-element at much the same time as the Northern Australian basins developed on the Northern Australian mega-element (Muhling and Brakel, 1985). These basins are complex polyphase structures which continued to subside for more than 200 My, preserving in excess of 10 km of sediment, all with similar basin-fill architectures and including significant thicknesses of platform carbonates (Lindsay and Brasier, 2000). The basins have experienced only mild and often localised tectonic activity since 1.8 Ga (Plumb et al., 1990), and are thus well preserved and provide an ideal setting for a detailed study of Palaeo- to Mesoproterozoic isotopic signatures (Lindsay and Brasier, 2000). The Central and Southern Australian mega-elements were amalgamated somewhat later than their northern and western counterparts. However, the process was probably completed in the Late Mesoproterozoic, by approximately 1.1 Ga (Myers et al., 1994, 1996; Camacho and Fanning, 1995; Clarke et al., 1995). The final amalgamation occurred as part of the aggregation of the Rodinian (sections 3.10 and 3.11) supercontinent. The crust de-

5.3. Evolution of the Precambrian Atmosphere

393

veloped during this period was pervaded by north-dipping, thick-skinned thrust faults; it was thick (40-50 km: Lindsay and Leven, 1996) and strong and able to support stress over long geological time periods (Haddad et al., 2001). Thus, by 1.1 Ga, the crustal substrate was in place, setting the stage for the evolution of the central Australian Neoproterozoic basins. Following a period of stability (1.1-0.8 Ga) a large area of central Australia, in excess of 2.5 • 106 km 2, began to subside in synchroneity (Fig. 5.3-2). This major event was due to mantle instability resulting from the insulating effect (e.g., sections 3.2 and 3.11) of Rodinia. Initially, beginning at c. 900 Ma, a rising superplume (sections 3.2 and 3.3) uplifted much of central Australia (Zhao et al., 1994; Lindsay, 1999) leading to peneplanation of the uplifted region and the generation of large volumes of sand-sized clastic materials. Ultimately, the decline of the superplume led to thermal recovery and the development of a sag basin (beginning at c. 800 Ma), which in turn resulted in the redistribution of the clastic sediments and the development of a vast sand sheet at the base of the Neoproterozoic succession (Lindsay, 1999). The superbasin generated by the thermal recovery (Fig. 5.3-2) was short lived (c. 20 My) but, in conjunction with the crustal fabric developed during supercontinent assembly, it set the stage for further long term basin development that extended for half a billion years, well into the Late Palaeozoic (Lindsay, 2002). Following the sag phase at least five major tectonic episodes influenced the central Australian region. Compressional tectonics reactivated earlier thrust faults that had remained dormant within the crust, disrupting the superbasin, causing uplift of basement blocks and breaking the superbasin into the four basins now identified within the central Australian Neoproterozoic succession (Officer, Amadeus, Ngalia and Georgina basins). These subsequent tectonic events produced a distinctive foreland basin architecture and were perhaps the trigger for the Neoproterozoic ice ages (Lindsay, 1.989) (see also sections 5.6-5.8). The reactivated basins became asymmetric with major thrust faults along one margin, parallelled by deep narrow troughs that formed the main depocentres for the remaining life of the basins (c. 290 My). The central Australian basins are the product of events surrounding the assembly and dispersal of Rodinia (Lindsay, 2002). The development and later dispersal of Rodinia (McMenamin and McMenamin, 1990), beginning at around 1 Ga, has been broadly outlined elsewhere (e.g., Bond et al., 1984; Dalziel, 1991, 1992; Li et al., 1996) (sections 3.1.0 and 3.11). More specifically the connection between the Australian craton and Rodinia is discussed in Lindsay et al. (1987), Powell et al. (1994) and Lindsay (2002).

The Carbon Isotopic Record A comprehensive sedimentary record is preserved in the three temporal groupings of sedimentary basins that overlie the Australian craton (Fig. 5.3-2). The basins are largely intracratonic in nature and contain a significant proportion of platform carbonates, often occurring at the top of upwards-shallowing eustatically generated depositional sequences (Lindsay and Korsch, 1989; Lindsay and Leven, 1996; Lindsay, 2002). More than 2000 analyses of the isotopic composition of these carbonates, along with associated major and

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trace element studies, are now available to outline broadly the secular stable carbon isotope curve for much of the Australian Precambrian (e.g., Calver and Lindsay, 1998; Brasier and Lindsay, 1998; Lindsay and Brasier, 2000, 2002). Analytical techniques developed for these early samples are discussed in Lindsay and Brasier (2000, 2002). Sampling was preferentially undertaken at 5-10 m intervals on carbonates from the less deformed parts of the basins and on lithologic intervals free from evidence of secondary alteration. Selected portions of carbonate were cleaned and analysed using a VG Isomass PRISM mass spectrometer attached to an on-line VG Isocarb preparation system in the Oxford University laboratories (cf. Brasier et al., 1996; Lindsay and Brasier, 2000). Major and trace element analyses were carried out using XRF and ICP-MS. The extreme age of the samples required further careful analysis of possible diagenetic alteration even where thin section evaluation suggests diagenetic effects are minimal. Covariance in ~ 13Ccarb/~ 18Ocarb cross-plots were used as an indicator of diagenetic alteration. Overall, we found that the primary fabrics of these ancient carbonates are well preserved, especially in the major platform carbonate units. We take this to suggest that early diagenesis, including dolomitisation and silicification, was predominant (cf. Veizer et al., 1990, 1992; Buick et al., 1995a; Lindsay and Brasier, 2000) (see also section 6.4). Thin section analysis suggests that the carbonate rocks were largely sealed against the passage of fluids during later diagenesis, thereby preserving their fabric and retaining the primary 13Ccarb signatures.

The Archaean and early Palaeoproterozoic basins The late Archaean and early Palaeoproterozoic basins that rest upon the west Australian mega-element (Fig. 5.3-2) form a time-series associated with the formation and ultimate disassembly of one of the Earth's earliest major continental masses (Tyler and Thorne, 1990; Thorne and Seymour, 1991; Gee and Grey, 1993; Occhipinti et al., 1997, 1998; Pirajno et al., 1998; Pirajno and Adamides, 2000; Lindsay and Brasier, 2002). These basins contain an important and ancient sedimentary record of the early Earth including some of the earliest carbonate platform deposits (Simonson et al., 1993a, b). Most of the carbonates, especially in the Hamersley basin, form the highstand systems tracts of major upwards shallowing depositional sequences. The sequences begin with black shale and grade upwards to BIF. The sharp transition to carbonates above the BIF suggests that the ocean was stratified such that surface waters were oxygen-rich but depleted in iron thus allowing carbonate deposition (see also sections 5.2 and 5.4). At greater depths the water column was oxygen deficient as iron was rapidly precipitated to form massive deposits of BIE In all, a total of 474 carbonate samples were analysed from these early basins (Lindsay and Brasier, 2002) (Fig. 5.3-3). In the latest Archaean (c. 2.6 Ga) the secular 613Ccarb curve (Fig. 5.3-3) is flat, much like that seen in the later Palaeoproterozoic basins of northern Australia (< 1.8 Ga). However, in the early Palaeoproterozoic, beginning after 2.5 Ga and continuing until at least 1.9 Ga, the 613Ccarb curve is much more dynamic, with significant positive and negative excursions, including a major positive excursion (+9%0 PDB) close to 2.2 Ga (see also section 5.2). These excursions can be correlated with the Lomagundi event identified in

5.3. Evolution of the Precambrian Atmosphere

395

Africa, Europe and North America (see Karhu and Holland, 1996; Bau et al., 1999). Previously published studies of the overlying Meso- to Palaeoproterozoic Bangemall basin and of 1.8-1.5 Ga old basins in northern Australia suggest that the 613Ccarb curve became relatively monotonic again after c. 1.8 Ga and, as discussed below, remained so for most of the following Mesoproterozoic (Buick et al., 1995a; Brasier and Lindsay, 1998; Lindsay and Brasier, 2000).

The Palaeoproterozoic-Mesoproterozoic basins Shallow marine Palaeo- and Mesoproterozoic sedimentary successions, including widespread platform carbonate intervals, are widely distributed in several major basins across northern Australia (Fig. 5.3-2). The successions are only gently deformed and their stratigraphy is relatively continuous, thus offering an ideal opportunity to document secular variations in carbon isotopes. Marine carbonate intervals from two of these major basins, the McArthur and Mount lsa basins, have been sampled to document secular variation in S13Ccarb from approximately 1700-1575 Ma (Brasier and Lindsay, 1998; Lindsay and Brasier, 2000). The 576 samples were tied to a well dated sequence stratigraphy (see chapter 8). Isotope curves for the two basins (Fig. 5.3-4) show that S13Ccarb values lie within a relatively small range throughout most of the time interval (Lindsay and Brasier, 2000). Values from both basins average -0.6%o and seldom lie more than 1%0 either side of the mean. The mean ~13Ccarb values for the two basins are statistically the same at the 95% confidence level, with means of-0.59%c and -0.65%o in the McArthur and Mount Isa basins, respectively. The data show that the curve is essentially flat, indicating that following the Lomagundi event and prior to the Neoproterozoic isotopic excursions (see also section 5.8), the 313Ccarb record entered a long term biogeochemical stasis. The Neoproterozoic basins Neoproterozoic sedimentary rocks are widely distributed across central Australia (Fig. 5.3-2). Locally these intracratonic basins extend to depths of 15 km and include rocks as young as Carboniferous (Lindsay, 2002). The Neoproterozoic is well represented in the section and includes a significant thickness of platform carbonate and evidence of both the Sturtian and Marinoan glaciation (Lindsay, 1989) (see also section 5.8). Published carbonate carbon isotope data are available for both the Amadeus and Officer basins (Walter et al., 1995; Calver and Lindsay, 1998; Hill and Walter, 2000). Here we present unpublished data from the Amadeus basin (Fig. 5.3-5) which is central to the area and contains one of the thickest and best preserved Neoproterozoic successions (Lindsay, 1987, 1989, 1993). The data show two major positive excursions extending to +5 and +6~ in the preglacial part of the section with a negative excursion of -5%0 lying between them (Fig. 5.3-5). Erosion during the Sturtian glaciation has removed a significant thickness of section (Bitter Springs Formation), perhaps as much as 200 m (Lindsay, 1989), before the curve again becomes negative in the post-glacial cap carbonates (see also sections 5.6, 5.7 and 5.8) (Areyonga Formation) reaching minimum values of -4%o. Following a largely clastic succession, the final carbonates (Julie Limestone) directly beneath the Precambrian-

396

Chapter 5: Evolution of the Hydrosphere and Atmosphere

~ ,.,~

5.3. Evolution of the Precambrian Atmosphere

Fig. 5.3-3. Composite secular carbon isotope curve for the time interval 2.6-1.8 Ga based on data from Western Australian basins (Lindsay and Brasier, 2002). Note the abrupt increase in the S13Ccarb values at c. 2.3 Ga as the Pilbara and Yilgarn blocks converged, marking the transition from a passive to an active margin setting and the subduction of intervening ocean floor sediments beneath the Yilgarn craton. Tectonic outline based on Tyler and Thorne (l990), Occhipinti et al. (1998) and Pirajno et al. (1998). Black diamonds (2.47-2.45 Ga) indicate the abrupt period of mantle overturn proposed by Kump et al. (2001).

Fig. 5.3-4. Composite carbon isotope (6'3Ccart,) curve for the McArthur and Mount Isa basins in northern Australia. Note the small amplitude of deviations compared to Figure 5.3-3. Shaded areas indicate periods of non-deposition (see Brasier and Lindsay, 2000).

397

398

Chapter 5: Evolution of the Hydrosphere and Atmosphere

-6

-4

-2

a ~3Cpoe (%0) 0 2 4

6

8

10

600

LLI I_ "--~00 :3._J

MARINOAN GLACIATION < CO

Z

O~ ~.u_ ILl n,"

L

STURTIAN GLACIATION

t~

uJ I-

z

0 I--< n,," 0 LL O3 (.9 Z n" 13.. O0 n" ILl I-I-rn

700

Fig. 5.3-5. Composite carbon isotope (S13Ccarb) curve for the Amadeus basin in central Australia. Data for the Bitter Springs and Areyonga Formations come from drill core recovered from the Wallera #1 well (24~ 132~ Julie Limestone data derived from samples collected at Ross River (23~ 134029 ' 15.89"E) (C. Calver, 2002, pers. comm.). Shaded area indicates a period of erosion and non-deposition.

5.3. Evolution of the Precambrian Atmosphere

399

Cambrian boundary produce an oscillatory curve peaking at approximately +5%o. Frimmel (section 5.8) presents carbon isotope data for southern Africa. Discussion

There is a growing body of evidence to suggest that there is a periodic cycle of supercontinent coalescence and dispersal (Worsley et al., 1984; Murphy and Nance, 1992; Duncan and Turcotte, 1994; Veevers et al., 1997), driven by large scale mantle convection (Anderson, 1982; Gurnis, 1988; Kominz and Bond, 1991; Tackley, 2000) (sections 3.2, 3.3 and 3.9). The development and dispersal of a Neoproterozoic supercontinent, Rodinia (McMenamin and McMenamin, 1990), beginning at around 1 Ga, whilst not as well documented as Pangaea (cf. Veevers, 1988, 1989), has been broadly outlined (e.g., Bond et al., 1984; Lindsay et al., 1987; Dalziel, 1991, 1992; Li et al., 1996) (sections 3.10 and 3.11). Similarly, Rogers and Santosh (2002) have recently reconstructed a supercontinent now called Columbia which assembled in the Palaeoproterozoic. As discussed previously there is also evidence for the development of a significant continental mass, perhaps a supercontinent, at some time just prior to 2.8 Ga. The supercontinents assembled over geoid lows, mantle downwellings, and dispersed over the geoid highs at mantle upwellings. This cycle is likely to be continuous because the same forces that fragment the old supercontinent over the geoid high are effectively assembling the next supercontinent over the associated low (Condie, 1998) (sections 3.2, 3.9 and 3.11). Evidence therefore suggests three supercontinent cycles from the Archaean to the Neoproterozoic, with crustal assembly at approximately 2.8, 2.0 and 1.0 Ga. In each case, continental assembly appears to have been associated with mantle instability resulting in either mantle overturn and the development of superplumes or partial melting of the crust and upper mantle (sections 3.2-3.4). This in turn resulted in the development of a broad regional sag or superbasin (Fig. 5.3-2) which, once established, persisted as a depocentre for 200-500 My as subsidence was reinvigorated by more localised interplate tectonism. Each supercontinent event thus resulted in the development of broad shallow intracratonic basins which provided an ideal setting for the accumulation of extensive sheets of platformal carbonates (Lindsay, 2002). At the same time the extrusion of large volumes of volcanics and associated carbon dioxide must also have had a significant impact on the atmosphere (Davies, 1995b) (discussion in section 5.2). The overall basin-fill architecture of these basins is broadly similar, consisting of a series of unconformity-bounded megasequences, each reflecting a major basinal episode (Lindsay and Leven, 1996; Lindsay and Brasier, 1998, 2002; Lindsay, 2002). When the Australian Precambrian carbon isotope data are compared with the global curve they match closely, indicating that it is likely to reflect a global signal (Figs. 5.3-3-5.3-6). The Precambrian secular curve for 313Ccarb is bimodal with major well defined but oscillatory peaks at c. 2.3-2.2 Ga and at c. 0.65 Ga (Fig. 5.3-6). In between the two modes, the secular carbon isotope curve is almost flat. A number of threads of evidence are converging to suggest that the evolution of the biosphere and atmosphere were driven forwards by the evolution of the planet through the release of endogenic planetary

400 C h a p t e r 5: E v o l u t i o n o f t h e H y d r o s p h e r e a n d A t m o s p h e r e

Fig. 5.3-6. Summary of carbon isotopic, tectonic and biospheric changes through time. Black bars = supercontinent events; hachured bars = intracratonic basins. Carbon isotopic curve is based on data from Schidlowski (1988), Straws and Moore (1992). Veizer et al. (1992), Des Marais et al. (1992), Kaufman and Knoll (1995), Buick et al. (1995), Knoll et al. (1995b), Karhu and Holland (1996). Brasier et al. (1996), Bartley et al. (1998), Brasier and Lindsay (1998) and Lindsay and Brasier (2000, 2002). For other sources see text.

5.3. Evolution of the Precambrian Atmosphere

401

energy involving mechanisms such as hydrothermalactivity, plate tectonics and the supercontinent cycle (see Lindsay and Brasier, 2002; Brasier et al., 2002) (sections 3.6 and 5.2). Data presented here, while not ruling out other possibilities, support a tectonic mechanism as being important. It has been argued that the conspicuously bimodal nature of the secular carbon curve indicates that the global reduced carbon reservoir grew episodically and this, in turn, may indicate that the atmosphere was oxygenated in a stepwise fashion (Des Marais et al., 1992) as a result of episodic burial of carbon during large scale tectonic cycles (Des Marais, 1994a, 1997) (see detailed discussions, section 5.2). In between the episodes of oxygenation, it has been suggested that tectonic activity was low and that CO2 in the ocean-atmosphere system reached a state of near equilibrium with respect to mass balance of the carbon cycle (Brasier and Lindsay, 1998; Lindsay and Brasier, 2000). It may be significant that the first major isotopic excursion follows the first appearance of intracratonic basins on the Australian continent (i.e., the Hamersley basin) and the first major platform carbonate intervals and that, associated with this basin, is the earliest known evidence of glaciation (Trendall, 1976) (sections 5.6 and 5.7). There is also a striking coincidence between the deposition of BIF and the first isotopic excursion indicating that iron was being rapidly removed from the ocean as oxygen began to circulate through the oceanic water column (Fig. 5.3-6) (see, however, models in sections 5.2 and 5.4). When the Western Australian Archaean and early Palaeoproterozoic data are placed in their stratigraphic and tectonic framework (Fig. 5.3-3) we find that the monotonic latestArchaean curve coincides with a tectonically quiescent period in which carbonates formed in a basinal setting on a craton surrounded by passive margins. The data are consistent with an Earth in which the carbon mass balance was in equilibrium (Lindsay and Brasier, 2002). T h e 613Ccarb curve began to oscillate following the onset of glaciation as the Pilbara and Yilgarn cratons began to converge during the Capricorn orogeny, suggesting periods of rapid carbon burial during continental dispersal. However, the major positive excursion is preserved in carbonates from back-arc basins formed as the ocean closed and subduction began. Because similar tectonic processes can be rccognised, not only in northern Australia but also on other early cratons, it can be argued that the carbon isotope excursions relate to supercontinent cycles (section 5.2) and to major periods of mantle overturn and superplume development (sections 3.2-3.4 and 3.9). In between c. 1.9 and c. 1.0 Ga, the secular carbon isotope curve is almost fiat (Fig. 5.3-6) (Des Marais, 1994a; Buick et al., 1995a; Kaufman et al., 1997; Brasier and Lindsay, 1998; Lindsay and Brasier, 2000). During this period supercontinent assembly was associated with small-scale mantle convection that resulted in vertical accretion of the crust and emplacement of anorogenic granites. While the release of endogenic energy in the form of heat was just as significant during this period, plate interactions on the Australian craton were far less energetic and there is little evidence of the formation of foreland basins or of crustal uplift. Consequently, there was little carbon burial and the interaction with the biosphere was minimal such that it entered a period of stasis from approximately 1.8-0.8 Ga. Current models of the ocean (see also section 5.2) suggest that to maintain the carbon mass balance requires relatively low levels of tectonic activity, which in turn suggests that availability of nutrients, such as nitrate, iron and phosphorus, were stable and

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low (see also section 3.2). Prolonged nutrient stability may therefore have exerted a major influence upon the evolution of the biosphere over this time interval. The pattern of secular variation of carbon for the last billion years of earth history is now relatively well known (Fig. 5.3-6) (see Kaufman and Knoll, 1995; Kaufman et al., 1997). Beginning in the Late Mesoproterozoic at some time between 1.3 and 1.2 Ga, the carbon isotopic record gradually became more active and the curve oscillated from around -1%o to as much as +4%o (Knoll et al., 1995a; Knoll and Canfield, 1998). Activity gradually increased until, by 800 Ma, in the early Neoproterozoic, the curve had returned to exceptionally high values (up to + 11%0, Brasier et al., 1996), similar to those seen during the Lomagundi event a billion years earlier. These developments appear to occur in parallel with the assembly and ultimate dispersal of Rodinia (sections 3.10, 3.11 and 5.8). The complex Neoproterozoic carbon isotope record has been attributed to the rapid expansion of the metazoa (section 6.2), to nutrient flux and to the locking up of lighter 12C during anoxic events (e.g., Berner and Canfield, 1989; Schopf and Klein, 1992; Knoll and Walter, 1992; Derry et al., 1992, 1994; Brasier and Lindsay, 1998; Brasier and Sukhov, 1998). The overall trend of high positive 613Ccarb values associated with Neoproterozoic rocks was first documented in Spitsbergen where the record was seen to be punctuated several times by major negative swings, which, in part, could be associated with glacial intervals (Knoll and Canfield, 1998). These major isotopic excursions have since been documented globally as well as in Australia (Fig. 5.3-5) (see also section 5.8) and have been used to delineate both significant events in earth history, and the Earth's biogeochemical history (e.g., Schidlowski et al., 1983; Hayes, 1983; Schidlowski and Aharon, 1992; Strauss et al., 1992a; Des Marais et al., 1992; Brasier et al., 1994, 1996; Des Marais, 1994a; Kaufman and Knoll, 1995; Calver and Lindsay, 1998; Knoll and Canfield, 1998). A final major negative swing in the curve occurs near the Precambrian-Cambrian boundary after which the amplitude of the oscillations declines (Brasier and Sukhov, 1998) before settling into the more modest oscillatory pattern of the Phanerozoic (Fig. 5.3-6). The secular carbonate carbon curve thus provides evidence of two major isotopic excursions separated by a billion years, both of which suggest major changes in the carbon cycle and hint at biospheric revolution (Brasier and Lindsay, 1998). The major isotopic excursions occurred during periods of crustal revolution as supercontinents assembled and dispersed suggesting that they were driven by large-scale tectonic processes. This further suggests that the biosphere and atmosphere were also driven forwards in a general way by the same large-scale processes (Lindsay and Brasier, 2000, 2002) (section 5.2). Conclusions

The Australian craton has evolved over a very long period of time (Archaean to Mesoproterozoic). During that period a series of basins developed on the craton as a response to the supercontinent cycle, preserving a record of the evolution of the planet's Precambrian atmosphere and biosphere. Carbon stable isotope data derived from platform carbonate rocks preserved in the basinal successions show that the oxygenation of the atmosphere occurred in two steps (c. 2.0 and c. 0.7 Ga) separated by at least a billion years. The data

5.4. P r e c a m b r i a n I r o n - F o r m a t i o n

403

suggest that evolution of the Precambrian atmosphere and biosphere was linked to the supercontinent cycle and that in a general way the evolution of both was driven forwards by the release of the Earth's endogenic energy resources (section 5.2). Without this release of endogenic energy the biosphere would have entered an evolutionary stasis and ultimately faced extinction (Brasier and Lindsay, 1998; Lindsay and Brasier, 2000).

5.4.

PRECAMBRIAN IRON-FORMATION

A.E TRENDALL AND J.G. BLOCKLEY

Introduction Iron-formation (IF), especially the form characteristic of the early Precambrian known as banded iron-formation (B IF), is unique among sedimentary rocks in that nothing closely resembling either it, or a plausible sedimentary precursor, is being formed in modern environments; most aspects of its origin therefore remain speculative. Even such basic parameters as its rate of deposition, and the depth of water in which it formed, remain unclear, and a range of disparate depositional mechanisms for it can be reasonably defended. Its relationship to palaeoclimate, palaeolatitude, and volcanism are also unknown. IF is not, of course, simply a challenging sedimentological problem. Since the late 19th century it has been the world's main source of iron ore, and has thus made a major contribution to the fabric of industrial society through the iron and steel industries which depend upon it.

Iron-Formation: Definition, Nomenclature and Classification Iron-formation is an iron-rich sedimentary rock mainly confined to the early Precambrian stratigraphic record. The name originated in the Lake Superior area as a contraction of the "iron-bearing formation" of Van Hise and Leith (1911). James (1954) later defined it as "a chemical sediment, typically thin-bedded or laminated, containing 15 percent or more iron of sedimentary origin, commonly but not necessarily containing layers of chert". Despite James' (1954) lower limit of 15%, most IFs contain about 30% of iron by weight (Davy, 1983), usually as oxides, so that hematite (Fe203) or magnetite (Fe304) together constitute roughly half of the rock. Most of the remaining half consists of silica; this normally occurs as microcrystalline quartz, usually called chert. Carbon dioxide is present as a significant minor constituent in many BIFs, and is a major constituent in some, but all other oxides (e.g., A1203, MgO, alkalies) are typically minor, and "trace" elements are just that--these are chemically very "clean" rocks. Apart from carbonates (dolomite, ankerite and siderite), other minerals that may be present in minor to locally significant quantities are silicates (stilpnomelane, chlorite, greenalite, minnesotaite, riebeckite) and sulphides (pyrite, pyrrhotite), while progressive metamorphism may produce assemblages containing cummingtonite-grunerite, clino- and orthopyroxene, fayalite and almandine (Klein, 1983). The Precambrian Earth: Tempos and Events Edited by EG. F+riksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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All the mineral components of IF (and particularly of BIF, see below) are typically finegrained, which accounts for its characteristic hardness, and resistance to both hammering and weathering. In landscapes cut into Precambrian rocks throughout the world, IFs characteristically form conspicuous resistant ridges. The two major mineral constituents (quartz and iron oxides) are normally concentrated in alternating iron-rich and silica-rich bands on the mesoscopic scale; often these are brightly coloured--red, black, or white. A number of names have been applied to IF in different continents (Trendall, 1983). Examples include the "itabirite" of Brazil, the "BHQ" (banded hematite quartzite) of India, the "taconite" of the Lake Superior ranges, the "ironstone" of South Africa, and the "jaspilite" of Australia. All these are now subsumed under the generic name IF, but will no doubt continue to be applied locally. A number of classification systems have been proposed for IF but none has proved fully satisfactory. James (1954), distinguished four "facies" of IF in the Lake Superior area: oxide facies, carbonate facies, silicate facies and sulphide facies. He suggested that these are lateral depth-related stratigraphic equivalents, but this conceptual relationship has neither been demonstrated in the Lake Superior area nor been shown to hold elsewhere. While these names will remain useful for IF lithologies with the corresponding mineral and chemical compositions, it should be understood that the environmental implications initially associated with them are speculative. Gross recognised two types of siliceous IF, a Lake Superior type and an Algoma type, "based on the characteristics of their depositional basins and the kinds of associated rock" (Gross, 1980, p. 215). He described the Lake Superior type IFs as "deposited in near-shore continental-shelf environments and ... associated with dolomite, quartzite, black shale and minor amounts of tuffaceous and other volcanic rocks", whereas Algoma type IFs were "apparently formed close to volcanic centres" and "are consistently associated with greywacke sedimentary units and volcanic rocks". Gross (1980, his table 2) made it clear that the classification is one of depositional basins in which IF was deposited, rather than of IFs as lithological types. Misunderstanding of this point has resulted in the allocation of some IFs to different types by different authors. For example, such major IFs as the Dales Gorge Member of the Brockman Iron Formation, in the Hamersley basin, have been classed both as Lake Superior type (Gross, 1980) and Algoma type (Dimroth, 1976). This confusion is compounded by Gross's ( 1991 ) later reclassification of the Hamersley basin BIFs as Algoma type. The continued use of "Lake Superior type" and "Algoma type" is not recommended, as the terms carry no clear meaning (supported by Ohmoto, section 5.2). Other classifications have been put forwards (e.g., Kimberley, 1978; Beukes, 1980) but have not become widely used. Trendall (2002) suggested that the most significant division of IF is that between banded iron-formation, or BIF, which includes most occurrences older than c. 2.0 Ga, and the type of IF characteristically present in the circum-Ungava belt of North America, which is distinguished as granular iron-formation (GIF). This twofold lithological division is followed here, and is emphasised under a later heading. Most stratigraphic units of IF consist predominantly of either BIF or GIF, although a few contain a mixture of both types. Unfortunately, in much current and historical literature on IF the term BIF is used as a synonym for IE

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Global Occurrence

IFs are present on all continents, in all major areas of Precambrian rocks, and IFs closely similar in composition and lithology to those typical of the Precambrian also occur in the Palaeozoic (Kalugin, 1969, 1973). IFs of the older cratons include, notably, the oldest known BIF, at Isua (section 2.2) in Greenland, whose age is about 3.8 Ga. BIFs also occur in the early greenstone belt (sections 2.4, 4.4 and 7.3) sequences of all the main old cratons. Examples include the Abitibi belt (sections 2.4 and 4.3) of the Superior Province, the greenstone belts of the Yilgarn and Pilbara cratons of Australia (sections 2.5-2.7), the greenstone belts of the Baltic shield (Finland and Karelia) (section 3.9), those of the North China craton, the Amazon craton of Brazil, and the Kaapvaal and West African cratons. The BIFs of Krivoi Rog and Kursk in the Ukraine probably also belong here. Older Precambrian BIFs that are not components of greenstone belts occur in four of the Gondwana continents (South America, southern Africa, India and Australia). These comprise gently dipping sequences that form extensive and conspicuous topographic plateaus, and Trendall (2002) has called them the "Great Gondwana BIFs". They include, in Brazil, the Carajfis Formation of the Grfio Pardi Group of the Amazon craton and the Caua Itabirite of the Itabira Group of the Silo Francisco craton. In South Africa the Kuruman Iron Formation and some overlying units of the Transvaal Supergroup in the Griqualand West basin and the Penge Iron Formation of the Transvaal basin belong in this category. In India the Mulaingiri Formation of the Bababudan basin, in the Karnataka craton, also belongs here. And finally the BIFs of the Hamersley basin of Western Australia are also included among the Great Gondwana B IFs. IFs also occur in basins associated with younger Precambrian terranes. Of these the best known are the circum-Ungava IFs of Canada and the United States, which include those of the Lake Superior ranges (Mesabi, Cuyuna, Menominee, Gogebic and Marquette). These consist mainly of GIF, although some also contain BIE Finally, there is a special category of latest Precambrian IFs, which include Urucum in Brazil, Rapitan in the Yukon, and those of the Damara Belt in southern Africa. Description and Documentation: The Available Literature

Morey (1983) has traced the history of North American publication on IE Iron mining following the discovery of IF in northern Michigan in 1844 established the Lake Superior area as the birthplace of IF studies. By the early 20th century the outstanding monographs of the United States Geological Survey (e.g., Van Hise and Leith, 1911) had established a tradition of excellence in IF publication which was followed in the publications of Gruner (e.g., 1922, 1924, 1946) and James (e.g., 1954, 1958, 1966), and in those of Gross (e.g., 1965, 1967, 1968, 1973) covering IFs in Canada. Earlier literature from other areas is referred to in Trendall and Morris (1983).

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After the end of the Second World War in 1945 iron ore became a globally transportable commodity, and interest consequently increased in the world wide occurrence of IF, especially in southern hemisphere continents. The lifting in 1960 of an Australian embargo on the export of iron ore, which had been imposed in 1939 for political reasons, led to intense interest in the Hamersley Range area of Western Australia. This proved to be the largest known occurrence of IF on any continent, with significant differences from IF of the Lake Superior area, and led to the Hamersley basin becoming the new focus of IF research. Although the total literature of IF is now vast, a relatively small number of dedicated and relatively recent publications are available that condense and summarise it; the enthusiasm of Harold James led to the appearance of many of these. Extensive references to work prior to their publication dates can be found in James and Sims (1973), UNESCO (1973), Mel' nik (1982), Trendall and Morris (1983), Radhakrishna (1986) and Appel and La Berge (1987). Later papers that provide excellent brief accounts of most aspects of IF occurrence and deposition include those of Beukes and Klein (1992), Klein and Beukes (1992) and Morris (1993), while a contribution by Trendall (2002) includes more recent developments. Gross (1991) has provided a comprehensive analysis of the published literature on IF prior to that date. First-Order Genetic Issues

Although some early papers on the Lake Superior area referred to IFs as rhyolitic (Wadsworth, 1880; Winchell, 1900) they were accepted as sedimentary rocks by the early 20th century (Van Hise and Leith, 1911). Most subsequent workers have also accepted that, for B IF, the sedimentary precursor was a chemical precipitate of essentially similar composition to the final lithified rock. But this acceptance has not been universal. Lepp and Goldich (1964), argued from bulk chemical compositions that the silica component of many Precambrian IFs represented replaced primary carbonate. Dimroth (1975) similarly used the textural similarity of IF to carbonates to suggest that many IFs were diagenetically replaced carbonates, a concept echoed by Kimberley (1974). And more recently, Krapez et al. (2002) have suggested that the sedimentary precursors of some Hamersley Group BIFs may have been "hydrothermal muds" derived from plumes associated with sea-floor volcanism. Trendall and Blockley (1970, p. 268) argued that, for the BIFs of the Hamersley basin, three factors were inconsistent with extensive post-depositional chemical modification. These were: (a) the enormous amount of iron makes it unlikely that this, at least, was derived from some extraneous source; (b) the chemical uniformity of all the BIFs of the Hamersley Group; and (c) post-depositional modification would be unlikely to produce the remarkable lateral homogeneity of the BIFs. These points remain valid; and in particular if any IF represents sedimentary material of radically different composition which has suffered gross chemical transformation into IF after deposition, then some example should by now have been found and described where some vestige of the precursor which has locally escaped modification can be shown to grade laterally into IF.

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Stratigraphic Issues BIFs and GIFs: banding bedding, and lateral continuity A distinction has already been noted between BIF, or banded iron-formation, and GIF, or granular iron-formation. The characteristic features of BIF confer on it an identity that is never in doubt. It is called "banded" because of the concentration of the two major mineral constituents (quartz and iron oxides) into well defined iron-rich and silica-rich bands on the mesoscopic scale. Trendall and Blockley (1970) introduced the term "mesobanding" for such bands in the course of a detailed description of the BIF of the Dales Gorge Member of the Brockman Iron Formation of the Western Australian Hamersley basin. Silica-rich mesobands (Fig. 5.4-1a) are usually called "chert", and consist mainly of a tight mosaic of microcrystalline quartz. Some fine-grained iron carbonates, silicates or oxides are normally also present, and most chert mesobands have a significant (5-25%) iron content. Chert mesobands of the Dales Gorge Member are mainly between 5 mm and 15 mm thick, with a mean thickness of about 8 mm, and they make up about 60% of the total B IF volume. The name "chert-matrix" was applied to the mesobands of finegrained iron-rich material which alternate with those of chert, and form a matrix for them; the mesoband contacts are usually quite sharply defined (Fig. 5.4-lb). In the Dales Gorge Member chert-matrix mesobands have a mean thickness of about 10 mm, and make up about 20% of the total B IF volume. They have a mean iron content of about 40%, and consist of a fine-grained aggregate of quartz, iron oxides (magnetite or hematite) and other Fe-rich carbonate and silicate minerals. The aggregate has a finely streaky, irregularly laminated texture parallel to the mesoband margins. By decrease in silica content chert-matrix grades into mesobands of magnetite, and there are rare mesobands of iron-rich carbonate or silicate. Mesobands of chert, chert-matrix and magnetite together make up over 90% of the rock by volume. Trendall and Blockley (1970) also introduced the term "microbanding" for a regular small-scale lamination within many, but not all chert mesobands of the Dales Gorge Member. Microbands are defined by a concentration of some Fe mineral (either hematite, magnetite, carbonate, stilpnomelane or some combination of these) within the pervasive silica framework, a single microband being defined by an iron-rich component in which these minerals are concentrated and an iron-poor component from which they are effectively absent (Fig. 5.4-la and b). Microbanded chert mesobands commonly contain up to 40 microbands, with a maximum recorded of 236. The thicknesses of successive microbands vary little within any one microbanded chert mesoband, but microband thickness may differ substantially from one chert mesoband to another; microband thicknesses are mainly in the range 1.6-0.2 mm, with a mean about 0.5 mm. Following the introduction of the term microbanding for the type of lamination described above within chert mesobands of the Dales Gorge Member, and its wider application to other Hamersley Group BIFs, there was a tendency to apply it to any very thin lamination within B IE In response to this misuse, the alternative terms "aftband" (e.g., Trendall, 1983b), "aftvarve" (e.g., Morris, 1993) and "BIF-varve" (Morris, 1993) have been used to emphasise that the laminations to which Trendall and Blockley (1.970) originally applied the name "microband" were of a special

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and restricted type. Morris (1993, p. 261-263) has discussed this nomenclatural point with great clarity, and on that basis the future use of "BIF-varve" for the "microband" of Trendall and Blockley (1970), and of this review, would seem a good choice. The mesobanding of most B IFs follows the pattern described above for the Hamersley BIFs rather closely in respect to thickness, sharpness of mesoband boundaries, and alternation of Fe-poor (chert) mesobands and Fe-rich (chert-matrix) mesobands. Microbanding is generally less common, but is present in equal abundance in the Transvaal Supergroup BIFs of South Africa (e.g., Beukes, 1973). Among other BIFs it is present in chert mesobands of the Carajas Formation of Brazil and the Mulaingiri Formation of India, although in neither is it as conspicuous as in the Hamersley BIFs. Trendall (1973a, his Fig. 5) has shown it in other BIFs, and Gole (1981) has also figured excellent examples from BIFs of the Yilgarn craton of Western Australia; Matin and Mukhopadhyay (1992) describe microbands from an IF in the Sandur Schist Belt of the Indian Karnataka craton. The characteristic mesobanding of BIF, with or without microbanding, is not present in GIE There is a comparable alternation of iron-rich and iron-poor bands, but these are typically coarser and much less regular, and resemble the bedding of many epiclastic sediments where coarser and finer components are intercalated (Fig. 5.4-I c). The coarsely crystalline cherts tend to be wavy or lenticular, probably reflecting the rippled nature of the primary sandy material. Both iron-rich and silica-rich bands may be granular, more particularly the latter. The iron-rich bands of GIF, as the name implies, often consist of a close-packed and lithified mass of granules or ooliths, about l mm across (Fig. 5.4-ld). These are made up of iron oxides with or without quartz, and their interstices are filled by the same minerals, but usually with a lower iron content. The granules have the appearance of primary depositional components, and the material has been referred to as "a special type of sandstone" (Mengel, 1965). Apart from the obvious differences in stratification between B IF and GIF, these lithologies also differ in other important respects. Firstly, is the presence or absence of currentgenerated structures: whereas current-generated structures such as cross-bedding and ripple marking are commonplace in GIF, unequivocally current-generated structures have yet to be described from any BIE Secondly, is the lateral continuity of banding. Trendall and Blockley (1970) described and illustrated lateral correlation of subcentimetre mesobands of the Dales Gorge Member over distances of up to 300 km, and of microbands over

Opposite: Fig. 5.4-1. Types and scales of banding in BIF and GIE (a) Drillcore of BIF of the Dales Gorge Member, Brockman Iron Formation, Hamersley Group, Western Australia, showing silica-rich (light) mesobands of chert alternating with iron-rich (dark) mesobands of chert-matrix. (b) Thin-section of drillcore similar to that in (a). Note the relatively sharp edges of mesobands, and the regular microbanding within chert mesobands. (c) Drillcore of GIF from the Frere Formation, Earaheedy Group, Western Australia, showing interlaminated mesobands of granular chert (light) and dark iron-rich material. (d) Thin-section photograph of a granular chert mesoband, showing rounded iron-rich (hematitic) ooliths or granules.

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80 km, and Ewers and Morris (1981) later documented mesoband correlation over 130 km. It remains a reasonable supposition for the Hamersley basin that some microbands have basin-wide correlatability. Comparable lateral continuity of small-scale stratification has not been described from GIE Thirdly, there are major differences in the association with clastic material; and finally, BIF and GIF are largely time-restricted; both of these aspects are expanded under later headings.

Relationship to other lithologies BIFs generally form discrete, sharply bounded units mostly less than 100 m thick, as distinct from being thinly interstratified, or interdigitating, with other lithologies. No BIF of substantial thickness has ever been shown to be laterally gradational into another sedimentary lithology. They are commonly associated with a variety of volcanic rocks, as well as shales and carbonates, but also occur in association with clastic sequences. Isley and Abbott (1999) have convincingly documented a temporal distribution of BIF deposition with mafic volcanic rocks (see also section 3.2). There are exceptions to the generalisation that BIF is never thinly interstratified with other rocks. Thus Eriksson's (1983) idealised cross-sections of early Archaean sequences of the Pilbara and Barberton greenstone belts, show IF intercalations less than a metre thick capping the graded turbidite beds of clastic units, and in other special situations. Similarly, in the IF ("banded ironstone") of the Mount Belches area of the Western Australian Yilgarn craton, Dunbar and McCall (1971) used a discrete BIF unit ("Santa Claus Ironstone Member") as a structural marker, although noting that minor occurrences of IF less than a metre thick are intercalated within the adjacent, mainly clastic, sequence. Similar thin intercalations of IF have also been noted within the clastic sediments of the Beardmore-Geraldton greenstone belt of Ontario by Barrett and Fralick (1985, 1989). Despite these examples, it remains the case that major IF units of both areas are discrete, well defined, and relatively thick and extensive. GIF also tends to form discrete well-defined units. However, in terms of lithological association, GIF is commonly interstratified with coarse- or medium-grained epiclastic sediments, and a volcanic association, although usually present, is smaller relative to the volume of IF present.

The interpretation of banding The interpretation of the banding of BIF is crucially important for an understanding of its depositional environment, and the topic is therefore treated here at some length. Prior to the 1960s, apart from the work of Moore and Maynard (1929), few workers had focused directly on the problem of the origin of the banding, and none had linked a hypothesis for the origin of banding to textural characters of particular BIFs. Based on the extraordinarily regular spacing of microbands within microbanded chert mesobands, Trendall and Blockley (1970) postulated that each iron-rich/iron-poor microband couplet might have resulted from one year of chemical deposition: each microband was an annual layerw a chemical varve. The primary material was seen as a layer of finely particulate, possibly colloidal, water-rich material about 5 mm thick consisting largely of silica and iron

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hydroxide, with the iron component already differentiated within the silica. Based on the wide variation of microband thickness in different chert mesobands, and an inverse correlation in microbanded chert mesobands between microband thickness and total iron content, they then proposed, firstly, that finely microbanded chert mesobands were simply the more compacted equivalents of coarsely microbanded cherts, and secondly, that while the iron had remained fixed during the compaction process some of the more soluble silica had been removed in the course of diagenetic dehydration. From the marginal relationship of microbands within discontinuous nodule-like chert bodies, which they called "pods", they argued that the chert-matrix mesobands similarly represented the result of extreme compaction, and associated silica depletion, of identical primary precipitate to that from which chert mesobands were formed. Thus Trendall and Blockley (1970) saw mesobands as secondary structures formed during burial and compaction of a succession of annual layers of primary precipitate. Ewers and Morris (1981) agreed with Trendall and Blockley's (1970) interpretation of microbands as varves, and suggested also that fine irregularities within microbands may correspond to minor seasonal fluctuations in such factors controlling deposition as solar radiation, temperature, and biological activity. But they rejected the concept that mesobanding developed during early diagenetic compaction and removal of silica. Their argument was based on three main points (Ewers and Morris, 1981, p. 1945): (i) vertical (upwards) passage of silica-bearing fluids would probably have caused some disruption of the fine structures of the banding within the overlying BIF; (ii) insufficient connate water would have been available within the BIF to dissolve all the silica required to be removed from the BIF in Trendall and Blockley's (1970) model; (iii) silica removed from lower mesobands would be unlikely to reprecipitate at higher levels, to cause thickening, without some disruption of internal structure. They consequently proposed that mesobanding was controlled by primary deposition, with each mesoband representing "a period of several years, perhaps tens of years, in which the conditions for precipitation, water composition, etc., were reasonably stable", "the transition from one mesoband to another 'arising' from a change in these conditions" (Ewers and Morris, 1981). They did not discuss in detail the possible controls of those changes, but Morris and Horwitz (1983) extended the earlier hypothesis to include the concept that they were related to "the pulsed output of a large oceanic rift or hot spot" (see analogous ideas in Ohmoto, section 5.2). In a later and more complete presentation of these ideas, Morris (1993, p. 268) specified that at times of low hot-spot or midocean ridge (MOR) hydrothermal activity, water in the BIF depository had a relatively low iron content, and the precipitated material was mainly silica; the microbands within consequent chert mesobands were caused mainly by direct photo-oxidation. By contrast, "during periods of more violent MOR or hot-spot activity, higher levels of Fe(II) reached the depository by convection-driven upwelling, with increased nutrients ... possibly triggering a parallel growth of organisms" (Morris, op. cit.); the iron-rich chert-matrix mesobands were deposited during these periods. Although the interpretation of microbands as varves is now generally accepted, it remains the case that this interpretation is still only a hypothesis based on the perception that such regularly repetitive layers reflect some equally regular rhythm of the depositional

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environment; and that the year seems the most likely control of such a rhythm. Trendall and Blockley (1970, p. 257) considered the possibility that microbands were diurnal, but dismissed this idea on the joint grounds that it would imply an unrealistic depositional rate (6 years per foot, or 5.08 km/Ma in the units now preferred) (section 7.11), and that daily effects throughout such a large depositional basin were unlikely. Cisne (1984, p. 484) revisited the possibility of microbands being diurnal, and suggested that the gravitational effect of the weight of BIF on the underlying mantle may have caused exceptionally fast deposition. A feature of microbanding particularly well displayed in one BIF unit (Weeli Wolli Formation) of the Hamersley basin, is a cyclic variation in thickness of the individual microbands. Trendall (1973b) suggested that the cyclicity may represent a 23.3-year cycle equivalent to the modern double sunspot (Hale) cycle, but Walker and Zahnle (1986) reinterpreted it as an expression of the lunar nodal cycle. Most recently, Williams (2000) has pointed out that microband sequences in the Weeli Wolli Formation are remarkably similar, in terms of their cyclic patterns, to finely laminated late Neoproterozoic clastic sediments from South Australia; and in respect to those he has shown (Williams, 2000, his Fig. 6) that the correspondence of their cyclicity with modern tidal records argues strongly for their identification as tidal rhythmites (see also sections 5.9 and 7.5). He concluded that microbands of the Weeli Wolli Formation may represent either diurnal increments (following Cisne, 1984) grouped in monthly cycles, or semidiurnal increments grouped in fortnightly cycles, or fortnightly increments arranged in annual cycles. Apart from the need for more spectral analyses of BIF microband and other laminated sequences throughout the stratigraphic record, critical evidence for the origin of microbands will be provided by a precise determination of the depositional rate of BIF; this is discussed later. The mesoscopic scale banding of GIF is generally accepted (e.g., Zajac, 1974; Morey, 1983) to have been generated by reworking, in a shallow-water high-energy environment, of fragmented iron-rich precipitates of uncertain original nature. Although there has been no detailed modelling, it is usually implied that the process of fragmentation would have produced a mix of coarse and fine debris, and that the sorting processes that led to the preferential concentration of these components into bands were similar to those normally operating in other clastic sedimentary environments. Detailed discussion of these is outside the scope of this review. Distribution of IF in time During the 1960s and 1970s it was widely believed that most major IFs were deposited around 2000 Ma (e.g., Goldich, 1973; Cloud, 1973; James and Sims, 1973). Advances in geochronology, as well as greater knowledge of the global occurrence of IF, made this position untenable, and it was replaced by the modified concept that, although there was indeed an early Precambrian peak, relatively small amounts of IF, heralding their later abundance, were deposited locally from about 3.8 Ma (Isua) onwards. A virtual end of significant IF deposition by 1.5 Ga was also sustained within this concept, with a minor peak in latest Neoproterozoic time. Gole and Klein (1981), James (1983) and Klein and Beukes (1992) are among the authors who have published diagrams illustrating this change of IF abun-

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dance with time, and Trendall (2002, his Fig. 6) has shown a slightly modified diagram, reproduced here as Figure 5.4-2a. In almost all such diagrams, including Figure 5.4-2a, the Y-axis is unquantified, indicating that it represents a subjective estimate of relative abundance. James (in James and Trendall, 1982) made an attempt at quantification, illustrating the difficulties of estimating accurately the total deposited iron content of most IFs. A further difficulty is that very few IFs were (and still are) precisely dated, estimates of their age being largely based on the ages of datable rocks above and below them in the sequences in which they lie. Though subjective, Figure 5.4-2a is a reasonable representation of the distribution of Precambrian IF distribution in time. The subjectivity is emphasised in Figure 5.4-2b, as explained in the caption. The apparent objectivity and quantification given by the compilation of Isley and Abbott (1999) is fatally flawed, in that any named lithostratigraphic unit of IF is accorded equal status, regardless of size; their contribution is significant, however, in demonstrating the association of many IFs with mafic igneous rocks (section 3.2). Figure 5.4-2a shows characteristic differences in the IF deposited at different times during the Precambrian. Throughout early Precambrian time individual IFs tend to be thinner, laterally less extensive, and often closely associated with volcanism in greenstone belts. There then seems to be a peak at c. 2.5 Ga, to which a very significant contribution is made by two of the Great Gondwana BIFs: those of the Hamersley and Transvaal-Griqualand West basins. There also appears to be a later period of abundant IF deposition, possibly around 1.8 Ga, to which the main contributors are the GIFs. After this period there is a long hiatus in the later Precambrian when IF deposition ceased; and finally there is a latest Precambrian (Neoproterozoic) scattering of mainly small IFs of various types (see section 5.6 for discussion of the temporal association of BIF and Precambrian glaciations).

Depository Issues Basins of iron-formation deposition The title "Three great basins of Precambrian banded iron-formation deposition: a systematic comparison" (Trendall, 1968) was chosen to emphasise that the study of IF required attention to IF depositories, not to the IF alone. This approach was continued by Trendall and Morris (1983), who allocated 300 pages of their book to six selected basins in which IF was a significant component. It was based on the search for common characters of basins with significant IF deposition, in the hope that these might indicate why the IF occurred in those basins, and not in others. Trendall (2002), taking into account that Precambrian IF seemed to have occurred in a great variety of tectonic basin types over a vast period of time, concluded that it may be more productive to postulate a minimal set of necessary and sufficient conditions for IF deposition to take place in any early Precambrian volcanosedimentary depository. He suggested that, for older Precambrian BIF (Fig. 5.4-1a), these might be: (a) tectonic stability for long (c. 106 years) periods during their evolution, to allow enough time for the typically discrete units of BIF to be formed; (b) sufficient water depth to avoid contamination with epiclastic material, and to be free of bottom disturbance; (c) shapes that permitted deep ocean water to circulate freely into and out of them. For the

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Fig. 5.4-2. (a) Schematic abundance of IF through Precambrian time, after Trendall (2002). Time runs from left to right on the X-axis, and the numbers show Ga before present. The Y-axis shows abundance. It is important to realise that such a representation is highly subjective, as illustrated in (b). (b) Expanded part of diagram (a), between 2.4 and 2.7 Ga, with actual tonnages of Fe laid down originally in the individual BIF units of the Hamersley Group, in the Hamersley basin of Western Australia. This diagram is included to indicate that actual IFs, if plotted on (a), would be represented by a series of individual hair-line spikes. The letters above the columns indicate: M--Marra Mamba Iron Formation; D--Dales Gorge Member; J--Joffre Member; W/B--combined Weeli Wolli Formation and Boolgeeda Iron Formation. The X-axis length of each rectangle represents duration.

younger GIFs, conditions (a) and (c) were accepted as necessary, but water depth for GIF formation was accepted as shallower. This change of viewpoint, in which a "basin of iron-formation deposition" is not seen as a depository with a unique set of characteristics, any more so than a "basin of sandstone deposition", has much to c o m m e n d it. Instead, just as sandstone is a c o m m o n lithology present in a wide range of basin types, so also is IF, the key to its presence in any specific basin being an architecture which provided adequate water depth, and appropriate water

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circulation to and from the deep ocean. The significance of these aspects is discussed under other headings.

The depositional rate of BIF The uniformity in banding characteristics of BIF suggest that it may have a characteristic rate of deposition (section 7.11 gives a more general discussion of sedimentation rates; see also section 5.8). From the genetic hypothesis for banding already outlined, in which microbands of the Hamersley Group BIF represent chemical varves, and mesobands were produced by differential compaction during diagenesis, Trendall and Blockley (1970) calculated that 225 g of Fe were precipitated per square metre of basin area each year. Using this figure, and also the known total Fe content and density of BIF, 1 cm of (compacted) BIF of the Dales Gorge Member would have taken 44 years to accumulate, equivalent to a depositional rate of 270 m Ma - l . Morris (1993, p. 276) reviewed those estimates and recalculated them as between 227 m Ma -l and 87 m Ma -1, compared to his own preference of 893 m Ma -1 , based on a different interpretation of microbanding (and mesobanding). Klein and Beukes (1989), also accepting the varve hypothesis for microbands, estimated a rate of 568 m Ma -! for the closely similar BIF of the Kuruman Iron Formation of South Africa. Trendall (2002) later revisited the Dales Gorge Member evidence, and concluded that from the data of Trendall and Blockley (1.970) a better value lay between 23 and 230 m Ma -1. Trendall (1998) compared these rates with those derived from high-precision U-Pb zircon (SHRIMP) ages from intercalated tufts within the Hamersley Group, and concluded that the data available imposed constraints of between 19 and 225 m Ma -1 . Pickard's (2002) SHRIMP data from the Joffre Member suggested a rate of 33 m Ma -1 for the BIF of that unit; but unfortunately the errors associated with the SHRIMP data mean that they set no upper limit on depositional rate, but did set a lower possible limit of 15 m Ma -l . Hopefully, further technical advances will reduce these limits, but at present the fact must be faced that the depositional rate of Hamersley Group B IF cannot be determined with sufficient precision to discriminate between the different hypotheses for origin of microbands, already discussed. Basin water depth Taken together, the general absence of clastic material within BIF, and the lack of currentgenerated structures, constitute a first-order argument for their deposition in deep water, distant from land. Neither of these features is definitive. For example, Trendall and Blockley (1970) proposed a desert climate, and consequent lack of surface drainage, as a reason for the paucity of clastic material associated with the BIF of the Hamersley Group; Morris and Horwitz (1983) have suggested deposition on an offshore platform as an alternative explanation of this feature. Trendall (2002, p. 44, his Fig. 5) has drawn attention to a further argument for deep-water deposition of BIFs in the Hamersley basin, based on the total depositional history of the basin. Throughout deposition of the c. 6 km-thick lower volcanic succession of the basin, it is demonstrable from combined field and geochronological evidence that deposition was both relatively rapid (c. 90 m Ma - l ) and synchronous with basin sinking, since facies are either shallow-water or terrestrial. The overlying

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Hamersley Group has a much lower integrated depositional rate, so that if the floor of the basin continued to sink at the same rate, an increase in water depth, coinciding with the first appearance of BIF, was inevitable. Simonson and Hassler (1996) have argued independently on the need for deep water for early Precambrian B IFs more generally, and have related their deposition to global sea level high-stands (sections 8.1 and 8.2 for general concepts). Their argument hinges on "the widely accepted idea.., that the deposition of large iron formations was made possible by a reservoir of dissolved iron in the deep oceans, while surface waters contained relatively low concentrations of dissolved iron" (Simonson and Hassler, 1996, p. 666). This concept fits the stratigraphy of the Hamersley basin very well, where clastic and shale units between the major BIF units may be credibly interpreted as related to periods of shallower basin water.

Miscellaneous Matters The supply of iron, and the volcanic association The quantities of iron present in Precambrian IF defy the imagination: the largest single lithostratigraphic unit of BIF known (the Joffre Member of the Brockman Iron Formation of Western Australia) contained at least 4.3 x 1013 tonnes of iron at the time of deposition; this is roughly as much as would be produced in 80,000 years of global iron mining at the present annual rate. Where this came from is a key question for understanding IE There have been four main answers: wind-blown dust, dissolved iron in rivers, volcanic sources close to the sites of IF deposition, and global seawater. Two arguments weigh strongly against Carey's (1976, 1996) suggestion of windblown dust. Firstly, it is hard to explain why iron, any form of which would be expected to be in a heavy component of dust, should be preferentially winnowed from a land surface. And secondly, Carey (1996) linked the formation of all the major Precambrian BIFs to a single brief wind-dominated climatic episode, and this is inconsistent with the time-distribution of IE The derivation of iron by weathering of the rocks of continental areas adjacent to the depository, and associated transport into it by rivers, was widely assumed in early papers on the Lake Superior IFs (Leith, 1903), but Van Hise and Leith (1911) were already aware that such a model involved great difficulty, in that if enormous quantities of iron were to be selectively extracted from continental crust then a much larger quantity of remnant material would be produced, and it is not clear how this could have been disposed of. The problem has been debated repeatedly in later literature without a compelling answer to the problem (e.g., Lepp and Goldich, 1964, and discussion by Trendall, 1965; Trendall and Blockley, 1970, pp. 273-275; Garrels, 1987, and discussion by Morris and Trendall, 1988). A terrestrial source of iron does not seem feasible, even without the positive evidence in support of a volcanic association. There is a wealth of data from both the Sm-Nd isotopic system and from REE content (Fryer, 1983; Miller and O'Nions, 1985; Gerlach et al., 1988; Klein and Beukes, 1989; Derry and Jacobsen, 1990; Bau and M611er, 1991, 1993; Alibert et al., 1991 ; Danielson

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et al., 1992; Alibert and McCulloch, 1993; Morris, 1993; Manikyamba et al., 1993; Arora et al., 1995) consistent with the derivation of the iron of IF from a volcanic source. Even before this evidence was available, such a source had been widely suggested (see Gross, 1991, p. 55), and if it is accepted there remains the question whether the volcanism was immediately associated with particular basins of IF deposition, or whether it was buffered through the global ocean. Trendall and Blockley (1970) preferred a local source for Hamersley Group BIFs, on the basis of the abundance of associated igneous activity (over 20% of the thickness of the group consists of igneous rocks), and the quantitative feasibility of deriving the iron from the evidence of the rate of iron effusion from some recent volcanic areas; Barley et al. (1997) have recently re-emphasised the igneous association of the Hamersley Group BIFs. But there is no compelling geological evidence to link either the Hamersley Group B IFs, or indeed any other IF, directly to a specific volcanic centre. In addition, there remains at least one major basin of IF occurrencewthe Quadrilatero Ferrifero--with no volcanic rocks yet identified in close association with the IE Finally, it is clear that modern volcanic-associated iron-rich sea-floor deposits, such as those of the Red Sea, are "utterly unlike" IF (James, 1969). For this combination of reasons ocean water (see also section 5.2) appears to be the most likely source of iron for global IF deposition.

The precipitation mechanism and relationship to life Trendall and Blockley (1970, p. 272) argued, for the BIFs of the Hamersley basin, that "all minerals in the existing iron formation are secondary, and that little can be deduced about their parent materials from their present textures", although this is disputed by Morris (1993). They used the term "secondary" to indicate a distinction from the unknown "primary" materials that were precipitated. Klein (1983, p. 422) expressed the same view in the words "it is almost impossible to obtain first-hand information on the primary phases that were originally precipitated...". The most credible first-order model is that the parent precipitate of the finally lithified BIF was a hydrous silica-iron gel. What then was the mechanism of precipitation? Four main possibilities have been proposed. The first is evaporative concentration, and the remaining three all involve the oxidation of dissolved ferrous iron, but by three different mechanisms--photosynthesising microbiota, direct oxidation by bacteria, and oxidation by ultraviolet solar radiation. Trendall and Blockley (1970, p. 283) referred to the possible role of evaporation in maintaining a high iron concentration in the water of the Hamersley basin, and Trendall (1973a) argued that BIFs could represent a style of varved iron-bearing evaporite restricted to the Precambrian, analogous to the varved saline evaporites restricted to the Phanerozoic. However, the wide time spread of BIF deposition in a variety of basin types decreases the likelihood of such a relationship. Trendall and Blockley (1970, p. 283) also discussed the possible role of photosynthesising microbiota in both effecting, and mediating the annual precipitation of iron. Cloud (1973) argued for an interaction between "primitive oxygen-releasing photosynthesisers" (cyanobacteria) and dissolved ferrous iron in the formation of IFs. He presented an elegant and detailed hypothesis which represented the appearance of such organisms as an evolutionary step which followed from the very early appearance of chemoautotrophs.

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The hypothesis went on to suppose that the earliest photoautotrophs which evolved from these would have had their development restricted because they lacked (like their chemoautotrophic forbears) buffering mechanisms against the oxygen which was produced by their metabolism. When they acquired such buffering mechanisms, Cloud (1973) argued, the immediate near-explosive increase in their numbers led to the deposition of vast amounts of IF at 2.1-2.0 Ga. Cloud's hypothesis has been re-examined, and summarised in greater detail, by Trendall (2002), who points out that although some features of the Cloud (1973) model are incompatible with improved knowledge of the time-distribution of IF, his belief in an organic role in IF formation, and the need to interpret IF deposition holistically, taking the chemical evolution of both the atmosphere and the oceans (see section 5.2, for extensive discussion of these topics), and reconciling these with the parallel evolution of life, give Preston Cloud's idea a unique place in IF study. Anoxygenic bacterial photosynthesis has also been proposed (e.g., Kump, 1993; Widdell et al., 1993) as a mechanism for IF precipitation, and seems equally consistent with the isotopic evidence and indirect evidence for an organic agency in B IF deposition. An abiogenic precipitation mechanism based on photo-oxidation of iron by sunlight has been suggested (e.g., Braterman et al., 1984; Francois, 1986; Braterman and Cairns-Smith, 1987a, b); Draganic et al. (1991) have proposed the decomposition of early ocean water by potassium-40 radiation as a source of oxygen for the precipitation of some early Precambrian BIFs. More generally, Klein and Beukes (1989) have described a depositional model for microbanded BIF in which the depositional mechanism has no direct biological control, but is reliant on the annual overturn of basin water (ibid., p. 1771), with the iron supplied by upwelling from a stratified ocean. Klein and Beukes (1989, p. 1768) contend that "the lack of organic carbon in iron-formation is a serious problem in any model that couples the deposition of BIF to microbial activity, especially photosynthetic activity... ". Trendall (2002) has discussed their reasons for this, and has noted, for the Hamersley basin at least, some of the evidence for the presence of abundant photosynthesising cyanobacteria immediately before, and during earliest, B IF deposition. Of particular interest is the recent discovery by Brocks et al. (1999) of hydrocarbon markers from shale within the Marra Mamba Iron Formation, which they interpret as firm evidence for the presence of photosynthesising cyanobacteria (discussion on the latter in section 5.2). Although more evidence is needed before the precise role of microbiota in the precipitation of iron during BIF deposition can be established, it is worth noting finally that Konhauser et al. (2002), have shown that the concentrations of P and trace metals (V, Mn, Co, Zn and Mo) within Dales Gorge Member BIF, as well as the rate of iron deposition, are consistent with precipitation by iron-oxidising bacteria like those found in modern Fe-rich aqueous environments.

A second-order genetic model The following integrated depositional model links various topics mentioned in isolation under earlier headings; it is called "second-order" because it builds on the "first-order" case, already argued, that IF is a lithified chemical precipitate whose primary composition was close to that of the present IF.

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Trendall and Blockley (1970)envisaged a depositional model for BIFs of the Hamersley basin in which the precursor iron- and silica-rich sediments were laid down annually in a gently sinking basin of c. 100,000 km 2, offshore from a continental area with an arid climate, and with a water depth of 50-250 m. They suggested that the Fe concentration of the basin water was 10-20 ppm, and that this level was maintained by fumarolic volcanicity along the western edge, where there was likely to be a shallow connection with the open ocean. Many features of this early model remain valid, although ocean water as a source of iron and silica, rather than a local volcanic source, seems to account more credibly for the strikingly uniform character of BIF over a long period of Precambrian time. The isotopic and geochemical evidence that the iron is of volcanic origin is more simply accounted for by its primary derivation from global seafloor volcanic rocks. The long-term retention of iron in solution in ocean water assures a stable annual supply and also provides a buffer against large concentration fluctuations. The Hamersley basin during BIF deposition was probably a deep (> 500 m) offshore shelf, not significantly barred or restricted, and with access to circulating ocean water containing no more than about 10 ppm dissolved ferrous iron. A suggestion that the "deeper waters of the early oceans" were richer in iron than surface water was made by Holland (1973), who saw the upwelling of deep water as the source of iron for BIF deposition in "shallow marine areas". In a later discussion Holland (1984) drew attention to the low iron content of early Precambrian shallow-water carbonates to support the restriction of iron-rich water to the deep oceans. The significance of a stratified ocean in relation to the deposition of the Transvaal Supergroup was discussed in detail by Klein and Beukes (1989), and this concept is now widely accepted, for example, Eriksson et al. (1997, p. 49) comment that "Archaean sea-water also was enriched in Fe ++ but only below the pycnocline" (see also discussion in section 5.2). We now accept the application of this iron-stratified ocean model for the Hamersley basin, and its implication that the supply of iron for BIF deposition was either related to upwelling, or that the pycnocline was close to the level of the basin floor. Iron and silica were precipitated together annually from the basin water, the former probably by photosynthesising organisms, and the latter by uncertain means. The inclusion in the model of an ocean stratified in both Eh and iron content make it difficult to use the presence or absence of IFs at any point in the Precambrian stratigraphic record as an index of atmospheric content at that time (see also discussion by Ohmoto in section 5.2). In principle, IFs could have been deposited when oxygen was either virtually absent or relatively abundant in the atmosphere, since the lower, anoxic, and iron-bearing deep waters would have been protected by the upper levels from direct involvement in the processes of IF deposition. The simplest explanation for the mid-Precambrian cessation of major IF deposition (Fig. 5.4-2) is the gradual lowering of the pycnocline through a steady increase in ocean oxygenation as living organisms, and photosynthesis, became more abundant and effective. This revised model for the Hamersley basin, which is in close accord with the general model for IF deposition outlined by Button (1982), is one that may be applied generally to a wide variety of depositories containing IE This point was emphasised by Trendall (2002)

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in association with the concept of "necessary and sufficient" conditions for IF deposition: there is no need for models which closely constrain basin architecture, when all that is required is a geometry which permits the annual circulation of ocean water from adequate depth. From the earliest Archaean basins onwards, IF could be formed in a variety of basinal settings. Trendall (2002) linked the diversity of basinal settings with a preferred hypothesis for the origin and development of continents; the further implications of that hypothesis (see section 3.6 for additional discussion) are outside the mainstream of current thinking, and await further testing.

The Neoproterozoic Iron Formations The Neoproterozoic IFs are only referred to briefly here, for two reasons. Firstly because the history of IF deposition before the c. 1 Ga late Precambrian hiatus (Fig. 5.4-2) is an issue that needs to be independently understood; and secondly because the late Neoproterozoic IFs seem to be part of an exceptional "superevent" whose exact nature and significance is the subject of intense current scrutiny (see sections 3.10, 3. l 1 and 5.6-5.8). Such IFs have been described from Australia (Braemar Iron-Formation: Whitten, 1970; Holowilena Iron-Formation: Dalgarno and Johnson, 1965), northwest Canada (Rapitan Group: Young, 1976; Klein and Beukes, 1993), Brazil (Jacadigo Formation: Dorr 1973; Urucum: Walde et al., 1981), and Namibia (Damara Supergroup: Martin, 1965); other occurrences are listed by Yeo (1986). They have ages in the approximate range 800-600 Ma, and most have some evidence of glacial association (sections 5.6-5.8). The Rapitan IF is one of the most closely studied, and may be taken as representative. Like other Neoproterozoic IFs it is essentially a fine-grained, thin-bedded, hematite-quartz rock lacking either the sharply defined mesobanding of early Precambrian BIF or the coarser banding of GIE Its sedimentology was described by Young (1976), who also provided a review of its possible significance, including the proposal of Williams (1972) based on the secular variation in the inclination of the rotational axis of the Earth (see section 5.9); Young (1976; section 5.6) concluded that there was no definitive explanation for the presence of these Neoproterozoic IFs. The later detailed study by Klein and Beukes (1993) showed that the simple mineralogy and major element chemistry are distinctly different from those of most early Precambrian IFs. They concluded that the Rapitan IF was deposited during a major transgression after a glacial event, and drew attention to the suggestion of Kirschvink (1992) that it could be related to development of anoxic, and iron-rich, ocean bottom water consequent upon an almost completely ice-covered ("snowball") Earth (see discussions in section 5.2, and particularly by Young, section 5.6, and by Williams, section 5.7). Hoffman et al. (1998b) have subsequently developed Kirschvink's (1992) idea of a snowball Earth, which was based on an earlier idea of Harland (1964), into an integrated hypothesis in which the IFs are only one geological outcome of rapid and extreme oscillations of global climate. During cold excursions ice covered much of the land, and the oceans were also insulated from the atmosphere by thick ice. There was a consequent fall in dissolved oxygen level, resulting in an increase in dissolved iron. Meanwhile, continuing volcanism built up atmospheric carbon dioxide levels in the atmosphere, to a point

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where greenhouse warming led to rapid melting of both terrestrial and oceanic ice. The resultant melt-related deposition of glaciogene sediments at the start of the warm period was coeval with precipitation of dissolved ferrous iron from the re-oxygenated oceans, leading to a global association of IF with glacial material. The hypothesis provides an elegant working model: Schmidt and Williams (1995), for example, have demonstrated that the glaciogene deposits of South Australia were formed in association with grounded glaciers near sea level at near-equatorial latitude, although they have pointed out that explanations other than greenhouse oscillation need to be considered. The snowball Earth hypothesis is at an early stage of testing, and the emphasis placed by some authors (e.g., Breitkopf, 1988; Young, 1988; Trompette et al., 1998) on the relationship between rift-related mafic volcanism and some Neoproterozoic IFs indicates that the evidence for a purely climatic control of their deposition is not yet definitive (for contrasting views on the snowball Earth hypothesis, see sections 5.6-5.8).

Concluding Comments Although understanding of IF has increased substantially in the last few decades, many problems concerning it still need to be resolved. Examples of these include the status of silica vis-a-vis the iron that is the natural and historical focus of attention, the exact role (indeed, if any) of organisms in the depositional process, and the more precise determination of the depositional ages of individual IFs. A fruitful field of future research is the conduct of controlled laboratory experiments where postulated conditions for both deposition and diagenesis of IF can be reproduced.

5.5.

THE PRECAMBRIAN SULPHUR ISOTOPE RECORD OF EVOLVING ATMOSPHERIC OXYGEN

T.W. LYONS, L.C. KAH AND A.M. GELLATLY

Introduction This section will focus on the most recent developments in Precambrian sulphur geochemistry. Sulphur, along with carbon (section 5.3), has long provided the most effective isotopic tool available for deconstructing biospheric evolution (see chapter 6). Such value is even more apparent today, as increasing sophistication in sulphur-based palaeoenvironmental reconstruction mirrors major advances in instrumentation, analytical approaches, microbiology, and chemical oceanography. While many questions and debates persist, sulphate and sulphide S isotope ratios in marine sediments show great promise in faithfully tracking even the most subtle details of oxygen's role in the evolving Precambrian oceanatmosphere system and the coupled pathways of bacterial respiration (see also detailed discussions by Ohmoto, section 5.2). The Precambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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In this review paper we first summarise the background necessary for interpreting sulphur isotope trends in Precambrian sediments. This background leads naturally into interpretations of Archaean S cycling and its temporal relationships to the earliest signs of life. We then highlight recent methods and models for analysing Proterozoic S isotopes, including case studies based largely on iron sulphides in mineralised regions and sulphate bound within gypsum and carbonate minerals. A discussion of the global implications of the Proterozoic records emphasises ongoing efforts to estimate sulphate concentrations in sea water and oxygen-versus-sulphide availability in the deep ocean. Concentrations of sulphate in the Proterozoic ocean--like oxygen--were intermediate relative to those of the Archaean and Phanerozoic (see, however, section 5.2). This intermediate ocean chemistry may have favoured a globally anoxic and sulphidic (euxinic) deep ocean. Proterozoic S isotope records of this transition provide a clear environmental context for the diversification of eukaryotes, culminating in the rise of metazoans. The challenge in reviewing this topic is to find and synthesise the common threads, the consistencies and inconsistencies, and the controversies encapsulated in a complex and rapidly evolving field and, in doing so, portray the state of the art.

BackgroundDSulphur Isotope Geochemistry Interpretations of Precambrian sulphur isotope trends require an understanding of how sulphur in the ocean is cycled and ultimately sequestered in the geologic record, including the burial of reduced S as iron sulphide. Sedimentary pyrite formation begins with bacterial reduction of sulphate under conditions of anoxia in the water column or within sediment pore fluids. The kinetic isotope effect associated with bacterial sulphate reduction (BSR) results in hydrogen sulphide (and ultimately pyrite) that is depleted in 34S relative to the 348/328 ratios of residual, coexisting sulphate (Goldhaber and Kaplan, 1974). On geologic time scales, the balance between net burial versus oxidative weathering of pyrite controls the 348[328 ratio in the global oceanic sulphate reservoir and, along with the redox cycling of organic carbon, is the principal modulator of pO2 in the atmosphere (Claypool et al., 1980; Berner and Petsch, 1998) (sections 5.2 and 5.3). Continental oxidation of pyrite and other metal sulphides, a major contributor of sulphate to the ocean throughout the latter half of Earth history (Berner and Berner, 1996), facilitated the buildup of oceanic sulphate during the Proterozoic under an increasingly oxidising atmosphere (Canfield, 1998). Oxidation of continental igneous/magmatic sulphides would have dominated initially the sulphate flux prior to deposition and recycling of sulphate- and sulphide-rich sedimentary rocks later in the Proterozoic and Phanerozoic. It is recognised widely that pyrite burial and weathering largely control the concentrations and isotopic compositions of oceanic sulphate, but it has been more challenging to reconcile experimental results for BSR with the complex range of 634S values observed in sedimentary pyrite (see also section 5.2). Dissimilatory sulphate reduction under pureculture laboratory conditions can produce sulphide depleted in 34S by roughly 2-46%0 relative to the parent sulphate (Chambers et al., 1975; Canfield, 2001; Detmers et al., 2001). Although this range is generally accepted, the controls on the magnitude of this fractionation

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are subjects of renewed study and debate. For example, contrary to a long-held assumption, some workers are now suggesting that the isotopic offset between parent sulphate and product HS- during BSR (A34S) does not vary with a simple inverse relationship to the rate of sulphate reduction (cf. Kaplan and Rittenberg, 1964; Canfield, 2001; Detmers et al., 2001; Habicht and Canfield, 2001). Nevertheless, isotope fractionations during BSR appear to be unaffected by sulphate concentration at levels > 1-2 mM (relative to c. 29 mM in the modem ocean) (Canfield, 2001). In light of the significantly smaller isotope effects attributable to BSR under pure culture conditions, and assuming that the experiments mimic nature (cf. Wortmann et al., 2001; Ohmoto, section 5.2), recent studies have addressed the fractionations of up to and exceeding 60%o (e.g., Lyons, 1997) that abound in the Phanerozoic. One model with important Precambrian implications (Canfield and Thamdrup, 1994; Habicht and Canfield, 2001) invokes bacterial disproportionation of elemental sulphur and other S intermediates as a means of exacerbating the 3as depletions observed in HS- and pyrite (Fig. 5.5-1). Questions remain, however, as to why such redox cycling (with disproportionation) gives rise to the commonly observed A348 "ceiling" of c. 60-70%0 and how prevalent these pathways are in the subsurface, where sulphide concentrations can be high enough to be toxic to the disproportionating bacteria and where the availability of S intermediates is generally low. Ultimately, net isotopic fractionations preserved in geologic systems reflect both the magnitudes of bacterial fractionations and the properties of the sulphate reservoir as recorded in the integrated history of pyrite formation (Zaback et al., 1993). Even in the presence of large fractionations during BSR and coupled disproportionation, high 6348sulphide values occur in (1) pore-water systems with restricted renewal of sulphate relative to the rate of bacterial consumption (i.e., under conditions of rapid sediment accumulation) or (2) through deposition in a sulphate-limited basin or ocean. Conversely, low 634S values typically represent marine systems where sulphate availability does not limit BSR. As a result of these multiple controlling factors, bacteriogenic pyrite can display a broad range of 634S values that are often very low (34S-depleted) relative to coeval sulphate. These broad ranges and 34S depletions are the oft-cited fingerprints of BSR, although strongly positive values--particularly common in the Proterozoic---characterise conditions of sulphate limitation. The possibility of abiotic sulphate reduction at elevated temperatures must also be considered in Precambrian studies (Machel et al., 1995). Independent evidence for hydrothermal mineralisation (e.g., fluid-inclusion temperature estimates and pyrite textural relationships) is invaluable in distinguishing thermochemical pathways from BSR.

In the Beginning--Sulphur Cycling in the Archaean Until very recently most geochemical arguments implied a temporal correlation between the first hints of oxygenic photosynthesis in the Archaean and the primitive origins of bacterial sulphate reduction. In a review paper published in 1983, addressing the antiquity of BSR, Schidlowski et al. noted: "Between 1.8 and 3.8 Ga, the sulfate record is virtu-

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S042- - .....................................

8349

i \ i BSR

~,.. | | |

SQ2S O 4 2-

| l

IA2 ......

H2S~ S ~

H2S"-I"S ~

al = BSR (2CH20 + S042- ~ " H2S + 2HC03-), <_40-45%(based on laboratory results) A2 up to and exceeding 60%o a small for H2S --> S~ a small for H2S --~ "FeS" ~

| I I I l I I

H2s . . t .

FeS2

Fig. 5.5-1. Schematic (qualitative) representation of S isotope fractionations resulting from bacterial sulphate reduction (AI) and elemental S disproportionation. Note the increasing depletion in 34S in the H2S reservoir with progressive oxidation-disproportionation steps, which can result in overall fractionations between sulphate and sulphide (a2) that exceed 60%o. The transformation of H2S to SO reflects partial oxidation. Isotope effects associated with this oxidation and with the formation of iron sulphide from H2S are minor compared to those associated with BSR and disproportionation. This figure is modified slightly from Canfield and Thamdrup (1994). ally nonexistent, but sedimentary sulfides often display what might be accepted as bacteriogenic patterns. The 634S distributions displayed by the 2.7-Ga-old Michipicoten and Woman River Iron Formations have come to be widely accepted as the oldest presumptive evidence of bacterial sulfate reduction". As recently as 1999, Canfield and Raiswell were still suggesting that the first S isotope evidence for BSR is found in sedimentary rocks dated at c. 2.7 Ga. As outlined in the previous section, large spreads in sulphide (dominantly pyrite) S isotope data, abundant negative values, and large isotopic offsets between coeval sulphate and sulphide are the generally accepted isotopic signatures of BSR. The paucity of these bio-indicators in the early record were (and largely still are) attributed to the absence of sulphate in a dominantly oxygen-deficient Archaean ocean-atmosphere system (see, however, Ohmoto, section 5.2). In the absence of sulphate, BSR may have evolved much later, in concert with increasing seawater sulphate concentrations. Contrary opinions have been advanced, however, and the Archaean records have received renewed and refined attention over the past few years (see Canfield and Raiswell, 1999; Canfield et al., 2000, for comprehensive surveys) (see also discussions in section 5.2). The oldest evidence for cyanobacteria and oxygenic photosynthesis is lipid biomarker data from 2.7 Ga shales of the Pilbara craton, Western Australia (Brocks et al., 1999; additional evidence reviewed in Canfield and Raiswell, 1999; Summons et al., 1999; cf. Blank, 2002) (see also section 6.2). Nevertheless, new data from the Warrawoona Group, also

5.5. Precambrian Sulphur Isotope Record

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part of the Pilbara craton, show S isotope fractionations of up to 21.1%0 (mean of 11.6%0) for coexisting c. 3.47 Ga barite and pyrite inclusions (Shen et al., 2001; related molecular phylogenetic arguments are further developed in Canfield and Raiswell, 1999), suggesting BSR far earlier than the first known record of appreciable oxygen production (see also Ohmoto, section 5.2). (Morphological microfossil evidence in the Warrawoona sediments (Schopf, 2000; see also section 6.2) does not point uniquely to cyanobacterial, O2-producing photosynthesis.) Oxygenic photosynthesis is critical because it drives sulphate delivery to the ocean through oxidative weathering on the continents, thus supporting BSR in the anaerobic portions of the sediment and water column. In the possible absence of oxygenic photosynthesis, photochemical oxidation of volcanogenic SO2 and anoxygenic photosynthesis may have been the dominant sources of sulphate to the ocean for early BSR (Canfield and Raiswell, 1999; Farquhar et al., 2000a; Habicht et al., 2002). Canfield and Raiswell (1999) argued that the early rise of anoxygenic photosynthesis is suggested by bacterial molecular phylogeny. Also, 13C depletions in the earliest known sedimentary rocks (c. 3.7-3.8 Ga; Schidlowski, 1988, 2001; Mojzsis et al., 1996; Rosing et al., 1996; Rosing, 1999; Mojzsis and Harrison, 2000; cf. Fedo and Whitehouse, 2002; van Zuilen et al., 2002) and in the kerogen residues of what appear to be the earliest known cellular fossils of any type (c. 3.47 Ga Apex chert, Warrawoona Group; Schopf, 1993; Kazmierczak and Kremer, 2002; Schopf et al., 2002; cf. Brasier et al., 2002; Pasteris et al., 2002) (see also sections 6.2 and 6.4) may reflect the anoxygenic photoautotrophic pathway. As an analogy, many modem purple and green sulphur bacteria store elemental sulphur generated during anoxygenic photosynthesis, which can later serve as a photosynthetic electron donor during further oxidation to sulphate. Any sources of sulphate that were decoupled from O2-producing photosynthesis, however, would have been quantitatively minor--perhaps yielding only local "oases" of elevated marine sulphate in isolated, evaporative ponds (e.g., Shen et al., 2001; Pavlov and Kasting, 2002). The above interpretation of the c. 3.47 Ga barite/pyrite record is controversial and is predicated on the assumption that original sedimentary gypsum was later replaced pseudomorphically by barite (Buick and Dunlop, 1990). The gypsum argument precludes thermochemical pathways of sulphate reduction (Machel et al., 1995) in the low temperature setting of original sulphate deposition and pyrite formation. Runnegar et al. (2002) challenged the gypsum precursor for the Warrawoona Group barite described by Shen et al. (2001) in favour of direct hydrothermal precipitation of barite. The isotopic offset between the co-existing sulphate and sulphide was thus interpreted as equilibrium fractionation under hydrothermal conditions rather than kinetic isotope effects during BSR. At the same time Runnegar et al. (2002) invoked recently discovered pathways of mass-independent fractionation unrelated to BSR (Farquhar et al., 2000a, b) to explain pyrite S isotope records in Archaean black cherts and shales, and in banded iron-formations. Despite existing debates over the antiquity of bacterial sulphate reduction and its relationship to oxygenic photosynthesis (e.g., discussion in section 5.2), there is a widely held view that the Archaean ocean contained very low concentrations of sulphate that mirrored the lack of oxidative weathering beneath a reducing atmosphere. Under such conditions,

426

Chapter 5: Evolution of the Hydrosphere and Atmosphere

sulphate became enriched only locally, and the isotopic difference between contemporaneous sulphate and sulphide resulting from BSR would otherwise have been suppressed even if the BSR pathway had evolved. Because of sulphate limitation, fractionations of < 4%0 result from BSR under sulphate concentrations of < 1 mM (Harrison and Thode, 1958; Canfield and Raiswell, 1999; Canfield et al., 2000; Canfield, 2001; cf. Shen et al., 2002). (Most recently, Habicht et al. (2002) proposed sulphate concentrations of < 200 ~tM as the critical threshold in limiting S isotope fractionation during BSR to a few per mil or lessmthus further constraining Archaean oceanic sulphate concentrations to extremely low values.) Consequently, BSR may have occurred throughout the Archaean without the emergence of the telltale isotopic signatures well expressed by c. 2.3 Ga (see Proterozoic discussion below). Older isotopic signals of BSRmsuch as those at approximately 3.47 and 2.7 Ga (see also section 5.2)~might record locally elevated sulphate concentrations. Undoubtedly, the photosynthetic production of 02 at c. 2.7 Ga would have enhanced sulphate accumulation, although global oceanic levels appear to have remained low until c. 2.3 Ga.

Alternative interpretations of the Archaean sulphur record In contrast to the low-sulphate Archaean ocean model, the absence of large fractionations between contemporaneous sulphate and sulphide and the predominance of data clustering around 0%0 have been cited historically as evidence against the presence of BSR and in favour of S mineralisation controlled by magmatic processes. Such abiotic mantle fluxes can yield metal sulphides (and sulphate minerals) that range typically over only a few per mil either side of 0%0 (Ohmoto and Goldhaber, 1997). Distinguishing between the low-sulphate (bacterial) and high-temperature (mantle-buffered) models necessitates careful, independent evaluations of the modes and temperatures of mineralisation--keeping in mind that large abiotic fractionations (e.g., 20%0 at 100~ are also possible during thermochemical sulphate reduction and are subject to the same reservoir considerations (e.g., sulphate limitations) as BSR (Machel et al., 1995). Archaean sediments almost certainly bear the signatures of all these processes. The Archaean S isotope record is complicated further by mass-independent fractionations recorded in sedimentary sulphate and sulphide minerals. These fractionations prevailed under the low pO2 conditions of the Archaean (Farquhar et al., 2000a, 2002; Mojzsis et al., 2002, 2003; Pavlov and Kasting, 2002; Runnegar et al., 2002; Farquhar and Wing, 2003; cf., however, Ohmoto et al., 2001; Ohmoto, section 5.2). More specifically, widespread mass-independent fractionation stems from a sulphur cycle dominated by UVdependent, gas-phase atmospheric reactions--such as photolysis of volcanogenic SO2 and HzS. These signals are best generated and preserved in the absence of (1) appreciable atmospheric ozone and (2) the large-scale, homogenising, mass-dependent effects of (bacterial) S processing within a sulphate-rich marine reservoir beneath an oxidising atmosphere (Farquhar et al., 2000a, b, 2001; Pavlov and Kasting, 2002). The loss of preserved massindependent behaviour is expressed across a transition spanning from c. 2.45 to 2.09 Ga (Farquhar et al., 2000a). Although much is still unknown about these mass-independent fractionations and their capacity to mask and mimic the mass-dependent isotope effects of

5.5. Precambrian Sulphur Isotope Record

427

possible BSR in early Precambrian sediments (Runnegar et al., 2002), this timeframe for the loss of mass-independent effects is generally consistent with multiple proxy evidence for a fundamental shift in the level of atmospheric O2 (see details below). A controversial alternative model for the Archaean sulphate/sulphide S isotope record argues for an Oz-rich atmosphere/ocean and correspondingly high concentrations of sulphate in Archaean seawater (i.e., more than one third the present level by 2.5 Ga and "appreciable" sulphate by 3.4 to 3.2 Ga) (Ohmoto et al., 1993; Kakegawa et al., 1998; Kakegawa and Ohmoto, 1999; discussions in Ohmoto, section 5.2). This interpretation is based on microscale 6348pyrite variation in sediments and metasediments dating from more than or equal to c. 2.5 Ga. Although the isotopic data span up to c. 14%0, many of the data cluster around 0%0 and show inter- and intra-sample variation smaller than 14%0--not unlike the ranges possible within complex hydrothermal systems. Assessments of fractionations between contemporaneous Archaean sulphate and sulphide are almost always compromised by poor control on the marine 6348sulphate. Nevertheless, fractionations were estimated at less than or equal to c. 20%o for c. 2.5 Ga pyrite (Kakegawa et al., 1998) and only c. 2-5%0 for 3.4-3.2 Ga pyrites (Ohmoto et al., 1993; Kakegawa and Ohmoto, 1999). Ohmoto et al. (1993) and Kakegawa et al. (1998) suggested that the comparatively small inferred fractionations, relative to values up to and exceeding 60%0 for the Phanerozoic, may reflect high rates of BSR in a warm Archaean ocean (section 5.2). As discussed above, the relationship between sulphate reduction rates and isotope effects is more ambiguous than previously supposed. Most importantly, however, S isotope fractionations are still in the range of 20-40%o in modern settings characterised by high concentrations of sulphate, warm temperatures, and extremely high natural rates of BSR (Habicht and Canfield, 1996; Canfield et al., 2000). It is difficult therefore to envision BSR rate-control for the inferred small fractionations in an Archaean ocean assumed to be rich in sulphate. Furthermore, the model of Shen et al. (2001), based on localised sulphate enrichment in evaporated Archaean sea water, argues for early BSR with fractionations as high as c. 20%0, despite an Archaean ocean that otherwise contained very low levels of sulphate. Such a model for locally elevated sulphate could apply to a temporally and spatially broad range of Archaean deposits (e.g., Ohmoto et al., 1993; Kakegawa et al., 1998; Kakegawa and Ohmoto, 1999).

A Time of Transition--Sulphur Cycling in the Proterozoic Introduction Although the differences in ~34S values for sedimentary sulphides and inferred seawater sulphate indicate BSR tentatively by 3.47 Ga and convincingly by c. 2.7 Ga, the fractionations are generally ~< (frequently <<) 15-20%0 prior to c. 2.3 Ga (e.g., Ohmoto et al., 1993; Kakegawa et al., 1998; Canfield and Raiswell, 1999; Kakegawa and Ohmoto, 1999; Shen et al., 2001) (section 5.2). These early fractionations are appreciably smaller than the A34S values of up to c. 60%o commonly observed in Phanerozoic sedimentary systems-which are thought to record additional fractionation during disproportionation of interme-

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

diate S species (Canfield and Thamdrup, 1994)--and the approximately 45%o achieved via BSR alone in laboratory experiments where sulphate is non-limiting. Canfield and Teske (1996; as further developed in Canfield, 1998; Canfield and Raiswell, 1999; Canfield et al., 2000; Shen et al., 2001) argued for a two-step evolution in the Proterozoic sulphur isotope record. According to this model, small isotopic fractionations prior to c. 2.3 Ga were constrained by marine sulphate concentrations of < 1 mM (< 200 I.tM; Habicht et al., 2002). Sulphate enrichment occurred in isolated, evaporative settings, yielding only local S isotope fractionations that are typical of BSR not limited by sulphate availability (e.g., Shen et al., 2001). Between c. 2.3 Ga and 0.64-1.05 Ga (or 0.8 Ga; Canfield, 1996), the upper limit for fractionation was fixed at 40 to 45%o via BSR in an ocean where the critical threshold in sulphate availability (i.e., concentrations more than c. 0.2-1 mM) was exceeded globally because of rising atmospheric oxygen. Atmospheric 02 concentrations are thought to have increased dramatically at c. 2.3 Ga (e.g., Pavlov and Kasting, 2002) (see, however, discussion in Ohmoto, section 5.2, and in Lindsay and Brasier, section 5.3). Empirical data from this interval, however, commonly record maximum net 345 depletions of c. 20-25%o relative to coeval sulphate (e.g., Kah et al., 2001; Luepke and Lyons, 2001) despite the potential for larger instantaneous fractionations during BSR. As discussed below, these net fractionations are influenced also by the local and global availability of sulphate. By c. 0.8 Ga oxygen availability in the ocean-atmosphere system may have increased to the point of greatly enhancing the oxidative portion of the sulphur cycle, including evolutionary development of non-photosynthetic sulphide-oxidising bacteria. Under such oxidising conditions, and the associated biotic and abiotic production of intermediate sulphur species, disproportionation reactions could have increased fractionation to values typical under open reservoir conditions throughout the Phanerozoic (e.g., Lyons, 1997). Canfield and Teske (1996; also Canfield and Raiswell, 1999) used controversial molecular phylogenetic approaches to argue for evolutionary radiation of non-photosynthetic sulphideoxidising bacteria during the Neoproterozoic, which they attributed to rising atmospheric oxygen concentrations, leading ultimately to the appearance and rapid radiation of metazoans. Increases in atmospheric oxygen--resulting first in increasing sulphate in the ocean followed by enhanced redox cycling of Suare a product of large scale burial of organic carbon (i.e., net photosynthesis on a global scale) (Lindsay and Brasier, section 5.3; Ohmoto, section 5.2). Temporal trends in the growth of the crustal inventory of reduced carbon (versus loss via weathering) are reflected in the 313C chemostratigraphy of sedimentary carbonates and organic matter (Kump and Arthur, 1999, among others) (section 5.3). Using these isotopic proxies, Des Marais et al. (1992) argued for "stepwise oxidation" of the Proterozoic atmosphere, with oxygenation steps corresponding generally with the c. 2.3 Ga and c. 0.8 Ga transitions in the sulphur isotope record noted above (Canfield and Teske, 1996; Canfield, 1998; Canfield and Raiswell, 1999) (see, however, arguments in section 5.2). Models for S cycling throughout the Proterozoic, while elegant and inclusive, generally suffer from our weak understanding of the concentrations and 8345 values of sulphate

5.5. Precambrian Sulphur Isotope Record

429

in the ancient ocean. Such deficiencies are even more pronounced in studies of Archaean palaeoenvironments. For example, fractionations between contemporaneous sulphate and sulphide are often only approximated tenuously given the paucity of Archaean and Proterozoic sulphate isotope data--many of which derive from barite. Hydrothermal processes can obscure barite records of original seawater chemistry (cf. Strauss, 1997 and Paytan et al., 1998). Furthermore, sedimentary (evaporative rather than hydrothermal)origins are often uncertain for early anhydrite/gypsum. Finally, it is difficult to establish temporal equivalence for Proterozoic geochemical data given the discontinuous records and limited age controlmall of which make Proterozoic global ocean chemistry elusive. The remainder of this section will emphasise our recent work on carbonates, gypsum, and fine-grained siliciclastics from the Mesoproterozoic (1.6-1.0 Ga), particularly as related to our efforts to constrain the $345 and availability of sulphate in the early ocean. Several case studies highlight where we are, and where we still need to go, with Precambrian S geochemistry. Isotope recordsBsulphide case studies Our comprehensive sulphur isotope data set for Mesoproterozoic sulphides from the northwestern U.S.A. helps constrain many of the issues discussed in the previous section, with added complexity stemming from both local and global oceanic controls. Specifically, Lyons et al. (2000) and Luepke and Lyons (2001) presented sulphur isotope results for pyrite from shales (Newland Formation) of the c. 1.47 Ga lower Belt Supergroup, Montana. These organic-rich shales accumulated in comparatively deep water within the Helena embayment, a fault-bounded eastern extension of the Belt basin, and were the hosts for multiple episodes of stratabound and commonly stratiform massive sulphide, sedimentary exhalative (SEDEX) mineralisation. Despite clear evidence for hydrothermal activity, host shales lying stratigraphically above and below the mineralised zones are dominated by low-temperature, bacteriogenic pyrite, as indicated by the presence of framboidal pyrite textures, diagnostic S isotope trends, and other criteria. The ~345 data from the shales show two clear relationships: (1) a preponderance of positive values ( - 8 . 7 to +36.3%o; mean = +7.6%0, n = 41), including overlap with the $345 of coeval sulphate as constrained by barite and carbonate-associated sulphate (see Strauss and Schieber, 1990; Gellatly and Lyons, submitted, for sulphate data), and (2) stratigraphic S isotope trends wherein $345 values decrease systematically up section by up to 45~00 over 10s to 100s of metres. Further to the west, in the "main" Belt basin, data for pyrite and pyrrhotite in the Prichard Formation--argillites roughly coeval to the Newland shalesmshow similar S isotope ranges ( - 0 . 8 to -I-22.8%0; mean = +7.6%0, n = 42) and similar stratigraphic variation (c. 20%o) over 100s of metres (Fig. 5.5-2). In contrast to an earlier lacustrine model (Winston and Link, 1993), Lyons et al. (2000) used the 34S-enriched and stratigraphically variable isotope data from the Prichard and Newland Formations to argue for at least episodic fluxes of open marine water into a restricted early Belt basin, wherein the strength of the marine connection varied over time. While this interpretation is consistent with the rift basin setting independently indicated for Belt deposition, Lyons et al. (2000) and Luepke and Lyons (2001) also suggested that Belt S isotope trends should bear the effects of smaller Mesoproterozoic fractionations in

Chapter 5: Evolution of the Hydrosphere and Atmosphere

430

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Fig. 5.5-2. Stratigraphic distributions of S isotope data for the c. 1.47 Ga Prichard and Newland Formations of the Belt Supergroup, northwestern U.S.A. Data for the Prichard represent fine-scale pyrite/pyrrhotite laminae, lenses, and disseminations within dark grey to black argillites. Despite low-grade metamorphism in the Prichard, the ~345 values are assumed to preserve original bacterial fractionations. The Newland data are from disseminated (in part framboidal) bacteriogenic pyrite within unmetamorphosed black shales. Also provided for comparison are the $34S ranges for Newland sulphate as determined from barite and carbonate-associated sulphate. Details are provided in Lyons et al. (2000) and Luepke and Lyons (2001).

the absence of disproportionation pathways (Canfield and Teske, 1996) and substantially lower global levels of marine dissolved sulphate relative to the Phanerozoic ocean. A number of authors have noted that enrichments in 345 are a common characteristic of sedimentary iron sulphides from the Proterozoic (e.g., Logan et al., 1995; Canfield, 1998; Canfield and Raiswell, 1999; Gorjan et al., 2000; Shen et al., 2002; Strauss, 2002), including large stratigraphic variations within a given sedimentary unit (Ross et al., 1995; Strauss, 1997, 2002; Lyons et al., 2000; Luepke and Lyons, 2001). In our ongoing survey of Proterozoic SEDEX deposits, we are seeing similar 345 enrichments and stratigraphic trends (Lyons and Gellatly, 2002) (Fig. 5.5-3). SEDEX mineral deposits, which form on or near the seafloor penecontemporaneously with host sediment deposition, result from submarine venting of hydrothermal fluids (Lydon, 1996). Traditionally, workers have tied high ~345 values within SEDEX metal sulphide mineralisation to bacterial and thermochemical reduction of seawater sulphate in evolving reservoirs isolated from the ocean via local re-

5.5. Precambrian Sulphur Isotope Record

431

striction within fault-bounded (rift) basins (see Lydon, 1996, for a review of sedimentary exhalative base metal sulphide deposits). It is our assertion, however, that 345 enrichments in the Proterozoic deposits are exacerbated by low sulphate concentrations in the Proterozoic ocean, and the stratigraphic trends may track rapid S345sulphatevariation in the global ocean rather than in local reservoirs. (Furthermore, the paucity of Archaean SEDEX mineralisation (Turner, 1992) could reflect very low sulphate concentrations in the early ocean). Isotope records--sulphate case studies

Previously untapped sulphate records from Proterozoic carbonates further suggest that sulphate concentrations and $345 in the global ocean fluctuated rapidly, but concentrations generally remained low. In their seminal S isotope compilation for marine sulphate over geologic time, Claypool et al. (1980) relied on evaporite (gypsum/anhydrite) data, which spanned the Phanerozoic but sparsely covered only the latest Precambrian. This scarcity of Precambrian data, which persists in more recent compilations (Strauss, 1993, 1999), reflects poor preservation potential of gypsum under surficial weathering conditions and the difficulty of achieving gypsum saturation through evaporation of low-sulphate Precambrian seawater (Kah et al., 2001, and discussions below). In the absence of gypsum data, particularly for the Archaean, ~345 compilations for marine sulphate have emphasised bedded barite (e.g., Schidlowski et al., 1983; Canfield, 1998). Evaluations of fractionations over time (Canfield, 1998; Canfield and Raiswell, 1999) rely on comparing these barite data to approximately contemporaneous ~345pyrite but are limited by the patchiness of the barite record and its varying capacity to record primary sea water chemistry. Because of the lack of gypsum data, carbonate-associated sulphate (CAS) has emerged as an attractive alternative. Although the presence of sulphate in carbonate minerals (calcite, aragonite, and dolomite) has long been known, is was not until 1989 that Burdett et al. demonstrated that CAS can record isotopic values equivalent to those of contemporaneous modern seawater and ancient evaporite deposits, thus confirming that CAS can approximate closely ancient seawater (see also Strauss, 1999; Lyons et al., 2002). Sulphate is present typically in modern and ancient sedimentary carbonate minerals in concentrations ranging from a few hundred ppm to extremes of 104 ppm (Staudt and Schoonen, 1995), although Meso-/Neoproterozoic levels of CAS typically fall in the range of 10~ ppm (Hurtgen et al., 2002; Gellatly and Lyons, submitted; Kah et al., submitted). Our CAS data for the Mesoproterozoic show (1) an abundance of highly variable 34S-enriched samples (Fig. 5.5-4), (2) overlap with penecontemporaneous barite and gypsum data (e.g., cf. with Fig. 5.5-2), and (3) stratigraphic trends analogous to those observed for Proterozoic sulphides. Most striking are the data from the 1.2 Ga Bylot Supergroup of northeastern Canada. This well-characterised locality affords a rare opportunity to compare CAS data from peritidal dolomite to isotopic results from interbedded, evaporative gypsum (Kah et al., 2001, submitted). Gypsum sulphur from the Bylot Supergroup shows a striking linear stratigraphic isotopic trend that increases up-section from roughly -t-22 to -t-32%0 over only 300 m of section. Complementary Sr isotope and trace element data argue for a global marine origin for the S isotope record (Kah et al., 2001 ), and the general

Chapter 5: Evolution of the Hydrosphere and Atmosphere

432

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834S (%0) Fig. 5.5-3. Summary of 634S ranges for metal sulphides relative to approximate deposit age for a broad temporal and spatial distribution of Proterozoic sedimentary exhalative (SEDEX) mineral deposits and host sediments. Numbers correspond to specific SEDEX deposits: (1) Aberfeldy (Scotland), (2) Balmat-Edwards (U.S.A.), (3) Jiashengpan (China), (4) Sullivan (Canada), (5) Sheep Creek (U.S.A.), (6) Amjore (India), (7) Gamsberg (South Africa), (8) Aggeneys-Black Mountain (South Africa), (9) Aggeneys-Broken Hill (South Africa), (10), Aggeneys-Big Syncline (South Africa), (11) McArthur River (Australia), (12) Mount Isa (Australia), (13) Hilton (Australia), (14) Dugald River (Australia), (15) Lady Loretta (Australia), and (16) Dariba-Rajpura (India). Deposits 7-15 are dated at c. 1.65 Ga (7-11) and 1.67 Ga (12-15). The sequence in the figure for 7-15 is not meant to convey a known chronological order. Amjore (6) is plotted as the mean of the reported 1.4-1.6 Ga age. Details and primary references are available in Lyons and Gellatly (in prep.).

correspondence between 8345 data from the gypsum and CAS in interbedded dolostones is unprecedented validation of the CAS method (Fig. 5.5-5; Kah et al., submitted). Also, the CAS data allow us to track the isotopic trends for marine sulphate stratigraphically above

433

5.5. Precambrian Sulphur Isotope Record

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834S (%0) Fig. 5.5-4. Isotope data for carbonate-associated sulphate (CAS) for four Proterozoic units: (a) Mescal Limestone, Apache Group, Arizona, U.S.A. (c. 1.2 Ga); (b) Helena Formation, Belt Supergroup, Montana, U.S.A. (c. 1.4-1.5 Ga); (c) Newland Formation, Belt Supergroup, Montana, U.S.A. (c. 1.4-1.5 Ga); and (d) Paradise Creek Formation, McNamara Group, Queensland, Australia (c. 1.7 Ga). Details are available in Gellatly and Lyons (submitted).

the highest gypsum layer. Consistent with the sulphide data described above, the combined gypsum/CAS profile from the Bylot Supergroup shows systematic variation of > 10%o over stratigraphic scales of 102 m. Similar "rapid" isotopic variability has been observed for the CAS from the c. 1.45 Ga Helena Formation of the Belt Supergroup (Fig. 5.5-6; Gellatly and Lyons, submitted) and the 1.3 Ga Dismal Lakes Group, Arctic Canada (Kah et al., submitted). The Helena data show 34S-depletions (~345CA S values as low as c. -1%0) outside those predicted for the global ocean--perhaps reflecting local reservoir effects superimposed on a global trend.

434

Chapter 5: Evo/ution of the Hydrosphere and Atmosphere

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Low sulphate concentrations in the Proterozoic ocean

Sulphide and sulphate S isotope data from temporally and spatially distinct units show systematic stratigraphic trends that are similar in magnitude and stratigraphic extent (i.e., variation by 10s of per mil over 10s to 100s of metres of section). This consistency in variation defines a style of rapid Mesoproterozoic isotopic variability that may mirror changes in the S isotope composition of the global ocean rather than a global distribution of similar local reservoir effects. (Where stratigraphic variation is not systematic, the sulphides are still consistently-34 S-enriched, and both the 6 34 Ssulphidc and (~345CAS data are highly variable.) When observed, the rapid, systematic variability in ~345 of marine sulphate, compared to variations spanning c. 20%0 over scales of 107-108 years for the Phanerozoic (Claypool et al., 1980), suggests a substantially reduced sulphate reservoir in the Mesoproterozoic

5.5. Precambrian Sulphur Isotope Record

435

Fig. 5.5-6. $34S variation in CAS in the Helena Formation (c. 1.45 Ga), Belt Supergroup, Montana, U.S.A. Details are available in Gellatly and Lyons (submitted).

ocean--as linked to lower pO2 in the ancient biosphere (Lyons et al., 2002; Gellatly and Lyons, submitted; Kah et al., submitted). Although a low-sulphate Proterozoic ocean is supported broadly by evidence for rapid change in marine sulphur isotope compositions and often observed low A34S values, there are few quantitative constraints on marine sulphate concentration after c. 2.3 Ga. In a recent model, Kah et al. (submitted) use available age constraints, theoretical platform subsidence rates, and observed stratigraphic variation in 34Sgypsum and ~348CAS (Fig. 5.5-5) to calculate rates of marine S isotope change in the Mesoproterozoic. The mass of the marine sulphate reservoir was then calculated using a time-dependent equation for isotopic change modified from Kump and Arthur (1999). Model results suggest that the Meso-

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proterozoic sulphate reservoir was c. 5-15% the size of the modern ocean reservoir (Kah et al., submitted). Although C-S redox cycling was important, at such reduced sulphate concentrations, isotopic perturbation of the global sulphate reservoir was decoupled from C isotope change--unlike the Phanerozoic (Berner, 2001). Variations in ~348sulphate were likely driven instead by the short-term relationships among weathering inputs, ocean ridgerelated hydrothermal activity, the extent of sulphate reduction, and pyrite burial--all of which have enhanced effects under the lower concentrations (and thus shorter residence time) of sulphate. Pyrite burial, in turn, was likely influenced by the extent of euxinic bottom waters. Ultimately, increased ocean-atmosphere oxygenation in the Neoproterozoic, driven by greater burial of organic matter, resulted in an increase in the size of the marine sulphate reservoir, possible evolution of the bacterial sulphate community with a corresponding increase in A34S, and increased coupling of the C and S isotope systems (Kah et al., submitted) (see also sections 5.2 and 5.3). As suggested above, low sulphate is reflected further in the abundance of 3aS-enriched Mesoproterozoic sedimentary and SEDEX pyrite. Although local basin restriction almost certainly played a major role (Lyons and Gellatly, 2002)--such restriction has also been invoked to explain 34S enrichments in Phanerozoic SEDEX mineralisation (e.g., Goodfellow and Jonasson, 1984)--extreme and frequent local reservoir isolation would have been required during much of the Proterozoic to be the sole cause of the pervasive positive S348pyrite values. By analogy, the Black Sea, with its active sulphate reduction in the water column shows open-system behaviour for the 634S of pyrite accumulating in the deep-basin sediments, despite its tenuous link to the Mediterranean (Lyons, 1997). Similar models for low ocean sulphate concentrations have been developed by other workers to explain the persistence, locally and globally, of 34S-enriched sulphides in the Proterozoic (Shen et al., 2002, 2003). Shen et al. (2002) argued for Mesoproterozoic seawater sulphate concentrations in the range of 0.5-2.4 mM, and further suggested that fractionations during BSR at sulphate concentrations < 1 mM might have been significantly greater than those proposed in recent models for the Archaean and earliest Proterozoic (e.g., Canfield et al., 2000). (Consistent with this interpretation, the critical lower limit for sulphate concentration able to yield comparatively large A34S during BSR was recently refined through experimental calibration to 200 laM (Habicht et al., 2002)). The temporal distribution of Precambrian gypsum is also linked to inferred seawater sulphate concentrations. Kah et al. (2001) noted that while evidence for evaporite deposition is not uncommon in rocks older than 1.2 Ga (e.g., section 7.12), extensive bedded marine CaSO4 evaporites are lacking. Kah et al. (2001) further suggested that a limited oxygenation event at c. 1.3 Ga may have increased marine dissolved sulphate to levels that favoured widespread evaporite minerals, such as the thick gypsum sequence observed in the Bylot Supergroup. Despite previous claims of chemical stasis during the Mesoproterozoic (Buick et al., 1995a), a carbon isotope excursion between c. 1.3 and 1.25 Ga suggests increased organic carbon burial and corresponding oxygenation of the biosphere (Kah et al., 1999, 2001; Bartley et al., 2001; Frank et al., 2003) (see also section 5.3). An additional possible factor in the paucity of early Precambrian CaSO4 evaporites is limited Ca 2+ availability under conditions of extreme CaCO3 saturation and precipitation (Grotzinger, 1989;

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Grotzinger and Kasting, 1993; as reviewed in Kah et al., 2001) (section 6.4). Extensive carbonate precipitation appears during much of the Proterozoic as unconventional depositional fabrics (Grotzinger and James, 2000), such as molar-tooth c a l c i t e t a product of rapid, early diagenetic filling of gas expansion cracks (Furniss et al., 1998). Under the hypothesised low oceanic sulphate concentrations, methane may have been the crack-forming gas (cf. Frank and Lyons, 1998). Concentrations of CAS are also a possible proxy for the amount of dissolved sulphate in ancient seawater. Precambrian CAS concentrations are often orders of magnitude below those observed in modern and Phanerozoic carbonates (Hurtgen et al., 2002; Gellatly and Lyons, submitted; Kah et al., submitted), which may reflect loss during diagenesis. Nevertheless, isotope relationships appear to remain intact (i.e., buffered to primary values), and concentration trends, while possibly shifted, still track independent (lithofacies) estimates of temporal variance in local evaporative enrichment (Gellatly and Lyons, submitted; Kah et al., submitted). Rather than diagenesis, the low Proterozoic CAS concentrations may record the low amounts of sulphate in the ocean (Hurtgen et al., 2002; Pavlov et al., 2003). Such low sulphate concentrations are consistent with a Proterozoic atmosphere comparatively enriched in methane due to the enhanced availability of reactive organic compounds for methanogenesis following limited remineralisation during BSR, reduced consumption of methane via anaerobic methane oxidation (AMO), and suppressed methane oxidation in a euxinic deep ocean beneath an atmosphere with 02 well below present levels (Pavlov et al., 2003; cf. Ohmoto, section 5.2). Interestingly, the extreme 345 enrichments that are so common in Proterozoic pyrite and are thought to reflect limited supplies of sulphate in the ocean, are rare today except at sites of extreme BSR linked to AMO (M. Formolo, 2003, pers. comm.). Models for evolving bacterial fractionations even during the Proterozoic (e.g., Canfield and Teske, 1996; Canfield, 1998) require careful considerations of the global and local availability of sulphate, and, ultimately a strong reliance on maximum observed fractionations between sulphate and sulphide, which may still be compromised by sulphate limitations and associated isotopic reservoir effects. As developed in an earlier section, net A34S can be small in a limited sulphate reservoir regardless of the magnitude of instantaneous bacterial fractionation. It is possible, for example, that the broadening in A34S observed at c. 0.8 Ga reflects increasing sulphate availability, as well as evolving bacterial communities. Finally, calculating A34S is hampered by the rapid A34S variation of ocean sulphate and the almost universal difficulty of establishing precise coevality of sulphide and sulphate isotope data (Hurtgen et al., 2002). Oceanic redox and cycling of trace metals and sulphur

An important implication of recent models for Precambrian sulphur bio-geochemistry is the suggestion that euxinic deep waters may have been present throughout much of the Proterozoic (Canfield, 1998). Despite growing evidence for low sulphate concentrations in Proterozoic sea water relative to more recent times, a substantial increase beyond the sulphate-poor Archaean ocean, in combination with a poorly oxygenated deep ocean, might have supported pervasive BSR in the water column and sediments. Canfield (1998)

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further suggested that cessation of extensive banded-iron formation (B IF) deposition at c. 1.8 Ga could reflect decreased dissolved Fe availability (solubility) in a sulphidic oceanmin contrast to conventional arguments favouring oxygenation of the deep ocean (see, however, Trendall and Blockley, section 5.4). This scenario requires a progression in seawater chemistry from the anoxic, Fe-rich ocean of the Archaean to sulphide-rich, Fe-poor deep waters during the Proterozoic when the atmosphere became sufficiently oxidising to increase the weathering flux of sulphate to the ocean (see Ohmoto, section 5.2, for a different view). The interval between 2.3 and 1.8 Ga was presumably marked by progressively increasing sulphate fluxes to the ocean such that sulphide production finally exceeded rates of reactive Fe input (dissolved and particulate) and shut down BIF formation. Proving a global distribution of Proterozoic deep euxinic waters remains the fundamental challenge, although studies emphasising iron geochemistry as a proxy for sulphide in the water column (Raiswell and Canfield, 1998; Lyons et al., 2003) are showing promise on the level of individual Proterozoic basins (Shen et al., 2002, 2003). (A sufficient number of such studies may establish a global redox context.) This model is roughly analogous to the "progressive ventilation" described by Berry and Wilde in 1978. Berry and Wilde attributed eventual deep ocean oxygenation to circulation patterns (high-latitude, deep-water production) during the glacial episodes of the late Neoproterozoic (sections 5.6-5.8). Canfield (1998) and Hurtgen et al. (2002) suggested, by contrast, that the Late Proterozoic glacial intervals were a time of intense oceanic stagnation, wherein sulphate (and correspondingly sulphide) levels were sufficiently low, given suppressed continental (riverine) inputs, to lead to renewed BIF formation (see also section 5.6). Hurtgen et al. (2002) further argued that high •345pyrite values and the high amplitude and frequency of S34ScAs variation associated with the snowball Earth deposits (sections 5.2, 5.6 and 5.7 offer different viewpoints) are consistent with almost complete reduction of sulphate in an isolated, anoxic glacial ocean. Ultimate oxygenation of the deep ocean was likely linked to a second stage of Earth-surface oxidation, which occurred late in the Neoproterozoic (Canfield, 1998, as described above). An intriguing implication of the sulphidic Proterozoic ocean model links possible effects on trace metal solubility/availability--for example, Fe and Mo--to nitrogen cycling (Anbar and Knoll, 2002). More specifically, low metal availability would limit prokaryotic N2 fixation catalysed by nitrogenase metalloenzyme systems and thus the oceanic supply of bio-available N. These redox-sensitive metals play other essential roles in the N cycle, such as during eukaryotic Mo-facilitated assimilation of nitrate. These connections between oceanic redox, trace metal solubility, and supplies of bio-available N might be expressed temporally in the ecological range and ultimately in the evolution of eukaryotic algae (Anbar and Knoll, 2002). While an attractive model, final validation may lie with our ability to confirm a globally euxinic Proterozoic ocean with sulphide concentrations sufficiently high (given low predicted sulphate) to limit metal availability. Such arguments for metal sequestration must be considered in light of current models for metal mobility (e.g., Helz et al., 1996; Raiswell and Canfield, 1998; Lyons et al., 2003). Alternatively, the global sulphidic ocean of Canfield (1998) can be viewed as a palaeoredox end-member.

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Whether global or not, widespread reducing conditions could have driven the metal-linked nitrogen limitations of Anbar and Knoll (2002), and abundant dissolved sulphide in the sediments alone could have yielded the detrimental sequestration of Mo (see Lyons et al., 2003)rebut regardless, conditions in the Proterozoic ocean were very different to the redox environment before and after.

Summary These highlights of recent progress provide a compelling illustration of what sulphur geochemistry has revealed about the Precambrian world. Many questions and controversies remain, but there is growing consistency among the interpretations of sulphur data and a broad array of complementary proxies. Through this consistency, a cohesive picture of evolving ocean-atmosphere oxygen availability and thus sulphate concentration and microbial ecology is emerging for the Precambrian (see, however, discussions in section 5.2 and chapter 6): 1. Despite controversial evidence for a very early origin (more than or equal to c. 3.47 Ga) of bacterial sulphate reduction (BSR), the Archaean ocean was dominantly low in sulphate as a product of the prevailing atmospheric 02 deficiency, also recorded in massindependent S fractionations and other palaeoredox proxies. Bacterial S fractionations would have been minimal in the low sulphate ocean except under local enrichments in sulphate "oases". 2. Arguments for widespread high rates of BSR in a warm, sulphate-rich Archaean ocean are difficult to support in light of recent studies of S isotope fractionations under high rates of BSR in modern natural settings. 3. By 2.7 Ga, corresponding to the first evidence of oxygenic photosynthesis, ~345 values for sulphate and sulphide and the offset between them argue for increasing sulphate in the ocean as recorded in the unambiguous signatures of BSR. These fractionations remained small, however, until c. 2.3 Ga when a critical threshold in sulphate concentration was exceeded. 4. By 2.3 Ga, corresponding to a time of intense organic carbon burial and thus atmospheric oxygenation (see, however, sections 5.2 and 5.3), continental weathering, and sulphate-rich runoff to the oceans, S isotopes expressed the full magnitude of fractionation observed experimentally via BSR in the absence of sulphate limitation. 5. Despite increasing oceanic sulphate concentrations and potential for large kinetic isotope effects during BSR, many pyrite samples from the Proterozoic show 34S enrichments consistent with sulphate reservoirs that were ultimately limited relative to those typical of the Phanerozoic. 6. The paucity of bedded gypsum prior to c. 1.2 Ga and the comparatively rapid S isotope variability observed in Proterozoic sulphate (including carbonate-associated sulphate) and sulphide minerals suggest global sulphate limitations that may have been exacerbated by local conditions. 7. Increasing Proterozoic seawater sulphate in a still poorly ventilated deep ocean may have led to sulphidic bottom waters, which impacted the bio-availability of essential

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trace metals and thus eukaryotic ecological expansion. A globally euxinic deep ocean remains conjectural pending further study. 8. By the Neoproterozoic, oxygen concentrations increased to levels that supported generally high seawater sulphate concentrations (except during glacial intervals) and a stronger oxidative loop in the sulphur cycle capable of driving S isotope fractionations to values as high as those observed today. Ultimately, the strength of these interpretations is severely limited by the present scarcity of high-resolution, well-dated, unaltered geochemical data, particularly for sulphate in the Precambrian ocean. But as for any good hypothesis, recent studies of Precambrian sulphur chemistry are providing a solid platform for ongoing and future work.

5.6.

EARTH'S TWO GREAT PRECAMBRIAN GLACIATIONS: AFTERMATH OF THE "SNOWBALL EARTH" HYPOTHESIS

G.M. YOUNG

Introduction There is evidence of two major periods of glacial activity, near the beginning and end of the Proterozoic Eon (2.5 Ga to c. 540 Ma) (see also Williams, section 5.7, which is a complementary paper to this one, and section 5.8). Glacial interpretations of Proterozoic diamictite-bearing successions have not gone unchallenged. Crowell (1957) and Schermerhorn (1974) both pointed out that "pebbly mudstones" or diamictites can also form by mass-flow processes. These studies, and others suggesting that some diamictites may be related to major impacts of extraterrestrial bodies, have led to refinement of glacial criteria and have encouraged closer inspection of the evidence for ancient glaciations. Re-investigation of many diamictite-bearing successions throughout the world culminated in the great compilation of pre-Pleistocene glacial occurrences by Hambrey and Harland (1981), which remains the most complete source of factual information on such deposits. Evans (2000) provided a recent assessment of geochronological and palaeomagnetic studies.

Emergence of the Snowball Earth Hypothesis Mawson (1.949) and Harland (1964) suggested the possibility of widespread Precambrian glaciations. Harland and Bidgood (1959) pioneered the use of palaeomagnetism in the study of Precambrian glacial deposits and suggested that some may have formed at low palaeolatitudes. Embleton and Williams (1986), Schmidt et al. (1991) and Sohl et al. (1999) presented further evidence of low latitude glaciation in the Neoproterozoic of Australia (section 5.7). Following these investigations, Kirschvink (1992) introduced the phrase "snowball Earth hypothesis". He also suggested that isolation of the ocean from the atmosphere could have caused build-up of hydrothermal iron in the oceans (section 5.5). 771e Precambrian Earth: Temposand Events FAited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Precipitation of the enigmatic Neoproterozoic iron-formations was considered to be the result of re-oxygenation of the oceans at the end of the glaciation. Hoffman et al. (1998b) drew attention to stable isotopic data (613C) (see also section 5.3) from carbonates above and below Neoproterozoic glacial deposits in Namibia (section 5.8). They suggested that the unusual association of carbonates and glaciogenic diamictites was best explained by dramatic and extreme climatic perturbations summarised in the snowball Earth hypothesis.

Neoproterozoic Glaciations in the Light of the Snowball Earth Hypothesis Geochronology Neoproterozoic glaciogenic rocks, in common with most sedimentary rocks, are notoriously difficult to date. Contemporaneity of each "phase" of the Neoproterozoic glaciations is a sine qua non of the snowball Earth hypothesis (SEH) but it remains to be demonstrated. Neoproterozoic glacial deposits range from about 800 Ma (Martins-Neto and Hercos, 2002) to the base of the Cambrian at c. 540 Ma. This huge time period of c. 260 My permits many possible interpretations. The Port Askaig Formation in the west of Scotland, long considered to be part of the c. 600 Ma Varanger glaciation, has recently been re-assigned to the c. 730 Ma "Sturtian" episode (Prave, 1999). The affinities and ages of glacial deposits in Namibia, which were the platform from which the promotion of the SEH was launched by Hoffman et al. (1998b), are contentious (Kennedy et al., 1998) (see discussion by Frimmel, section 5.8). There is current debate not only about the ages of such deposits but also about the number of glacial episodes that occurred. These deficiencies concerning the ages and number of Neoproterozoic glaciations do not disprove the SEH but their resolution is a prerequisite for its acceptance (section 5.8). The carbonate problem, including cap carbonate The nature and significance of carbonate rocks associated with many Proterozoic glaciogenic successions has long been debated (Schermerhorn, 1974; Williams, 1975; Roberts, 1976). Cap carbonates, which overlie some Neoproterozoic glacial successions were shown to have an unusual C-isotopic signature (Knoll et a1.,1986). Hoffman et al. (1998b) concluded that low 313C values in carbonates both below and above glaciogenic deposits in Namibia (section 5.8) could indicate the near-extinction of marine photosynthetic organisms during extreme glacial periods. The utility of carbon isotopes as a unique indicator of ancient life has recently been brought into question (Fedo and Whitehouse, 2002). Subsequently Hoffman and Schrag (2002) have retreated from this "extreme" interpretation and suggested that low 613C values in carbonates below Neoproterozoic glacials could be due to build-up of methane, whereas the post-glacial low values may be due to incorporation of large amounts of atmospheric CO2 into world oceans. The idea of a CO2-induced supergreenhouse causing the end of glaciation and alkalisation of the world's oceans is not in accord with studies of post-glacial siliciclastic rocks (e.g., Young and Nesbitt, 1999), which show a gradual upwards increase in weathering at the end of the Palaeoproterozoic Gowganda glaciation in Canada. This finding contrasts with the predicted highly weathered material, which should occur immediately above the glacial deposits as a residue from the

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supposed flushing of alkalis from the continental crust. Apart from sparse and in some cases, conflicting isotopic data (Kennedy et al., 1998), there is a dearth of major and trace element geochemical information from cap carbonates and carbonates that occur within glaciogenic sequences (see, however, Frimmel, section 5.8, for an example). Such data should shed light on their origins and affinities.

Indicators of strong seasonality at low palaeolatitudes Varved deposits imply significant annual temperature variations. Laminated sedimentary rocks of this kind have been described from Neoproterozoic glacial sequences in the North Atlantic area (Spencer, 1971; Hambrey, 1983) and by Pettijohn (1959), Jackson (1965) and Young (2001) from Palaeoproterozoic occurrences. Williams and Schmidt (1997) concluded from palaeomagnetic evidence that the Gowganda Formation was deposited at low latitudes, which today are characterised by small seasonal temperature variations. Williams (1975) proposed a significantly increased obliquity in order to accommodate these and other features of the Proterozoic glacial record (section 5.7). Large polygonal structures interpreted as fossil ice-wedges occur in Neoproterozoic glacial successions in South Australia (Williams, 1986), Scotland (Spencer, 1971), Ireland (Johnson, 1993) and elsewhere. These large sandstone-filled polygons are thought to indicate strong seasonal temperature differences over long periods of time (Williams, 1986, 1998a) (sections 5.7 and 5.9). Hoffman (1999a) contended that such features can form at low latitudes by referring to their reported occurrence today in Hawaii and on Mount Kilimanjaro. Many of these modern structures are smaller and different in configuration (linear as opposed to polygonal) from those reported in Neoproterozoic sequences. They result from diurnal temperature variations that are a reflection of their high altitude. These and other explanations for large and deep Proterozoic ice-wedge structures, which formed at low palaeolatitudes near sea level are unsatisfactory. Occurrence of such features at low pa!aeolatitude~ in Neoproterozoic glaciogenic successions presents a conundrum comparable to that posed by varved bedding in the Gowganda Formation.

Thickness and facies of Neoproterozoic glaciogenic successions According to the SEH, glaciation of the planet should have proceeded very rapidly (Hoffman et al., 1998b, 2000), as a result of "runaway albedo". Likewise, disappearance of global ice cover should have been equally quick once atmospheric CO2 concentrations reached critical levels. It is difficult to reconcile the supposed cessation of the hydrologic cycle with the preservation of many kilometres of glaciogenic deposits (e.g., Young and Gostin, 1989). Thick successions of glaciogenic rocks cannot be explained as being due to the simple existence and destruction of thin glacier ice. Ice advance on a large scale is also a prerequisite and such advances, at the same time as melting, can only be produced under an active hydrologic regime. Hoffman (2000) proposed that a vigorous hydrologic cycle could be maintained by sublimation of ice in low latitudes and precipitation at high altitudes (due to lapse rate). If most of the continental lithosphere occupied low palaeolatitudes then such sublimation would mostly have taken place from the thin and limited ice cover that existed there. Without a huge moisture reservoir such as the presumably

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frozen oceans such sublimation/precipitation events would have resulted in recycling of a meagre continental ice cover and could not have produced the thick and complex glacial successions observed in the record. Complex stratigraphies (Spencer, 1971; Yeo, 1981; Young, 1992) attest to a complex glaciological history. Thick dropstone-bearing successions (Young and Gostin, 1991; Condon et al., 2002) indicate a gradual retreat of glaciers rather than the rapid demise suggested by proponents of the SEH. Cross-bedded sandstone bodies associated with diamictite successions (Young and Gostin, 1991; Arnaud and Eyles, 2001) indicate glaciation under a temperate regime rather than the extreme conditions predicted by the SEH. Some of these sandstones contain large scale cross-beds that are interpreted to indicate the existence of large bodies of open water (Arnaud and Eyles, 2001).

Iron-formations and Neoproterozoic glacial deposits Kirschvink (1992), Klein and Beukes (1993) and Hoffman et al. (1998b) suggested that Neoproterozoic banded iron-formation (BIF) may be the result of isolation of the ocean and atmosphere by a more-or-less world-encircling ice cover (see also section 5.5). Following significant Fe-enrichment, as a result of oceanic hydrothermal activity, destruction of the ice cover would have caused iron precipitation. The distribution of Neoproterozoic BIF is, however, much more restricted than that of glaciogenic deposits. Yeo (1981, 1986) outlined a viable mechanism to explain the Neoproterozoic BIF (see also section 5.4). The model involves hydrothermal activity in Red Sea-type rift settings, accompanied by thermal overturn and precipitation of Fe. Glaciers descending from rift flanks provided icebergs, which melted and emplaced isolated clasts (dropstones) in the iron-formations. Positive Eu anomalies in shale-normalised REE plots, suggest hydrothermal influence (Yeo, 1981; Neale, 1993; Lottermoser and Ashley, 2000). Such chemical evidence is equally compatible with the glaciated rift model or the SEH but strong evidence of rift activity, provided by dramatic facies and thickness changes in Neoproterozoic successions such as the Rapitan Group in NW Canada (Yeo, 1981; Eisbacher, 1985) and in the Adelaide geosyncline in Australia (Young and Gostin, 1989, 1991) lends support to the former interpretation. The Neoproterozoic iron formations can be accommodated without invoking a totally frozen planetary surface. According to the SEH, precipitation of iron should have followed disintegration of the ice but iron-rich rocks in the Rapitan Group underlie the main body of glaciogenic diamictite. Neoproterozoic glaciogenic deposits are extremely widespread but it must be kept in mind that they may have had a much more restricted global distribution prior to the breakup of the supercontinent Rodinia (sections 3.10, 3.11 and 5.8). They may have formed over a period of about 300 My. Until more precise geochronological control is obtained, it is premature to interpret them as indicating that the entire planetary surface was frozen. Many puzzles remain, including evidence of strong seasonality at apparent low palaeolatitudes (increased obliquity?) (section 5.7). Strong evidence of thick glaciogenic successions resulting from prolonged glacial activity under temperate conditions is not explained by the SEH. The limited distribution of Neoproterozoic BIF, their variable stratigraphic

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position visa vis associated glacial deposits and their geochemistry all support deposition in glaciated rift basins.

Palaeoproterozoic Glacial Deposits Palaeoproterozoic glacial deposits are much rarer than those of the Neoproterozoic. Limited distribution of early Precambrian glaciogenic rocks could reflect the existence of a less extensive continental crustal area or it may be a function of preservation. There is very little evidence of glaciation in Archaean sedimentary rock sequences. (Page and Koski, 1973; Young et a1.,1998; Crowell, 1999). In North America, early reports by Coleman (1908, 1925) and Blackwelder (1926) focused attention on Precambrian glaciations. Young (1970) proposed that glaciogenic deposits in various localities in Canada and the U.S.A. may have been products of a single large continental ice sheet. Palaeoproterozoic glaciogenic rocks have been identified in Finland (Marmo and Ojakangas, 1984) (section 3.9), Western Australia (Trendall, 1981; D.M. Martin, 1999) and in South Africa (Visser, 1981; Eriksson et al., 2001 c). As in the Neoproterozoic, there have been several reports suggesting that glaciation in the Palaeoproterozoic took place in low palaeolatitudes (Evans, 1997; Williams and Schmidt, 1999). The Huronian stratigraphic succession contains evidence of three successive glacial intervals separated by significant thicknesses of non-glacial sedimentary rocks. The lower two diamictite-bearing units, the Ramsay Lake and Bruce Formations, are only developed locally and appear to contain a high proportion of reworked material from underlying Huronian formations, whereas the Gowganda Formation is much more widespread and is mainly composed of material from the underlying Archaean rocks of the Superior Province. The Medicine Bow Mountains of Wyoming contain a near-identical Palaeoproterozoic succession (Houston et al., 1992). In most other areas glaciation is signalled by the occurrence of a single glaciogenic unit. The age of the widespread glaciogenic rocks of the Gowganda Formation is not well constrained but the Huronian Supergroup is considered to have been laid down between about 2.4 and 2.2 Ga, the latter constraint being the age of the Nipissing diabase. In Finland and adjacent areas of Russia, glacial deposits have been identified in the Sariolian Group (section 3.9). The glaciogenic rocks overlie the Sumian Group, which includes 2440 Ma old mafic igneous rocks (Ojakangas et al., 2001). Palaeoproterozoic glacial deposits are represented in Western Australia by the c. 2.45-2.2 Ga-old Meteorite Bore Member of the Kungarra Formation (Trendall, 1981; D.M. Martin, 1999). The glacial rocks form part of the Turee Creek Group, which is considered to have formed in a back-arc compressive cratonic basin (Blake and Barley, 1992; Powell et al., 1999). The Meteorite Bore Member is about 270 m thick and occurs about 1800 m above the base of the Kungarra Formation. It is separated from BIF by a thick siliciclastic succession. In South Africa, the Palaeoproterozoic glacial Makganyene Formation rests with an angular unconformity on a thick siliciclastic Koegas Formation. This deltaic formation gradationally overlies a granular iron-formation (Griquatown IF), underlain by the Kuruman

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Banded Iron Formation (2480 Ma; Nelson et al., 1999). These chemical to clastic sedimentary rocks make up a large second-order shallowing-upwards succession of at least 50 My duration (Altermann and Nelson, 1998) and the glacial deposits are separated from this succession by an angular unconformity. The lavas of the Ongeluk Formation, overlying the glaciogenic deposits with a marked disconformity (Altermann and H~ilbich, 1990, 1991), have been dated at 2222 Ga (Cornell et al., 1996). Locally restricted manganiferous ironstones, covered by dolomites, follow the up-to-600 m-thick craton-wide lavas. There is thus no direct relationship between the glacial deposits and the BIF.

Tectonic setting of the Palaeoproterozoic glaciogenic rocks The Huronian Gowganda Formation on the north shore of Lake Huron has been interpreted as having formed at the time of transition from a rift to a passive margin (Young and Nesbitt, 1995). The lower Huronian rocks (up to the base of the Gowganda Formation) are mainly fluvial, deltaic and lacustrine (?) deposits with a restricted distribution. They display evidence of fault control on thickness and facies changes. The Snowy Pass Supergroup in SE Wyoming comprises a strikingly similar stratigraphic succession that probably formed in a similar tectonic setting to the Huronian (Karlstrom et al., 1984). Likewise in Finland, Ojakangas et al. (2001) inferred from facies, distribution and preservation of the Palaeoproterozoic glaciogenic rocks in the Sariolian Group that they were deposited on a subsiding continental margin (see also section 3.9). By contrast, glaciogenic rocks of the Meteorite Bore Member in Western Australia formed in a compressional setting, culminating in development of a foreland basin (Horwitz, 1982; Blake and Barley, 1992; Powell et al., 1999). ~

Comparisons with the Neoproterozoic Most Palaeoproterozoic glaciogenic successions have a much simpler stratigraphy than those of the Neoproterozoic. The Neoproterozoic glaciations may have spanned almost 300 My, whereas deposition of the entire Huronian Supergroup (c. 12 km of stratigraphy, of which a small proportion is interpreted to be glacial) only involved about 200 My. In the thick Huronian succession, only one of three glacial formations (the Bruce Formation) is overlain by carbonate-rich rocks. The widespread Gowganda Formation has no cap carbonate but shows a gradual upwards transition to deltaic sediments, followed by fluvial arkosic sediments of the lower Lorrain Formation, then into quartzarenites of the upper Lorrain. Thus, the cap carbonates predicted by the SEH are absent from the majority of Palaeoproterozoic glaciogenic successions and both the onset and end of the glacial episode appear to have been gradual.

Thickness and facies of glaciogenic deposits Thicknesses of up to 3000 m have been described from the Palaeoproterozoic Gowganda Formation in Ontario (Schenk, 1965; Lindsey, 1969). Some thickness changes appear to be related to contemporaneous faulting (Young and Nesbitt, 1985) and large scale slumping during establishment of a continental margin (Card, 1978; Young et al., 2001). Although down-slope movement of sediment may occur in subglacial environments (i.e., under

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a frozen ocean), the large thicknesses of many glaciogenic formations (whatever the mechanism of final emplacement) can only be explained under an active hydrologic regime. The presence of abundant sandstone and orthoconglomerate intercalations are more typical of a long-lived temperate style of glaciation acting under a vigorous hydrologic cycle than the totally frozen planet envisaged in the SEH.

Stratigraphic relationship to iron formations The Huronian Supergroup is interpreted by many as spanning the period when the Earth's atmosphere became oxygenated (Ohmoto, section 5.2, for full discussion of different hypotheses). The upper part of the glaciogenic Gowganda Formation includes red beds, providing part of the geological evidence for the presence of at least some atmospheric oxygen at that time. Under snowball Earth conditions iron formations should theoretically follow glacial deposits, as they are thought to be consequent on the disintegration of a worldspanning oceanic ice cover, permitting oxygenation of the Fe-charged oceans (Kirschvink, 1992) (section 5.5). Palaeoproterozoic glacial deposits do not appear to have a close association with iron-formations. The classical Superior-type iron-formations of western Lake Superior occur in a foreland basin setting (Young, 1983; Hoffman, 1987; Ojakangas et al., 2001) that formed more than 200 My after deposition of the glacial Gowganda Formation. In South Africa and Western Australia the stratigraphic association between Palaeoproterozoic glacial deposits and BIF is by no means more intimate, and does not appear to fit the predicted sequence of the snowball Earth hypothesis. Just as in the Great Lakes region of North America, the Palaeoproterozoic iron formations in Western Australia (and possibly in South Africa) were deposited in compressional plate tectonic regimes that culminated in production of a foreland basin (Visser, 1981; Blake and Barley, 1992; Powell et al., 1999). In Western Australia the glaciogenic Meteorite Bore Member overlies the associated BIF and is separated from them by a thick succession of siliciclastic rocks (D.M. Martin, 1999). In South Africa the glaciogenic Makganyene deposits were also laid down above the iron formation and are separated from it by the Koegas Formation clastic sedimentary rocks. Additionally, they are separated from the lavas of the overlying Ongeluk Formation (for which a low palaeolatitude has been established; Evans et al., 1997) by a disconformity (Altermann and H~ilbich, 1990, 1991). The stratigraphy of Palaeoproterozoic glacial deposits and BIF does not conform to the predictions of the SEH. The iron formations appear to have formed in response to uplift and oxygenation of large amounts of dissolved iron during periods of sediment starvation (for example, following the transition from passive margin to foreland basin) (see Trendall and Blockley, section 5.4, for full discussion of iron-formation genesis). Sudden onset and end of glaciation? An integral part of the snowball Earth model is a relatively sudden onset of glaciation due to runaway albedo and an equally sudden demise when atmospheric CO2 values reached threshold values. Rapid onset and demise of the global glaciations and the virtual cessation of the hydrologic cycle should have resulted in thin glacial successions and stratigraphically abrupt top and basal contacts to the glacial sediments. Detailed investigations of the

5.6. Aftermath of the "Snowball Earth" Hypothesis

447

Gowganda Formation show that none of these criteria is met. The Gowganda Formation is underlain in southern areas by a thick arenite unit, the Serpent Formation. Geochemical investigations (Fedo et al., 1997) show that this unit differs from other sandstonerich formations in the Huronian Supergroup in being mineralogically and chemically less mature. Fedo et al. (1997) interpreted this to indicate a gradual deterioration of climate prior to onset of the Gowganda glaciation. Young and Nesbitt (1999) showed, using a Chemical Index of Alteration (fully discussed in section 5.10; see also section 5.11), that the post-glacial part of the Gowganda Formation records a gradual upwards increase in weathering--a trend opposite to that predicted by the SEH. The transition to highly weathered materials above the Gowganda Formation (Young, 1973) is far from sudden, involving gradual upwards increase in weathering index through thick deltaic and fluvial deposits, prior to establishment of the highly weathering regime typical of the upper Lorrain Formation. Detailed sedimentological, chemical and mineralogical investigations of the Gowganda Formation and its enclosing strata support a long-lived glacial epoch with gradual onset and demise of glacial conditions. Conclusions

Poorly dated Neoproterozoic glaciations appear to range over almost 300 My. The number of glacial episodes is not known nor has world-wide contemporaneity of any one episode been demonstrated (see also section 5.8). It is therefore premature to interpret these longlived, widespread and complex glacial episodes as the product of global glaciation ("snowball Earth hypothesis"; SEH). Evidence of glaciation at sea level in low palaeolatitudes has been used to support the SEH, but these glacial sequences also contain evidence of strong seasonality. Alternative solutions (admittedly also speculative) include invoking a Precambrian Earth with a significantly higher obliquity (> 54 degrees) (Williams, section 5.7). Thick successions of diamictite, associated waterlain deposits and sedimentological and geochemical evidence of gradual climatic deterioration and amelioration at the beginning and end of the Proterozoic glaciations all provide arguments against the snowball Earth hypothesis (section 5.7). Explanations offered by supporters of the SEH for Neoproterozoic BIF are not in accord with their sporadic development, compared to associated glaciogenic facies. The stratigraphic distribution, facies associations and geochemistry of these BIFs are all explained by a more conservative model involving deposition in an extensional plate tectonic setting where glaciers debouched into developing rifts of Red Sea-type. Palaeoproterozoic glaciations appear to have been less widespread, less stratigraphically complex and possibly of shorter duration than those of the Neoproterozoic. They occurred in passive margin settings in North America and in foreland basins in Western Australia and in South Africa. In North America the glaciations are separated by more than 200 My from major BIE In Western Australia and South Africa, BIF accumulation took place prior to glaciation, not afterwards as predicted under the SEH. It is clear that the Earth underwent important climatic perturbations at the beginning and end of the Proterozoic Eon. The cause of the Earth's cold episodes is poorly understood but it is interesting that many of the ancient glaciations are preceded by periods of supercon-

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tinentality (sections 3.2, 5.3 and 3.9) and many glaciogenic deposits, particularly those of the Neoproterozoic, are preserved in rift basins heralding supercontinental break-up. Perhaps mountain building associated with the production of supercontinents led to enhanced weathering and CO2 drawdown. Location of a supercontinent in low palaeolatitudes would have enhanced albedo and thus contributed to global cooling. In spite of nearly 200 years of investigation of glacial rocks and processes the cause of these cool periods remains elusive (see further discussion in the next section, 5.7).

5.7.

THE PARADOX OF PROTEROZOIC GLACIOMARINE DEPOSITION, OPEN SEAS AND STRONG SEASONALITY NEAR THE PALAEO-EQUATOR: GLOBAL IMPLICATIONS

G.E. WILLIAMS

Introduction--Enigmatic Proterozoic Glaciations Proterozoic glaciogenic successions occur on all continents and exhibit enigmatic features that have prompted vigorous debate concerning the Proterozoic global environment (see also complementary section 5.6 by Young). The suggestion that Neoproterozoic glaciation occurred in low palaeolatitudes (Harland, 1964) was supported by a palaeomagnetic study of the c. 600 Ma Marinoan (Varanger) glaciogenic Elatina Formation in South Australia that indicated the magnetic remanence was primary (Embleton and Williams, 1986). Three fold tests on soft sediment folds in that formation (Fig. 5.7-1) all proved positive (Sumner et al., 1987; Schmidt et al., 1991; Schmidt and Williams, 1995), vindicating the conclusion that the remanence was acquired early. However, these studies sampled the field for < 100 years, prompting objections that the Elatina palaeopole was a "virtual geomagnetic pole". Regional palaeomagnetic data for the Elatina Formation overcame those doubts and identified magnetisation reversals within specimens and a rough magnetostratigraphy (Schmidt and Williams, 1995); this work was supported by Sohl et al. (1999). Combined data for all Elatina sites yielded a palaeolatitude of 7.9 -I- 3 ~ (Schmidt, 2001). Palaeomagnetic data for core from a drillhole in Western Australia supported a low palaeolatitude for Marinoan glaciation and implied a low palaeolatitude for the c. 750 Ma Sturtian glaciation (Pisarevsky et al., 2001). Low palaeolatitudes (6 -t- 4 ~ and 4 -t- 6 ~ were determined also for the c. 750 Ma Rapitan glaciation in Canada (Park, 1997). In addition, palaeomagnetic data for Palaeoproterozoic volcanic rocks in South Africa (Evans et al., 1997) and Huronian sedimentary rocks in Canada (Williams and Schmidt, 1997; Schmidt and Williams, 1999) suggested low palaeolatitudes for Palaeoproterozoic (c. 2.3 Ga) glaciation (section 5.6). Tidalites in Marinoan and Sturtian glaciogenic successions (Williams, 1994a, 2000) confirm glaciomarine deposition. Cold climates near sea level in low palaeolatitudes are all the more puzzling because the faster rotation of the Proterozoic Earth (Williams, 2000) (section 5.9) would have caused less efficient polewards transport of heat, resulting in The Precambrian Earth: Temposand Events Editcd by P.G. Eriksson, W. Altcrmann, D.R. Nelson, W.U. Mueiler and O. Catuncanu

5.7. Paradox of Proterozoic Glaciomarine Deposition

449

Fig. 5.7-1. A bed surface of the Marinoan Elatina Formation at Warren Gorge, South Australia, showing symmetrical ripple marks generated by wave action. The ripple marks drape cuspate anticlinal folds 30-50 cm apart caused by soft-sediment gravity sliding on a tidal delta (Williams, 1996). Comparable cuspate folds caused by gravity sliding occur in late Neoproterozoic deltaic deposits, Newfoundland (Myrow and Hiscott, 1991). Hammer is 33 cm long.

slightly warmer equatorial regions and substantially cooler poles (Kuhn et al., 1989). Curiously, however, unequivocal evidence for Proterozoic glaciation in high palaeolatitudes is lacking (Evans, 2000). Neoproterozoic glaciogenic successions have other enigmatic features. Carbonates commonly are interbedded with or cap glaciogenic rocks (Williams, 1979; Corkeron and George, 2001; Kennedy et al., 2001), and iron-formations occur with glaciogenic deposits of Sturtian-Rapitan age in Canada, South Africa and South Australia (Breitkopf, 1988; Young, 1988; Drexel et al., 1993) (see also sections 5.6 and 5.8).

Open Seas During Neoproterozoic Glaciations Much evidence indicates that seas were unfrozen across wide areas and for lengthy time intervals during Neoproterozoic glaciations. 1. Tidalites in the Elatina Formation display wave-generated ripples (Fig. 5.7-1) throughout a 20 m thick succession that records many years of deposition with virtually continuous wave activity and open seas (Williams, 1996).

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2. Glaciomarine deposits with thick (up to c. 2700 m), widespread mudstone-withdropstone facies occur in many glaciogenic successions of Sturtian and Marinoan age (Preiss, 1987; Drexel et al., 1993; McMechan, 2000; Condon et al., 2002). Such deposits attest to temperate glacial conditions during long intervals with voluminous sedimentladen meltwater plumes and icebergs calving into open seas. 3. Tidal rhythmites (see also section 7.5) in the Elatina Formation in South Australia record the annual (or seasonal) oscillation of sea level, which occurred continuously for at least 60 years (Williams, 2000) (section 5.9). Tidal rhythmites interbedded with diamictites in the Pualco Tillite of the Sturtian glaciogenic succession in South Australia (section 5.9) and tidal rhythmites in the Sturtian Chambers Bluff Tillite 700 km to the northwest also record this oscillation (G.E. Williams, unpubl, data). The annual sea-level oscillation is of non-tidal origin. Data for moderate to low latitudes indicate a direct relation between sea temperature and sea level (Roden, 1963; Wunsch, 1972), and Pattullo (1966) concluded that the annual oscillation of sea level between latitudes 45~ and 45~ is ascribable almost entirely to changes in heat content of the sea. Mellor and Ezer (1995) found that the annual variation of sea level for the Atlantic Ocean between latitudes 66~ and 66~ approximates the heating-cooling cycle of each hemisphere. If the sea had been frozen-over during Neoproterozoic glaciations (see also sections 5.6 and 5.8), the annual oscillation of sea level could not have occurred because the ice cover would have insulated the sea from seasonal changes of temperature. The occurrence of ripples throughout the Elatina rhythmites and locally abundant dropstones in the Pualco Tillite rhythmites provides independent evidence for long lasting, unfrozen seas during rhythmite deposition. The Seasonality Paradox The climatic paradox identified by Williams (1975)--that of large seasonal changes of temperature in low palaeolatitudes during Neoproterozoic glaciationmhas been amply verified and must be addressed. Large seasonal changes of temperature are indicated by spectacular Marinoan periglacial sand wedges that formed in a then-coastal area in South Australia (Fig. 5.7-2). Several generations of sand wedges, as much as 3+ m deep and marking polygons up to 30 m across, occur with other periglacial structures in a fossil permafrost regolith of brecciated quartzite and in overlying periglacial quartzose aeolianite (Williams and Tonkin, 1985; Williams, 1986, 1994a, 1998). Neoproterozoic periglacial sand wedges also occur in Mauritania (Deynoux, 1982) and Scotland, Norway and Spitsbergen (see summary in Williams, 1986). Sand wedges and ice wedges occur in Antarctica and ice wedges are widespread in the Arctic (P6w6, 1959; Washburn, 1980; Black, 1982; Karte, 1983). Such wedges are confined to periglacial regions marked by a strongly seasonal climate; importantly, sand wedges are best developed in Antarctica where the absolute seasonal air-temperature range exceeds 60~ and the mean monthly air-temperature range is c. 40~ Wedges show vertical lamination and in plan they define polygons c. 10-30 m across. It is widely agreed

5. 7. Paradox of Proterozoic Glaciomarine Deposition

451

Fig. 5.7-2. Marinoan periglacial sand-wedges, Stuart Shell South Australia. The large wedge (2) is 3 m deep, contains steeply dipping laminae of pebbly coarse sandstone, and is developed in a permafrost regolith of brecciated Mesoproterozoic quartzite. Two deformed sand wedges of an earlier generation (1) occur within the breccia, and a third-generation wedge (3) occurs in the upper part of the large wedge and in overlying Marinoan periglacial-aeolian quartzose sandstone. Upturning of material next to the wedges records summer expansion of the permafrost. The several generations of wedges indicate climate fluctuations on a 103-year time-scale. Identical sand wedges, including wedge-in-wedge structure, are forming in Antarctica under a strongly seasonal periglacial climate. From Williams (1993). that the wedges develop from thermal contraction cracks c. 1-5 mm wide and several metres deep that form in the upper part of permafrost with rapid drops of temperature during repeated severe winters. Ice wedges occur in humid periglacial areas where water freezes in the cracks, and sand wedges mark drier periglacial areas where the cracks are filled by drifting sand. Measurements across sand wedges in Antarctica over two decades indicated mean growth rates of up to 1 mm yr -l (Black, 1982), and estimated ages of ice wedges in Alaska based on measured growth rates of 1-3 mm yr -1 were verified by radiocarbon dating (Black, 1952, 1982). These observations confirm that periglacial wedges are forming actively in high latitudes under strongly seasonal climates. Strong seasonality clearly does not inhibit the development of glaciers, because such regions saw repeated glaciations during the Pleistocene.

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Periglacial sand wedges are reliable indicators of past climate because they formed through mechanical processes and so their interpretation avoids uncertainties in the nature of the former atmosphere and hydrosphere (see summary, Ohmoto, section 5.2) and later diagenetic alteration. The Marinoan sand wedges therefore imply an absolute seasonal airtemperature range exceeding 60~ and mean monthly air temperatures of - 3 5 ~ or lower in midwinter and ~< 4~ in midsummer (see Karte, 1983). This frigid, strongly seasonal climate occurred in a coastal area near the palaeoequator. Regarding Palaeoproterozoic glaciation, regularly laminated argillites with abundant dropstones and till pellets that occur in the glaciogenic Gowganda Formation in Canada have been interpreted as varvites (Young, 1981; Mustard and Donaldson, 1987) (section 5.6). Furthermore, the Ramsay Lake Formation in Canada displays structures interpreted as periglacial ice-wedge casts (Young and Long, 1976). A strongly seasonal glacial climate that caused repeated melting and runoff, and thermal contraction cracking, evidently occurred in low palaeolatitudes also during the Palaeoproterozoic. By contrast, the Phanerozoic is marked by equable climates near the palaeoequator. Pangaea had a seasonal temperature range of <~ 5~ in low palaeolatitudes during the Permian, whereas seasonal temperature ranges > 40~ were confined to middle and high southern palaeolatitudes (Crowley et al., 1989; Gibbs et al., 2002). In present low latitudes (<~ 10~ the mean monthly temperature range near sea level is < 2~ and the mean diurnal temperature range is < 10~ (Mtiller, 1982). Large seasonal temperature ranges near the Proterozoic equator seem especially strange because according to long-held belief the Earth's obliquity was smaller in the past (Goldreich, 1966), which would have rendered equatorial climates of the Proterozoic even more equable than those of the Phanerozoic. The Proterozoic Global Environment

What can the paradox of Proterozoic glaciomarine deposition, open seas and strong seasonality near the palaeoequator reveal about the Proterozoic Earth? A Proterozoic snowball Earth ?

According to the snowball Earth hypothesis (Kirschvink, 1992; Hoffman et al., 1998b) (section 5.6), negative ~13C values (see also section 5.3) for carbonates bracketing Neoproterozoic glaciogenic deposits in Namibia (section 5.8) reflect a collapse of biological activity in the surface of the world ocean for many millions of years, caused by a frozenover Earth. It was argued that the Neoproterozoic world ocean was frozen to an average depth of > 1 km, the mean global temperature was - 5 0 ~ with mean surface temperatures of - 8 0 ~ to - 110~ in high latitudes (Baum and Crowley, 2001), and the hydrological cycle virtually shut down. Coldest conditions occurred near the start of global glaciation, and the global freeze-over lasted for up to 30 My. In the absence of CO2 sinks such as silicate weathering, volcanic outgassing during glaciation raised atmospheric COe to 350 times the modem level. The ensuing extreme greenhouse conditions raised the mean global temperature to 40~ and abruptly (years to decades) ended the "snowball" state, leading to

5. 7. Paradox of Proterozoic Glaciomarine Deposition

453

the deposition of "cap" carbonates on glaciogenic deposits. Soluble ferrous iron that accumulated in anoxic seawater during the global freeze-over was oxidised by the atmosphere upon melting of the ice cover, causing the precipitation of iron-formations (section 5.6; cf. section 5.4) at the close of glaciation. Much evidence from geology, geochemistry and palaeontology refutes this scenario (see also section 5.6). 1. As discussed above, seas were unfrozen across wide areas and for lengthy time intervals during Neoproterozoic glaciations. These open seas were much more extensive and enduring than mere polynyas. The substantial geological evidence for extensive unfrozen seas during Neoproterozoic low-latitude glaciations is supported by a fully coupled ocean-atmosphere general circulation model (Poulsen et al., 2001). This study showed that even with a 5% reduction of the solar constant, low atmospheric CO2 (140 ppmv) and an idealised low-latitude continent, the sea-ice margin does not advance equatorwards of 46 ~ latitude. The large heat content of the ocean and ocean heat transport mechanisms inhibit a global freeze-over. 2. Thick (locally > 5000 m) Neoproterozoic glaciogenic successions on several continents contain glaciomarine tillites, mudstone-with-dropstone facies, fluvioglacial outwash deposits and striated pavements (e.g., Preiss, 1987; Yeo, 1981; McMechan, 2000; Condon et al., 2002). Such features indicate the operation of a vigorous and long-lived hydrological cycle. 3. Neoproterozoic glaciogenic successions commonly record numerous glacial advances and meltings or glacio-eustatic variations of sea level (Williams and Schmidt, 2000; Leather et al., 2002). Such high-frequency glacial cycles and eustatic oscillations conflict with the snowball Earth scenario, whereby each glaciation is marked by continuous severe cold, lasts for up to 30 My, and requires extraordinary circumstances for its initiation and termination. 4. CO2 sinks are not precluded during glacial intervals. CO2 sublimes at -78.5~ at atmospheric pressure, hence the mean surface temperatures as low as-110~ during a postulated snowball Earth episode would have caused atmospheric CO2 to precipitate and accumulate as the solid at high latitudes. Numerous Neoproterozoic continental regions were ice-free (Deynoux, 1982; Williams, 1986, 1998), so silicate weathering is possible. Furthermore, carbonate deposition is permissible in the widespread unfrozen Neoproterozoic seas and also may have occurred where the ice cover was thinMafter all, carbonates are forming below a permanent ice cover in saline lakes in Antarctica (Walter and Bauld, 1983). 5. Little seasonal variation occurs with global refrigeration because there is little precipitation or sublimation (Sellers, 1990). Hence large seasonal changes of temperature accompanying Proterozoic low-palaeolatitude glaciation (see above) appear in conflict both with a snowball Earth and a near-snowball or "slushball" Earth (see Crowley et al., 2001). Proponents of a snowball Earth accordingly have advanced other, albeit speculative, explanations for Neoproterozoic sand wedges. Hoffman (2001) suggested that glacial surge cycles cause the temperature changes required to produce such wedges.

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7.

8.

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

This idea is refuted by more than a century of observation and research on periglacial geomorphology and processes. Hoffman and Schrag (2002) argued that seasonal temperature changes would be enhanced globally for a snowball Earth because of the low heat capacity of the solid, ice-covered surface and weak lateral heat transfer, and made an invalid comparison with Mars, which is oceanless. This argument fails because of the compelling evidence that during Neoproterozoic glaciations the seas were unfrozen across wide areas and for lengthy time-intervals (see above) and many continental regions had patchy ice cover or were ice-free (Deynoux, 1982; Williams, 1998; McMechan, 2000). Maloof et al. (2002) claimed that diurnal temperature variations during a snowball Earth could produce sand-wedge polygons. It is highly dubious, however, that the likely mean diurnal temperature range of < 10~ for a 21.9-hour day (see section 5.9) in an equatorial coastal area near open seas, could produce cracks in brecciated quartzite that were sufficiently deep and wide and open long enough to permit the 3 m-deep Marinoan wedges of pebbly coarse sand (Fig. 5.7-2) to form. Positive 613Ccarb values for carbonates within and directly above Sturtian and Marinoan glaciogenic successions in Namibia (section 5.8), Australia and North America indicate a very normal, open ocean (Kennedy et al., 2001). Cap carbonates and their isotopic signature may be explained by ocean stratification followed by large-scale upwelling of deep waters that flooded continental shelves during deglaciation, where carbonates were precipitated in warm shallow seas (Grotzinger and Knoll, 1995). Furthermore, the record of 613Ccarb from c. 830-750 Ma for several continents shows numerous negative excursions (section 5.3) not associated with glaciogenic rocks, necessitating explanations other than a snowball Earth (Hill and Walter, 2000). The Rapitan iron-formation is succeeded by up to 600 m of tillite (Yeo, 1981) and the Sturtian iron-formation by 2700 m of mudstone containing dropstones and diamictite lenses (Drexel et al., 1993), indicating that glaciation continued long after iron-formation deposition had ceased. Geological and geochemical data indicate that Neoproterozoic iron-formations were deposited in rift-related, hydrothermally influenced basins (Yeo, 1981; Young, 1988; Breitkopf, 1988; Neale, 1993; Trompette et al., 1998) (see also sections 5.4, 5.6 and 5.8). Volcanism was associated with iron-formation deposition in Canada (Yeo, 1981) and Namibia (Breitkopf, 1988) and with Sturtian glacial deposition in South Australia (Preiss, 1987), which supports the hydrothermal interpretation. Significantly, the Sturtian-Rapitan iron-formations formed during Rodinian breakup (Powell et al., 1994) (see also section 3.10) or an episode of rifting related to that breakup. Marinoan glaciation did not cause any major change in the nature of Vendian acritarch populations in Australia (Grey, 2001, 2002), with pre- and post-glacial populations being almost identical. Grey (2001, p. 46) concluded that "The biotic evidence does not support the more extreme conditions predicted by a Snowball Earth". Argillites without cap carbonates conformably overlie the c. 2.3 Ga glaciogenic Gowganda and Ramsay Lake Formations in Canada (Young, 1991b) (section 5.6). Young and Nesbitt (1999) postulated a gradual transition to warmer conditions following the Gowganda glaciation.

5.7. Paradox of Proterozoic Glaciomarine Deposition

455

The above observations together argue that the snowball Earth hypothesis cannot account for the nature of Proterozoic glaciations. Other explanations must be considered.

Large non-dipole components of the Proterozoic geomagnetic field? Palaeomagnetism is based on the premise that the geomagnetic field has approximated a geocentric axial dipole (GAD) during geological history. However, recent studies suggesting the presence of significant non-dipole components of the geomagnetic field during the Proterozoic, Palaeozoic and Mesozoic (Kent and Smethurst, 1998; Van der Voo and Torsvik, 2001) (see also section 3.11) bring into question the assumed concordance of palaeomagnetic and true latitudes. Pleistocene periglacial sand wedges occur in North America at latitude 42.5~ (Wayne, 1990), so a discrepancy of c. 30 ~ between palaeomagnetic and true latitudes may resolve the paradox of a strongly seasonal periglacial climate near the Marinoan palaeoequator. Extremely large axial non-dipole components would be required to produce such a discrepancy (Schmidt, 2001). For example, an axial quadrupole component of 30% would offset equatorial palaeolatitudes by only 13 ~ However, evidence favours axial octupole fields in the Palaeozoic and Proterozoic (Kent and Smethurst, 1998; Van der Voo and Torsvik, 2001). Axial octupole fields would have no effect at equatorial (and polar) latitudes but a 50% octupole component would make true mid-latitudes appear to be palaeoequatorial. Such extreme non-dipole components have not been identified in the palaeomagnetic record, however, either by gross statistical analysis of Proterozoic data (Kent and Smethurst, 1998) or in data for the 1.2 Ga Mackenzie dyke swarm in Canada (Schmidt, 1999). Those two studies suggested an octupole component no greater than 30%. Schmidt (2001) concluded that "low latitudes alleged for some Proterozoic glaciations remain a first-order geological and geophysical paradox".

A Proterozoic large obliquity ? Obliquity and glacial climate. Planetary obliquity or axial tilt controls the seasonal cycle and climatic zonation, and cold climate may occur in low latitudes for a large obliquity. Mars has no large moon to stabilise its obliquity, which varies chaotically between about 0 ~ and 60 ~ on a c. 107-year time scale (Laskar and Robutel, 1993). At times of Martian large obliquity, water may sublime from the polar ice caps during the summer and move to low latitudes, where it may condense from the atmosphere and fill the top of the regolith with large amounts of ice (Jakosky et al., 1995). The obliquity of the ecliptic (the Earth's axial tilt, e) shows only small variation of +1.3 ~ around a mean of 23.3 ~ because of the moderate mean value and the stabilising influence of the Moon (Laskar et al., 1993) (see also section 5.9). A large obliquity would not be a primary cause of glaciation but could control the distribution and nature of glacial environments (Williams, 1975, 1993). 1. The ratio of solar radiation received annually at either pole to that received at the equator increases with increase in obliquity (Fig. 5.7-3). For e > 54 ~ latitudes < 40 ~ receive less radiation annually than high latitudes, and modelling with 70 ~ obliquity and 5% reduction of the solar constant showed that snow forms in low latitudes (Oglesby and

Chapter 5: Evolution of the Hydrosphere and Atmosphere

456

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Fig. 5.7-3. (a) Relation between the obliquity of the ecliptic (s) and the ratio of annual insolation at either pole to that at the equator. Points plotted for e -- 23.5 ~ (the present value), 54 ~ and 90 ~ (b) Latitudinal variation of relative mean annual insolation of a planet for various values of obliquity. The curves show that for s > 54 ~ glaciation would occur preferentially in moderate to equatorial latitudes. From Williams (1975, 1993).

Ogg, 1998). Because of the high albedo of snow, some may survive the equinoxes and form permanent ice in low latitudes. In high latitudes, by contrast, the cold, arid winter atmosphere would produce limited snow which would melt entirely during the very hot summer solstice. 2. The amplitude of the global seasonal cycle increases with increase in obliquity. For a large obliquity, high latitudes would endure greatly contrasting seasons and large seasonal changes of temperature would reach low latitudes. 3. Zonal surface winds such as the tropical easterlies and mid-latitude westerlies reverse for s > 54 ~ as the circulation in "Hadley cells" reverses (Hunt, 1982). In addition, strong surface winds would flow from the winter to the summer hemisphere. 4. Climatic zonation is relatively weak for e > 54 ~ Hence latitude-dependent climates would be unstable and any Milankovitch-band fluctuations in insolation may cause

5. 7. Paradox of Proterozoic Glacfomarine Deposition

457

abrupt and extensive climate changes and the stratigraphic juxtaposing of cold- and warm-climate deposits. Features of Neoproterozoic glaciations that are consistent with a large obliquity include: 9 glaciomarine deposition preferentially in low palaeolatitudes; 9 large seasonal changes of temperature in coastal areas, together with extensive and longlasting open seas and ice-free continental regions, near the palaeoequator; 9 palaeo-northwesterly winds in low palaeolatitudes directed obliquely towards or across the palaeoequator during Marinoan glaciation in South Australia (Williams, 1998); 9 the association of glaciogenic deposits and carbonates of apparent warm-water origin.

Obliquity origin and secular change. The large-obliquity hypothesis for low-latitude glaciation must include mechanisms to give an obliquity > 54 ~ during the pre-Ediacarian Proterozoic and then reduce it over an interval of c. 130 My to << 54 ~ prior to Late Ordovician polar glaciation. The widely accepted "single giant impact" hypothesis for the Moon's origin (Taylor, 1987; Canup and Asphaug, 2001) (see also Nelson, section 1.2, and Williams, section 5.9, for discussion of the origin of the Moon and the history of its orbit) requires a glancing impact with the proto-Earth and an impactor: target-planet mass ratio of ~>0.1 (but not an impactor with high orbital inclination), which would likely result in an obliquity of c, 70 ~ or larger for the Earth (Fig. 5.7-4). Dones and Tremaine (1993) confirmed that the obliquity of a planet resulting from impacts by a few large bodies "is likely to be substantial". A large early obliquity could have varied chaotically between 60--90 ~ (see Laskar et al., 1993). Modelling of the Archaean climate is consistent with a large early obliquity (Jenkins, 2000, 2001). Hence the long-held belief that the early Earth's obliquity was small (Goldreich, 1966) must be queried. The mechanism of "climate friction" for secular change of obliquity, identified by NASA celestial mechanicists in the 1990s, involves changes in the mass distribution of the Earth and Mars and the rate of spin-axis precession, with feedback loop, due to the waxing and waning of ice sheets (Bills, 1994; Rubincam, 1995, 1999; Ito et al., 1995). In principle, either secular increase or decrease of the Earth's obliquity may result, depending on the rate of ice-sheet variation and the rate of solid-earth deformation, which is controlled by mantle viscosity. D.M. Williams et al. (1998) investigated this mechanism as a possible cause of the postulated obliquity-decrease since the Neoproterozoic. While their study is inconclusive, the mechanism of climate friction is still in an early stage of exploration, As stated by Rubincam et al. (1998), "Unlike tidal friction, a topic two centuries old, climate friction is such a new field that we can't yet determine its importance for changing our planet's tilt. Only time will tell". Future Directions Both the snowball Earth and large obliquity hypotheses assume the Proterozoic geomagnetic field approximated a geocentric axial dipole (GAD). Available data do not contradict that assumption but much remains to be done in this area. More palaeomagnetic and

Chapter 5: Evolution of the Hydrosphere and Atmosphere

458

80

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geochronological studies of Proterozoic glaciogenic rocks also are required. If future work were to show that few if any continents occupied high palaeolatitudes during Neoproterozoic glaciations, the evidence for open seas and strong seasonality near the palaeoequator still may rule in favour of a large obliquity. On the other hand, demonstration of a reliable polar palaeolatitude for a pre-Ediacarian glaciomarine succession may rule out a large obliquity but would not confirm a snowball Earth or a slushball Earth because the paradox of open seas and strong seasonality would remain unresolvedma non-GAD may be preferred. Nor would the demonstration of synchronous glaciation in low palaeolatitudes favour a snowball Earth, because synchronous glaciation is possible with a large obliquity if glacial intervals saw little plate-motion (see also discussion by Lindsay and Brasier, section 5.3, on plate tectonic stasis).

5.8. Neoproterozoic Sedimentation Rates and Timing of Glaciations

459

Snowball, Slushball, non-GAD or Big Tilt? None of these hypotheses has been established firmly. On current evidence, however, resolution of the Proterozoic climatic paradox of glaciomarine deposition, open seas and strong seasonality near the palaeoequator may come via geomagnetism or celestial mechanics.

5.8.

NEOPROTEROZOIC SEDIMENTATION RATES AND TIMING OF GLACIATIONS--A SOUTHERN AFRICAN PERSPECTIVE

H.E. FRIMMEL Introduction

The Neoproterozoic Era is marked by the breakup of a 1.0 Ga supercontinent, with Rodinia (sections 3.10, 3.11) being the generally favoured choice of configuration, and subsequent closure of a number of oceanic and intra-continental basins that led to the assembly of the next supercontinent, Gondwana. This tectonic cycle (see also sections 3.2 and 3.9) from one to another supercontinent is characterised by large variations in the organic carbon cycle (section 5.3) and by possibly the most dramatic climate changes Earth has experienced over the past two billion years (sections 5.6 and 5.7). The climate changes are reflected by glacial deposits between marine carbonate successions and spectacular variations of the C isotopic composition of seawater proxies (e.g., Hoffman, 1999b). A number of hypotheses have been put forward to explain Neoproterozoic glaciations, with the snowball Earth hypothesis having gained particular popularity recently (Hoffman et al., 1998b; for a critical assessment, however, see Young, section 5.6, and Williams, section 5.7). It is presumably not coincidental that the end of this era of intense fluctuations in meteoric environmental conditions saw one of the most important changes in the evolutionary paths of life forms (section 6.3). Huge temporal variations of certain tracers, such as C and Sr isotopic compositions, within the Neoproterozoic sedimentary record have been recognised from many parts of the world and have been explained by the interplay between tectonic processes, climate and atmospheric conditions (e.g., Kaufman and Knoll, 1995; Jacobsen and Kaufman, 1999) (sections 5.2 and 5.3). The relationship between different chemostratigraphic anomalies and tectonic processes, both in space and time, remains, however unresolved. This is largely due to confusion in the global correlation of chemostratigraphic anomalies, the interpretation of such anomalies, and to a general lack of reliable age data on tectonic and climatic events that may have caused the chemostratigraphic anomalies observed. Neoproterozoic successions are known from all continents and many of them bear remarkable similarities. They typically start with continental rift deposits and contain one or more glaciogenic diamictite horizons, which are, in places, overlain by cap carbonates. Based on four distinct negative 813C excursions (note interpretation by Lindsay and Brasier, section 5.3) that are recorded within the Neoproterozoic carbonates (e.g., Jacobsen and Kaufman, 1999), two major glacial epochs have been recognised and are often The Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

460

Chapter 5: Evolution of the Hydrosphere and Atmosphere

referred to by their corresponding deposits in Australia as the Sturtian and Marinoan (section 5.7). Two distinct glaciations are distinguished within the Sturtian epoch. A maximum age for the lower Sturtian glaciation is given by U-Pb zircon dates of 755 4- 18 Ma for a granitic clast within a diamictite of the Rapitan Group in northwestern Canada (Ross and Vi!leneuve, 1997) and 758 + 4 Ma for a felsic tuff beneath an older diamictite in the Otavi Group (Damara Supergroup) in northern Namibia (Hoffman et al., 1998a). A minimum age is provided by a Pb-Pb zircon date of 741 4- 6 Ma (Frimmel et al,, 1996c) and a U-Pb zircon date of 752 4- 6 Ma (Borg and Armstrong, 2002) from felsic volcanic rocks above an older diamictite in the Port Nolloth Group in southern Namibia. The age of the upper Sturtian glaciation is constrained by a U-Pb age of 723 -+- 3 Ma obtained on dykes and basalts in the Rapitan Group (Park, 1997) and by a U-Pb zircon date of 712 -1- 2 Ma from a volcanic intercalation within an older diamictite of the Abu Maharah Group (Huqf Supergroup) in Oman (Allen et al., 2002). High resolution age constraints for the Marinoan glaciation are lacking but estimates range between 560 and 590 Ma (Saylor et al., 1998). The fourth negative 313C excursion is possibly related to a separate glacial event (Moelv glaciation) around 560 Ma (e.g., Brasier and Shields, 2000). A further negative ~ 13C excursion at the Precambrian/Cambrian boundary, dated at 543 Ma (Grotzinger et al., 1995), has so far not been unequivocally linked with a glacial event. Where absolute age data are absent, chemostratigraphic data, specifically 613C, and lithofacies are commonly used for stratigraphic correlation. Because of the relatively short residence time of C in seawater (Kump, 1991), 13C/12C ratios are very susceptible to local environmental conditions, such as water depth, evaporation rate, and mixing with meteoric water, and are therefore not necessarily a reliable proxy of secular variations in global seawater composition (see also section 5.3). As a consequence, attempts to correlate globally various post-glacial cap carbonate successions has led to ambiguity and confusion. This is exemplified by the correlation of some of the best studied Neoproterozoic glacial deposits, the Port Askaig Tillite of the Dalradian Supergroup in Scotland with the twin Varanger ti!lites of Norway and their correlatives in northeast Svalbard and East Greenland. They, in turn, have been correlated by some workers with the Marinoan glaciation of south and central Australia and with the Ice Brook Formation (Rapitan Group) of Northwest Canada (e.g., Saylor et al., 1998); in contrast, others (Kennedy et al., 1998) have assigned the upper Varanger glacial deposits to the Marinoan and the lower Varanger glacial deposits to the Sturtian glaciations, respectively, or suggested an entirely Sturtian age for these deposits (Prave, 1999; Brasier and Shields, 2000) (section 5.6). Similar problems exist with the correlation of the younger glaciogenic diamictite beds in the Damara Supergroup, which have been interpreted, based on two independently acquired similar data sets, both as synSturtian (Hoffman et al., 1998a) and as syn-Marinoan (Kennedy et al., 1998). Many of the Neoproterozoic successions, or at least parts thereof, were laid down in intracontinental basins and thus their rock record does not provide reliable proxies of contemporaneous seawater composition. Only a few successions meet the criterion of having exchanged fully with the open ocean, and even fewer successions provide reliable constraints on the timing of sediment deposition. Geochronologically reasonably well constrained, at least partly marine successions that span the period from about 750-540 Ma

5. 8. Neoproterozoic Sedimentation Rates and Timing of Glaciations

461

are the Huqf Supergroup (Brasier et al., 2000) and the Dalradian Supergroup (Brasier and Shields, 2000). Comparable successions occur in the Gariep and Damara orogenic belts in southwestern Africa. In particular the Gariep belt has the advantage of containing rocks that were deposited both on a passive margin, which eventually evolved into a syn-orogenic foredeep, and in an oceanic basin with oceanic islands or an aseismic ridge (Frimmel et al., 2002). Although the individual lithostratigraphic units were variably deformed and metamorphosed at low grade during the Pan-African orogeny, original thicknesses could be estimated for most of these units (Frimmel, 2000a). In this contribution, recently published geochronological and chemostratigraphic data are summarised with the aim of assessing integrated sedimentation rates for the three main tectonic phases of deposition, i.e., continental rifting, drifting and syn-orogenic basin closure. The Neoproterozoic successions of the Pan-African Gariep belt will serve as a reference, but an attempt to correlate crucial units, especially glaciogenic deposits, with those in other belts of southern Africa will also be made. It will be shown that the drifting phase is characterised by a lack of sediment deposition and by unusual isotopic composition of marine sediments. This conclusion will form the basis for speculation as to the possible causes of particularly positive 613C excursions.

The Gariep Basin Due to intense folding, top-to-southeast thrusting and wrench faulting during the PanAfrican orogeny, all direct evidence of the original extent, depth and structure of the Gariep basin has been obliterated. Only its eastern margin is well defined in the form of inverted structures, along which the Neoproterozoic rocks were thrust against their basement. Two main tectonic units are distinguished within the belt, the continental Port Nolloth zone and the oceanic Marmora terrane (Fig. 5.8-1). The Port Nolloth zone is a strongly shortened unit of predominantly siliciclastic rift to shelf carbonate deposits with minor volcanic rocks. Although the contact with the underlying Kibaran basement of the 1.0 Ga Namaqua-Natal belt (section 3.10) is, in most places tectonic, the Port Nolloth zone as a whole is regarded as para-autochthonous. The shape of the Port Nolloth zone most likely reflects the outline of the eastern basin margin (Fig. 5.8-1). The presence of two grabens, separated by a basement high, is inferred for the rifting phase. While the eastern rift graben (Rosh Pinah graben) failed, a half graben to the west evolved into an oceanic basin as recorded by the predominantly mafic rocks in the Marmora terrane (Fig. 5.8-2a). This terrane lacks basement inliers and in the lower part of its stratigraphy is devoid of significant continental siliciclastic deposits, but contains a relatively high proportion of mafic volcanic rocks. Geochemically, most of these rocks are comparable with oceanic island basalt, with some of them carrying a mid-ocean-ridge signature (Frimmel et al., 1996b). Based on a conservative estimate of east-west shortening and a width of 150 km of compressed oceanic rocks, estimated from aeromagnetic data (Corner, 2000), a minimum width of 1000 km is inferred for that part of the oceanic basin that is reflected by the Marmora terrane (Frimmel et al., 2002).

462

Chapter 5: Evolution of the Hydrosphere and Atmosphere

Fig. 5.8-1. (a) Distribution of Neoproterozoic tectonic belts in southern Africa and adjacent parts of South America. (b) Tectonic subdivision of the Gariep Belt (from Frimmel et al., 2002); 1-3--localities of profiles shown in Figs. 5.8-3a--c, respectively.

The lithostratigraphic subdivision of the c. 770-550 Ma Gariep Supergroup, likely depositional environments for individual units and available age controls have been reviewed by Frimmel (2000a) and Frimmel et al. (2002) and are summarised in Figure 5.8-2b. Three megasequences (M1-M3) are distinguished within the Port Nolloth zone: M1 comprises c. 770-740 Ma siliciclastic continental rift deposits, M2 c. 740-580 Ma passive margin

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8

464

Chapter 5: Evolution of the Hydrosphere and Atmosphere

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deposits that were laid down under highly variable eustatic conditions, and M3 c. 5 8 0 550 Ma syn-orogenic foredeep deposits (Frimmel and Frilling, in prep.). Onset of erosion into the corresponding foreland basin (Nama basin) overlaps with the age of metamorphism and continent-collision related deformation around 545 Ma (Grotzinger et al., 1995; Frimmel and Frank, 1998).

5.8. Neoproterozoic Sedimentation Rates and Timing of Glaciations

465

Chemostrati g raphy A series of chemostratigraphic profiles through carbonate-bearing sections, showing large variations in near-primary ~13C (relative to the PDB standard), 87Sr[86Sr, and in some cases even ~ 180 (relative to V-SMOW standard), has been established (Frimmel and Jiang, 2001; F611ing and Frimmel, 2002; Frimmel and F611ing, in prep.); a typical section is shown in Fig. 5.8-2a. Towards the upper Pickelhaube and Dabie River Formations, 313C increases to +9%0, only to drop again close to the top of the carbonate succession. The 87Sr]86Sri ratios are generally low and decrease up-section from around 0.7076-0.7071. Intense silicification and ferruginisation of the dolomite in the Gais and Sholtzberg Members in the Marmora terrane, probably caused by fumarolic degassing of the underlying volcanoes, makes chemostratigraphic comparison with similar stromatolitic carbonates in the Port Nolloth zone problematic, because of the lack of good, undisturbed profiles. Only a few samples were found to be fairly unaltered and they yielded 613C values of as much as +2.82%0. An isotopic profile through the type section of the Bloeddrif Member reveals a relatively uniform distribution of near-primary 613Ccarb, ranging between - 0 . 8 2 and + 1.01 %0 across the whole unit, except for the bottom 20 m of limestone above the Numees Formation diamictite, where the 313Ccarb values drop to as low as -4.63%0 (Fig. 5.8-3b). The marked depletion in 13C near the bottom of the succession has been explained by a strongly reduced input of organic carbon into the coeval seawater, as can be expected for cap carbonates above glacial deposits (see, however, sections 5.6 and 5.7). The extent of post-depositional alteration appears minimal as inferred from very low Mn/Sr and Fe/Sr ratios, and three distinct negative ~180 anomalies in the lower part of the succession might therefore be primary signals that reflect freshwater mixing with seawater during global melting of glaciers (F611ing and Frimmel, 2002). The 87Sr]86Srl ratios obtained for the whole succession are very consistent around 0.70824, except for the bottom 25 m, from where ratios of approximately 0.70852 have been obtained (F611ing and Frimmel, 2002). These higher ratios near the bottom also support an increased flux of melt water from the continents into the ocean, because river water and glacier ice draining older continental crust are enriched in radiogenic 87Sr relative to seawater. A profile through a comparable succession of the Bloeddrif Member (Frimmel and F611ing, in prep.) from a more proximal setting (Fig. 5.8-3c) differs lithologically by the partial dolomitisation of the limestone and the intercalation of thin sandstone beds. The isotope profile is similar to the previous one, except for a third negative 3180 outlier at 80 m stratigraphic height above the underlying Numees Formation diamictite, which might be an artefact of post-depositional alteration. Most importantly, the 613C values of the respective upper parts of the two sections differ significantly, with the second, more proximally positioned section showing a strong increase in 313C up-section (Figs. 5.8-3b, c). This discrepancy highlights the problems with chemostratigraphic correlation, that is a strong dependence of 613C on palaeodepositional environmental conditions. The stratigraphic equivalent to the Bloeddrif Member in the Marmora terrane (Dreimaster Member), is very condensed but displays a comparable C isotope trend but overall lower 6180 values, which is probably a result of post-depositional modification (Fig. 5.8-3a).

466

Chapter 5: Evolution of the Hydrosphere and Atmosphere

Fig. 5.8-3. Variations in 6180 and ~13C with stratigraphic height above lower contact with glaciogenic diamictite in foredeep carbonate deposits with increasing proximity to the continent: (a) section through a thin dolomite cover above mafic oceanic island basalt in the Marmora terrane at 27~ 15~ (Frimmel, 2000b); (b) section through a relatively distal succession in the Port Nolloth zone at 28~ 16~ (from F611ing and Frimmel, 2002); (c) section through a more proximal succession in the Port Nolloth zone at 29~ I~S, 17~ (Frimmel and F611ing, in prep.).

Correlation of Glacial Deposits Most attempts to correlate Neoproterozoic glacial deposits have relied heavily on chemostratigraphy, specifically on the interpretation of negative and positive 613Gearb excursions as indicators of global icehouse and greenhouse conditions, respectively (compare with Lindsay and Brasier, section 5.3). Some workers regard strong positive 613Ccarb excursions (in excess of 6%o) as indicative of a post-Sturtian and pre-Varangian/Marinoan age (e.g., Jacobsen and Kaufman, 1999), whereas others inferred an age between the older and younger Sturtian glaciation for carbonate successions that are characterised by strong enrichment in 13C (Hoffman et al., 1998b) (see also section 5.6). Relatively uniform 313Ccarb values between - 2 and +2%o are considered typical of post-Varangian/Marinoan carbon-

5. 8. Neoproterozoic Sedimentation Rates and Timing of Glaciations

467

ate rocks (e.g., Jacobsen and Kaufman, 1999). The strong susceptibility of 613Ccarb tO syndepositional environmental conditions, as exemplified by Figure 5.8-3, casts doubt, however, on the reliability of this particular indicator. Strontium isotopic composition seems more reliable for chemostratigraphic correlation, because of the longer residence time of Sr in seawater compared to that of C (Kaufman and Knoll, 1995). The existing data base suggests indeed that Sr isotope ratios permit a clearer distinction between pre-Marinoan and post-Marinoan marine carbonates, with the former being characterised by lower 87Sr/86Sr around 0.7070 and the latter by higher 87Sr/S6Sr around 0.7085 (Jacobsen and Kaufman, 1999). Based on such comparisons, the carbonate rocks of the Hilda Subgroup in the Port Nolloth zone, which display variable, but distinct positive 613C excursions and relatively low 87Sr]86Srl, have been compared with pre-Marinoan successions elsewhere, whereas the post-Numees carbonate rocks have been assigned to the post-Marinoan stage (F611ing and Frimmel, 2002). Correlation with the lower Sturtian glacial event is indicated for the older glaciogenic deposits of the Kaigas Formation, whose age is constrained between 771 + 6 and 741 -t- 6 Ma. A syn-Sturtian age is also supported by a Pb-Pb age of 728 + 32 Ma obtained on the cap carbonate above the Kaigas Formation, which has been interpreted as dating early diagenesis (F611ing et al., 2000). In contrast, the age of the younger glaciogenic deposits of the Numees Formation is less well constrained. The distinct negative 613C excursion in the lower part of the overlying Bloeddrif Member carbonates points to onset of carbonate sedimentation at the end of a glacial period. As the lower Bloeddrif Member bears all the hallmarks of a typical post-glacial cap carbonate, a significant hiatus between these carbonates and the, in places, conformably underlying Numees Formation diamictite is precluded. The latter should therefore be not much older than 583 Ma (555 + 28 Ma). Such an age would correspond well with the age estimated for the Marinoan glaciation elsewhere. In the light of the relatively large error that is invariably attached to a Pb-Pb carbonate age, the given geochronological constraint does not exclude the possibility of a correlation with the younger Moelv glaciation (around 560 Ma) elsewhere. Considering the glacial influence recorded in the upper Dernburg Formation of the Marmora terrane and oceanic crust formation there having lasted until around 600 Ma (Frimmel and F611ing, in prep.), a correlation between the Numees Formation on the continental margin and the oceanic Chameis Gate Member appears likely. Inversion from extension to contraction in the Gariep Basin at c. 575 Ma (Frimmel and Frank, 1998) coincided roughly with the end of the Numees glaciation. This agrees with a correlation with the Marinoan glaciation. Recovery to warmer climate might have continued until at least about 550 Ma, when a stormand wave-dominated carbonate ramp (Kuibis Subgroup) developed during sea level rise in the Nama basin. Regional correlation of the syn-Sturtian and syn-Marinoan glacial deposits, as recognised in the Gariep belt, across the Pan-African orogenic belts of southern Africa has been controversial. The proposed correlation (Fig. 5.8-4) is based largely on lithostratigraphy and trends in 613C, with absolute age control being fairly poor. For the Damara belt, the lithostratigraphic subdivision as proposed by Hoffmann (1989) and revised by Hoffmann and Prave (1996) was adopted here. In most sections, two glaciogenic units can be distin-

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3 Fig. 5.8-4. Stratigraphic subdivision and correlation of glaciogenic deposits (shown in italics) between the Damara,.Gariep and Saldania belts. Lithostratigraphic subdivision of the Damara belt after Hoffmann (1989; various tectonic subunits of the central and southern Damara belt are not distinguished) and Hoffmann and Prave (1996), that of the Gariep Belt after Frimmel (2000a) and Gresse (1992), and that of the Kango inlier after Le Roux and Gresse (1983). Absolute age data (U-Pb single zircon1Pb-Pb carbonate) from the following sources: I-(Hoffman et al., 1996), 2-(Hoffman et al., 1998a), 3-(Frimmel et al., 2001), 4--(Frimmel et al., 1996c), 5-(Folling et al., 2000), &(Grotzinger et al., 1995), 7--(Barnett et al., 1997), 8-Frimmel and Folling (in prep.).

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5.8. Neoproterozoic Sedimentation Rates and Timing of Glaciations

469

guished. In the Northern Platform of the Damara belt and in the Port Nolloth zone of the Gariep belt, geochronological data, supported by chemostratigraphic evidence, attest to a syn-Sturtian age for the older of these two glacial units, These older glaciogenic deposits were laid down towards the end of continental rifting, which took place more or less at the same time throughout the entire region. They are, therefore, not laterally continuous but confined to those areas along rift basin margins where mountain glaciers fed into the sea. Due to the lack of Ediacara fossils and to ~13C ratios above +8%0, Hoffman et al. (1998a) assumed a pre-Vendian/Marinoan age for the entire Otavi Group in the Northern Platform of the Damara belt. This would imply a Sturtian age also for the younger glaciogenic unit there (Ghaub Formation), with the Abenab Subgroup reflecting an interglacial period between the two Sturtian glaciations. The evidence is, however, equivocal. The post-Numees carbonate succession in the Gariep belt also lacks any Ediacara fossils and the presence or absence of stromatolitic bioherms (section 6.5) or algal mats (sections 7.9 and 7.10) are not considered reliable stratigraphic indicators but rather a reflection of certain depositional environments. An up-section trend from negative ~13C to ratios in excess of -+-4%0, and back to lower values, has been documented both from the post-Chuos carbonates of the Abenab Subgroup in the eastern (Kaufman et al., 1991" Frimmel et al., 1996a) and the western parts (Hoffman et al., 1998b) of the Northern Platform of the Damara belt, as well as from the post-Kaigas Hilda Subgroup in the Gariep belt (Frlling and Frimmel, 2002). Throughout the entire region, the isotope profiles above the younger diamictite units of the Ghaub and Numees Formations and the Chameis Gate Member start with a post-glacial negative ~13C excursion and continue with values around 0%0 (Kaufman et al., 1991; Frimmel et al., 1996a; Hoffman et al., 1998b; Fig. 5.8-3). Only in the uppermost Otavi Group (Htittenberg Formation) is a strong positive ~13C excursion recorded, reaching ratios as high as 10.7%0 (Kaufman et al., 1991; Frimmel et al., 1996a). That part of the isotope curve might be missing elsewhere, because siliciclastic flysch deposition took place in most parts of the Gariep and Damara depositories when the shallow marine platform carbonates of the Htittenberg Formation were laid down. The trend to higher ~13C ratios observed in the most proximal section through the Bloeddrif Member (Fig. 5.8-4c) might be a subdued equivalent to the strong positive ~13C excursion in the Htittenberg Formation. Bearing in mind that the various basins were closing at that time, they became progressively more restricted, which could account for the observed anomalous enrichment in 13C. Evidence for this exists in the form of evaporite beds within the Htittenberg Formation (Chetty and Frimmel, 2000). Correlation of the younger diamictite in the Northem Platform, and by analogy also in the other tectonostratigraphic units, with the syn-Marinoan Numees Formation (Fig. 5.8-4) is preferred therefore. In the central Damara Belt, only one diamictite is recognised at the base of the Swakop Group and it is tentatively correlated with the older glaciogenic unit elsewhere (Fig. 5.8-4). The absence of a syn-Numees (Marinoan) diamictite and corresponding cap carbonate succession there might be explained by that section having been already involved in contractional deformation (thrusting) at about 600 Ma, as it contains the fill of the central part of the Damaran basin (Khomas Sea) and represents the most intensely deformed and metamorphosed part of the orogen. The width of that basin is unknown, but the appearance of

470

Chapter 5: Evolution of the Hydrosphere and Atmosphere

syntectonic calc-alkaline granitoids around the same time has been interpreted as reflecting the beginning of subduction in the Damara belt (Miller, 1983). Correlation of the Damaran and Gariepian successions with those south of the Gariep belt is less clear. In the Vredendal outlier two diamictite beds occur and they are correlated, purely by analogy, with the Kaigas and Numees Formations of the Gariep belt. This is despite the fact that the intercalated Widouw Formation bears more isotopic similarities with the Bloeddrif Member than with the Hilda Subgroup (F611ing and Frimmel, 2002). To date no glacial deposits have been recorded from the Saldania belt. The C and Sr isotope trends and a Pb-Pb carbonate age obtained on the lower Kombuis Member of the Cango Caves Group are indistinguishable from those of the lower Bloeddrif Member, whereas an isotope profile for carbonates from the lower Matjes River Formation (Nooitgedacht Member) is similar to that from the Hilda Subgroup. A quartz arenite to arkose succession in the upper Matjes River Formation is possibly an equivalent to the Numees Formation (F611ing and Frimmel, 2002).

Neoproterozoic Sedimentation Rates The recognition of syn-Sturtian and a syn-Marinoan glaciogenic deposits in the Gariep belt, and by analogy also in the Damara and possibly in the Saldania belts, poses a problem with regard to sedimentation rates. The overall thickness of the Port Nolloth Group is estimated to be not more than 2 km (Frimmel, 2000a), which, taking into consideration the above age constraints, translates into an integrated sedimentation rate of 9 m My -l . This appears very low for the inferred depositional environments and types of sediment deposits (see also section 7.11, for a discussion of sedimentation rates in general), and it raises the question whether there are major gaps in the sediment record between about 770 and 550 Ma. The problem is not unique to southern Africa. A comparable situation exists in Oman, where the Huqf Supergroup with about the same total thickness covers roughly the same time interval (730-540 Ma; Brasier et al., 2000). Similarly, the Dalradian Supergroup represents the same time span as the Neoproterozoic successions in southern Africa, including the Mulden/Nama/Vanrhynsdorp Groups (Brasier and Shields, 2000). Continuous deposition of Dalradian sediments throughout this time interval has been rejected recently (Dempster et al., 2002). High sedimentation rates are assumed for the thickly bedded, coarse-grained alluvial fan and delta deposits of the Stinkfontein Subgroup and its equivalents in the Nosib Group. A minimum sedimentation rate of 27 m My-l (= 27 Bubnoff units, section 7.11) is derived from their thickness in the Port Nolloth zone and the above age constraints. Although the contact between the Stinkfontein Subgroup and the Kaigas Formation is unconformable in most places, reflecting deposition of the latter during rift-related extension and tilting of blocks along growth faults, examples of a conformable gradational contact between the two units exist (e.g., about 67 km east-southeast of Oranjemund at 29~ 17~ No major hiatus is envisaged for the boundary between M1 and M2 (Fig. 5.8-2) as the negative 613C excursion at the bottom of the post-Kaigas cap carbonates indicates onset of sedimentation during the waning stages of the glaciation. A carbonate accumulation

5. 8. Neoproterozoic Sedimentation Rates and Timing of Glaciations

471

rate of 52 m My-1 has been calculated for the Otavi Group (Hoffman et al., 1998a), and a similar value is probably applicable to the platform carbonates in the Hilda Subgroup and equivalents. A total thickness of 650 m for the Hilda Subgroup would correspond to a time span (12.5 My) that is less than the uncertainty on the age of these carbonates (728 -1-32 Ma). Yet, the overlying Numees Formation is some 150 My younger. A decrease from anomalously high 313C values in the carbonates immediately below the diamictite of the Numees and Ghaub Formations (Hoffman et al., 1998b; F611ing and Frimmel, 2002) is explained by worsening of the climate in preparation of the glaciation, thus precluding a major hiatus between the Hilda Subgroup and Numees Formation and between their equivalents in the Northern Platform, Damara belt. In the Port Nolloth zone, a major regressive event is recorded by the Wallekraal Formation (Fig. 8.5-2) and has been interpreted as reflecting eustatic sea level fall at the onset of the Numees glaciation (Frimmel et al., 2002). On the outer edge of the platform, in areas that were not affected by erosion in submarine channels breaking through the platform, carbonate deposition continued in the form of stromatolitic bioherms of the Dabie River Formation. There, however, the carbonate rocks of the Dabie River Formation rest apparently conformably on the mixed carbonate-mudstone-siltstone succession of the Pickelhaube Formation (F611ing and Frimmel, 2002). In contrast to M2, for which an integrated anomalously low sedimentation rate of less than 6 m My-I is indicated from the maximum thicknesses of the Hilda Subgroup (650 m) and the Numees Formation (500 m), higher sedimentation rates of more than 20 m My -l are indicated for the syn-orogenic post-Numees deposits, assuming a time span from around 580-550 Ma (M3). With progressive shrinking of the Gariep basin and the advance of the orogenic wedge from the west, sedimentation in the closing foredeep of the retro-foreland basin accelerated to as much as 150 m My -l during deposition of the lower Nama Group between 550 and 540 Ma, as derived from stratigraphic thickness and geochronological data (Saylor et al., 1998). Discussion and Conclusions

Whereas estimated integrated sedimentation rates for the continental rifting phase (M l) and syn-orogenic deposition in the foredeep (M3) compare well with those in modern analogues (see also section 7.11), a conspicuous lack of sediment for the interval between the syn-Sturtian and the syn-Marinoan glacial deposits is evident. The sediment present between the two glacial units should have been deposited, logically, over a few tens of millions of years, but the age difference between them is close to 150 My. The only potential gap that could account for this is at the base of the regressive Wallekraal Formation within the Hilda Subgroup, assuming that the contact between the Pickelhaube and Dabie River Formations is a paraconformity. If there was a regressive phase that lasted for some 100 My, it would have taken place during the opening of the Gariep basin (Adamastor ocean) between the Kalahari craton in Africa and the Rio de la Plata craton in South America. As an evolving passive continental margin should be progressively flooded during drifting, the speculated pause in sedimentation during Hilda Subgroup times would logically reflect

472

Chapter 5: Evolution of the Hydrosphere and Atmosphere

eustatic fall. A prolonged cold period with considerable icecaps at high latitudes but no ice cover at lower latitudes would thus be indicated for the 700-600 Ma interval (cf. section 5.7). Such an interpretation appears to be at variance with the greenhouse conditions generally inferred from the positive 613C excursion for this period. Some of the highest 13C ratios have been reported from the Zavkhan basin in western Mongolia (Brasier et al., 1996), the Olenek uplift in northeastern Siberia (Knoll et al., 1995a), the Yangtze platform in south China (Lambert et al., 1987), Oman (Burns and Matter, 1993), and northwestern Canada (Narbonne et al., 1994). In all of these areas, the depositional basins evolved from Neoproterozoic continental breakup and the possibility exists that none of these basins fully exchanged with the open ocean and/or that they were all relatively shallow at the time of deposition of the 13C-rich carbonates. If that is the case, no climatic implications can be deduced from the C isotopic record (see similar viewpoint, Lindsay and Brasier, section 5.3). An alternative to a 100 My cold period and low sea level stand can be found in the dominating tectonic regimes during the 700-600 Ma interval. The very low sedimentation rates could reflect correspondingly low subsidence rates. This would imply high spreading rates along mid-oceanic ridges. Fast spreading should be accompanied by intense hydrothermal activity due to high oceanic heat flux. Petrological evidence of extensive hydrothermal alteration during seafloor metamorphism at that time has been reported from the Marmora terrane (Frimmel and Hartnady, 1992). Independent evidence of enhanced hydrothermal input into the seawater during a time of little orogenic activity comes from low 87Sr/86Sr l ratios (< 0.7075) in carbonates of M2 in the Gariep belt and correlatable carbonate successions elsewhere (Jacobsen and Kaufman, 1999). The two alternative explanations for the lack of sediment deposition during the 700-600 Ma time interval are by no means mutually exclusive. Any change in the dominance of a given carbonic species dissolved in seawater will result in a change in ~13C, because of fractionation of C isotopes between CO2, HCO 3 and CO~- (see also section 5.2). Such a change could be effected by a change in pH. In the modem ocean, the pH of seawater (8.2) is buffered by its equilibrium with carbonates. As fast spreading of oceanic crust generally leads to a sea level rise, thus drowning possible pre-existing carbonate platforms, carbonate erosion rates would be low. In the absence of calcareous nanoplankton and thus pelagic carbonates in Neoproterozoic times, the buffering potential for the pH of seawater could have been exceeded by the high input of acidic hydrothermal fluids. Consequently, an increase in hydrothermal degassing on the ocean floor at a time of suppressed erosion of continental crust, as indicated by the low 87Sr/86SrI ratios, could potentially render seawater more acidic, thus favouring CO2 as dissolved carbonic species. This, in turn, would increase the 613C of dissolved HCO 3 by as much as 9%0 (Zeebe and Wolf-Gladrow, 2001), a difference that agrees well with the magnitude of the positive 613C excursions noted in post-Sturtian carbonate successions. It thus appears as if the sediment record and the trends in C as well as Sr isotopic compositions of seawater proxies for the post-Sturtian/pre-Marinoan period are controlled by the competing, and at times reinforcing responses to climatic, tectonic and hydrothermal processes, specifically the interplay between global dispersal of continents, lack of orogenic activity, largely peneplained land

5. 9. Earth's Precambrian Rotation and the Evolving Lunar Orbit

473

masses, intense hydrothermal degassing on the ocean floor and low sea levels related to cold climatic conditions (Lindsay and Brasier, section 5.3, have similar conclusions).

5.9.

EARTH'S PRECAMBRIAN ROTATION AND THE EVOLVING LUNAR ORBIT: IMPLICATIONS OF TIDAL RHYTHMITE DATA FOR PALAEOGEOPHYSICS

G.E. WILLIAMS

Introduction The theory of lunar tidal friction holds that the Moon's attraction on the Earth's tidal bulge causes the transfer of angular momentum from the Earth's spin to the lunar orbit, resulting in the slowing of the Earth's rotation and the recession of the Moon (Lambeck, 1980). The Sun's attraction makes a smaller contribution to the Earth's despinning (Brosche and Wiinsch, 1990). The implications of a backwards projection of the present rate of tidal energy dissipation are well known: around 1.5 Ga the Moon would have made a catastrophic close approach to the Earth. No evidence exists, however, for such a cataclysm. Early hopes that the Phanerozoic history of the Earth's rotation and the Moon's orbit could be determined from the study of skeletal growth increments in marine invertebrate fossils have not been fully realised (Lambeck, 1980; Williams, 2000). The Palaeozoic data, if taken at face value, imply little overall change in tidal friction in the past 500 My and an improbable close approach of the Moon between 1.5 and 2 Ga. Consequently, little work has been done on palaeontological "clocks" since the 1970s. Since the late 1980s the study of cyclic rhythmites of inferred tidal originmverticallyaccreted, laminated sediments that display periodic variations in lamina thickness reflecting a tidal influence on deposition (section 7.5)--has provided a new methodology for elucidating the history of the Earth's rotation and the Moon's orbit, that is applicable to the Precambrian. Here I point out features of tidal rhythmites and the cause of possible inaccuracies in rhythmite records, show how the validity of rhythmite data can be tested, and explore implications of the data for Precambrian length of day (LOD) and other aspects of palaeogeophysics (see also Nelson, section 1.2, for a review of the formation of Earth and the solar system).

Tidal Rhythmite Records Tidal rhythmites occur in modern marginal-marine deposits (Dalrymple and Makino, 1989; Tessier, 1993; Cowan et al., 1998) and are known from the Phanerozoic (Broadhurst, 1988; Kuecher et al., 1990; Miller and Eriksson, 1997) and Proterozoic (Williams, 1989a, 1991, 2000; Deynoux et al., 1993; Chan et al., 1994; Sonett and Chan, 1998). Such rhythmites form in a variety of settings including tidal flat, upper and deep-water estuary, glacial fjord, prodelta, delta plain and delta slope. The potential sediment load of tidal currents is related The Precambrian Earth: Temposand Events l-Xtited by P.G. l-riksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuncanu

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

directly to tidal range or height and consequent current speed. Hence sequential thickness measurements of laminae, usually of silt to fine sand grade, or groups of laminae deposited by tidal currents may provide a proxy tidal record that is suitable for time-series analysis. Tidal rhythmites (see section 7.5) can record semidiurnal, diurnal, fortnightly and longer periods that are ascribable to tidal pattern and type. Such rhythmites commonly show alternation of relatively thick and thin laminae (Fig. 5.9-1 a), recording the "diurnal inequality" of successive semidiurnal tides (de Boer et al., 1989). Failure to distinguish semidiurnal and diurnal laminae could result in misleading palaeotidal data. Tidal rhythmites also commonly show successive relatively thick and thin neap-spring (fortnightly) cycles (Fig. 5.9-1 c), reflecting the "monthly inequality" of spring-tidal ranges and current speed due to the elliptical lunar orbit. The monthly inequality varies from near-zero to a maximum and shows 180 ~ reversals of phase twice in just over a year, reflecting successive conjunctions and oppositions of the Moon and the Sun (Williams, 1991, 2000), and its apparent amplitude can be modified by sedimentary processes. Hence palaeogeophysical inferences (Eriksson and Simpson, 2000) should not be drawn from the amplitude of the monthly inequality shown by rhythmites. The accuracy of periods identified by spectral analysis of palaeotidal records depends on the length and continuity of the data (Williams, 1997, 2000). Sequential measurements of diumal or semidiurnal lamina thicknesses commonly provide an abbreviated palaeotidal record because of pauses in the deposition of laminae by tidal currents around times of neap tides, when only mud may have settled for several lunar days. Spectra for such short, incomplete or abbreviated records can have peaks shifted from true values. Any abbreviation of the raw data and consequent shifts of spectral peaks must be recognised if palaeotidal records are to be interpreted correctly. Hence, it is important to conduct tests where possible to determine the geophysical validity of claimed palaeotidal and palaeorotational values (see below). Precambrian Tidal Periods and Palaeorotation

Sequential thickness measurements of diurnal laminae and neap-spring cycles were obtained from ebb-tidal rhythmites in the Reynella Siltstone and coeval Elatina Formation (Figs. 5.9-1a, b) from the late Neoproterozoic Marinoan (Varanger) glaciogenic succession, South Australia (Williams, 1989a, 1997, 1998b, 2000). The Marinoan glaciation (see also section 5.8) has not been accurately dated, but c. 600 Ma is a fair estimate (Walter et al., 2000). The Elatina neap-spring cycles are abbreviated at positions of neap tides, and counts and spectral analysis of Elatina diurnal data do not yield meaningful palaeotidal periods. Hence the number of lunar days/synodic (lunar) month was obtained from counts of diurnal laminae in thick, evidently unabbreviated neap-spring cycles in the Reynella rhythmites. The rhythmites together record a wide range of palaeotidal and palaeorotational values (Table 5.9-1), including 13.1 synodic months/year and 29.5 lunar days/synodic month, which yield 400 solar days/year and a LOD of 21.9 hours at c. 600 Ma. The data indicate a dominantly synodic palaeotidal regime.

Table 5.9-1. Palaeotidal and palaeorotational values for Precambrian cyclic rhythmites and Modem values Parameter Lunar days/synodic month Solar days/synodic month Solar dayslsidereal month Synodic monthslyear Sidereal months/year Lunar apsides period (years) Lunar nodal period (years) Solar dayslyear Sidereal dayslyear Length of solar day (hours) Lunar semimajor axis ( R E ) Lunar recession rate (cm yr-' )

Age 2.45 Gaa (31.7 f 3.0) (32.7 f 3.0) (30.7 f 3.5) (1.7 h1.1 (16.7 f 1.1)

2.45 ~ a ' ( I . 1 f 1.5) (32.1 1 .5) (30.0 f 1.7) 14.5 f 0 . 5 * 15.5 f 0.5

c. 900 MaC (31.4 f 1.1) (32.4 f 1.1) (30.3 f 1.0) (14.3 f 0 . 6 ) (15.3 f 0.6)

c. 900 Mad (30.3) (31.3) (29.1) 13.5* 14.5

c. 750 MaC (29.9) (30.9) (28.7) 13.25* 14.25

23.3 f I S * (514f33) (515 f 33) (17.1 f I. 1 (51.9 k 3.3) (2.18 f0.86)

(21.6 f 0.7) (466f15) (467 k 15) (18.8 f0.6) (54.6 f 1.8) (1.47 f 0.46)

(21.5 f 0.3) 4644~13 (465 f 13) 18.9 k 0.5 54.7 f 0.7 (3.95 f 0.5)

(20.2) (422) (423) (20.8) (57.1) (2.25)

(19.8) (4 10) (41 1) (21.4) (57.8) (2.15)

*

c. 600 Maf 29.5 k 0.5* 30.5 f 0.5 (28.3 f 0.5) 13.1 f 0 . l " 14.1 f 0.1 9.7 f0.1" 19.5 f 0.5* 400f7 401f 7 21.9 f 0.4 (58.16 & 0.30) (2.24 & 0.31)

Modem 28.53 29.53 27.32 12.37 13.37 8.85 18.61 365.24 366.24 24.00 60.27 3.82 f 0.07

Tidal and rotational values vary with time, but the duration of the year is taken as constant (see text). The values in brackets were derived by the present author from respective primary or given values by applying equations (I), (2) and (3). as appropriate (see Williams, 2000). and make allowance for the solar tide's contribution to the loss of angular momentum of the Earth's rotation. Lunar recession rates are means from given ages to the present; the modem recession rate is obtained from lunar laser ranging (Dickey et al., 1994). "Cyclic banded iron-formation from the Weeli Wolli Formation, Western Australia, with the cyclicity viewed as annual bands grouped in palaeo-lunar nodal cycles (Walker and Zahnle, 1986). Error f l o for the primary value, determined by the present author from the data of Trendall (1973b). b ~ y c l i banded c iron-formation from the Weeli Wolli Formation, with the cyclicity viewed as fortnightly bands grouped in annual cycles (Williams, 2000). Error estimated for the primary value. 'Rhythmite from the Big Cottonwood Formation, Utah (Sonett and Chan, 1998). The given synodic period is 25.49 f 0.5 "present epoch days"; the primary neapspring period is unavailable. Errors f l u . d ~ h y t h m i t from e the Big Cottonwood Formation (Sonett et al., 1996a). Error for the primary value is unavailable. eRhythmite from the Pualco Tillite, South Australia. Error for the primary value is unavailable. f ~ h y t h m i t efrom s the Elatina Formation and Reynella Siltstone, South Australia (Williams, 1989a-c. 1991, 2000). The rhythmites also record semidiurnal and diurnal increments and fortnightly, semi-annual and annual periods. Errors f l a for primary values, except those for lunar dayslsynodic month and the lunar nodal period, which are estimates. 'Primary value determined directly from the rhythmite.

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

Fig. 5.9-1. Neoproterozoic tidal rhythmites, South Australia. Mudstone bands appear darker than sandy to silty laminae. Scale bar 1 cm, for all photographs. (a) The c. 600 Ma Reynella Siltstone, showing one fortnightly neap-spring cycle that contains c. 14 diurnal (lunar day) laminae of fine-grained sandstone; the thicker diurnal laminae have mudstone tops (arrows) and comprise two semidiurnal sublaminae. N marks mudstone bands deposited at neap tides. (b) The c. 600 Ma Elatina Formation, showing four fortnightly neap-spring cycles comprising graded diurnal laminae of very fine-grained sandstone and siltstone. N marks neap-tidal mudstone bands, where abbreviation of the cycles has occurred. (c) Rhythmite from the c. 750 Ma Pualco Tillite, showing 24 distal neap-spring cycles of finely laminated, very fine-grained sandstone that are bounded by thin mudstone bands deposited at neap tides. The specimen has split along one neap mudstone band (arrow) and a soft-sediment disruption occurs near the top.

The Elatina rhythmites also show strong cyclicity marking the annual oscillation of sea level (Fig. 5.9-2a), which has a mean period of 26.1-26.2 neap-spring cycles and is seen over the entire 60-year Elatina record. This non-tidal oscillation characterises modern sea-level records (e.g., Fig. 5.9-2c) and is driven by seasonal factors (Pattullo, 1966; section 5.7). Well preserved tidal rhythmites (Fig. 5.9-1 c) interbedded with diamictites in the Pualco Tillite of the c. 750 Ma Sturtian glaciogenic succession in central South Australia likewise record the annual oscillation of sea level (Fig. 5.9-2b). The Pualco Tillite rhyth-

5. 9. Earth's Precambrian Rotation and the Evolving Lunar Orbit

477

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Fig. 5.9-2. Neap-spring cycle thicknesses for (a) part of the c. 600 Ma, 60-year-long Elatina rhythmite record obtained from drill cores, and (b) the c. 750 Ma Pualco Tillite rhythmites (cycle number increases up the stratigraphic successions). The smoothed curves (5-point filter weighted 1, 4, 6, 4, 1) show annual peaks in the non-tidal annual oscillation of sea level that define the solar year. (c) Maximum height of the synodic fortnightly tidal cycle at Townsville, Queensland, from 19 October 1968 to 3 June 1970. The smoothed curve (5-point weighted filter) shows the non-tidal annual oscillation of sea level. The second-order peaks in the three curves reflect the semi-annual tidal cycle of the Sun's declination, which raises its highest spring tides around times of the equinoxes. Tidal data for Townsville provided by the Beach Protection Authority, Queensland Department of Transport.

Chapter 5: Evolution of the Hydrosphere and Atmosphere

478

mites contain 26.5 neap-spring cycles in an annual oscillation, which gives 13.25 synodic months/year and implies 410 solar days/year and a LOD of 21.4 hours at c. 750 Ma (Table 5.9-1). Rhythmites from the c. 900-Ma Big Cottonwood Formation in Utah have yielded inconsistent results (Table 5.9-1), and the Big Cottonwood data based on Sonett and Chan(1998) conflict with other Neoproterozoic data in Table 5.9-1. Big Cottonwood neap-spring cycles contain as few as 4-5 sandstone laminae and have thick mudstone bands at positions of neap tides (Sonett et al., 1996a), indicating that the rhythmites are strongly abbreviated. Palaeoproterozoic iron-formations evidently are the product of submarine hydrothermal activity (Barley et al., 1997), which may be modulated by such tidal processes as variations in earth tides and advection by tidal currents (Fujioka et al., 1997; Kinoshita et al., 1998). Table 5.9-1 gives palaeotidal values based on two different tidal interpretations of cyclic banding in iron-formation from the 2.45-Ga Weeli Wolli Formation in Western Australia (see also section 5.4).

Validity of Tidal Rhythmite Data Because tidal rhythmite records may be rendered inaccurate by abbreviation, the accuracy of derived data should be tested objectively where possible. The validity of palaeotidal and palaeorotational values in the same data set can be assessed by testing for self-consistency through application of the laws of celestial mechanics (Deubner, 1990; Williams, 1990, 1997, 2000). This procedure is demonstrated here for the Elatina-Reynella data set by using independent and widely separated primary values, i.e. values determined directly from the rhythmites, to calculate the lunar semimajor axis (mean Earth-Moon distance) at c. 600 Ma. The period of the lunar nodal cycle and the dimension of the lunar semimajor axis are related in the expression P - P0 (cos io/cos i)(ao/a)1.5,

(1)

where P is the past lunar nodal period, P0 is the present lunar nodal period of 18.61 years, a is the past lunar semimajor axis, a0 is the present semimajor axis of 60.27 Earth radii (RE), i0 is the inclination of the lunar orbit to the ecliptic of 5.15 ~ and i is the past lunar inclination (Walker and Zahnle, 1986). Assuming negligible change in lunar inclination i since the late Neoproterozoic, a lunar nodal period of 19.5 4- 0.5 years (Table 5.9-1) gives a/ao = 0.969 + 0.017 at c. 600 Ma (Williams, 1989a). Kepler's third law states that the square of the orbital period of a planet is proportional to the cube of the planet's mean distance from the Sun. For the Earth-Moon system, it follows that the square of the lunar monthly period is proportional to the cube of the lunar semimajor axis. Hence (2)

5. 9. Earth's Precambrian Rotation and the Evolving Lunar Orbit

479

where Ts is the length of the sidereal month in the past and To is the present length of the sidereal month. Employing values from Table 5.9-1, Ts/To = 13.37/(14.1 + 0.1), giving a/ao = 0.965 + 0.005 at c. 600 Ma. The loss of the Earth's rotational angular momentum resulting from tidal friction of the Moon and the Sun, and the concomitant increase in lunar orbital angular momentum, can be expressed as

1 w 1.219

4.93w0

-

(~00) 1/2 +

(0.46)2 (~00) 13/2 13 '

(3)

where w0 and w are the Earth's present and past rotation rates, and the present ratio of the Moon's orbital angular momentum to the Earth's spin angular momentum = 4.93 (Deubner, 1990). From Table 5.9-1, w / w o - (401 + 7)/366.24, which gives a/ao --0.968 + 0.007 at c. 600 Ma. The close agreement among these three independent determinations of the lunar semimajor axis, which make allowance both for lunar and solar tidal effects, demonstrates the self-consistency of the Elatina-Reynella data set and supports the validity of the implied LOD of 21.9 + 0.4 hours at c. 600 Ma (Table 5.9-1). The value a/ao -- 0.965 + 0.005, obtained from equation (2) using the best-constrained Elatina datum of 14.1 -1- 0.1 sidereal months/year, gives a mean rate of lunar recession of 2.24 • 0.32 cm yr -! since 600 Ma (the present lunar semimajor axis -- 384,400 km). By comparison, 14.25 sidereal months/year implied by the Sturtian tidal datum indicates a/ao - 0.958, giving a mean recession rate of 2.15 cm yr-1 since 750 Ma. This accordance of independently determined mean lunar recession rates holds for different estimated ages of Marinoan and Sturtian glaciations: taking a Marinoan age of 590 Ma gives a recession rate of 2.28 cm yr -l and a Sturtian age of 720 Ma gives 2.24 cm yr - l . Although the Sturtian datum is not validated, the accordance of these Neoproterozoic results provides further support for the veracity of the Elatina-Reynella data set. The validity of palaeotidal periods for c. 900 Ma (Big Cottonwood Formation" Sonett et al., 1996a; Sonett et al., 1996b" Sonett and Chan, 1998), 2.45 Ga (Weeli Wolli Formation: Walker and Zahnle, 1986; Williams, 1989c, 2000) and 3.2 Ga (Moodies Group: Eriksson and Simpson, 2000) (section 7.5) cannot be assessed by the above method because respective data sets do not contain two or more independent and widely separated primary values. Such unverified palaeotidal and palaeorotational values should be viewed with caution.

A Test of Earth Expansion Palaeotidal and palaeorotational data can be used to explore whether the Earth's moment of inertia has changed over geological time. Such analysis also can examine whether the Earth's radius has increased significantly with time, as required by the hypothesis of Earth expansion, because the Earth's moment of inertia would increase with secular increase in radius.

480

Chapter 5: Evolution of the Hydrosphereand Atmosphere

Employing the method of Runcorn (1964, 1966) and assuming no change in the universal gravitational parameter G, the Elatina-Reynella data set (Table 5.9-1 ) indicates that I/Io = 1 . 0 0 6 - 1.014 (+0.018), where I0 and I are the Earth's present and past moments of inertia (Williams, 1998b, 2000). These figures are the only direct estimates of I/Io using Precambrian data and argue against significant overall change in the Earth's moment of inertia since c. 600 Ma. They rule out hypotheses of Earth expansion since that time by endogenous mechanisms, including rapid expansion since the Palaeozoic (Carey, 1976), which requires I/Io = 0.5, and slow expansion during the Phanerozoic (Creer, 1965; Egyed, 1969), which gives I/Io = 0 . 8 9 - 0.94. The suggestion that substantial Earth expansion may have resulted from change of minerals in the Earth's interior to their less dense phases caused by a postulated secular decrease in G (Carey, 1976) is not supported by studies of the morphologies of other terrestrial planets, which show little or no evidence of expansion (Crossley and Stevens, 1976; McElhinny et al., 1978). Moreover, Mars Viking Lander and lunar laser ranging data indicate negligible change in planetary orbital radii, implying negligible change in the length of the year and in G (Hellings et al., 1983; Chandler et al., 1993; Dickey et al., 1994). The rhythmite data and astronomical observations therefore argue against significant change in the Earth's radius by any mechanism at least since the late Neoproterozoic. A Despinning Earth and Global Plume Events As tidal friction slowed the Earth's rotation over the eons, the Earth's forced nutations (periodic tipping) under the action of lunar and solar torques must have undergone resonances with the free nutation of the fluid core at annual, semiannual and higher frequencies (Toomre, 1974; Hinderer and Legros, 1988). Estimating the times of core resonance in the geological past requires accurate information on the Earth's past LOD. Using the ElatinaReynella datum for past LOD, which is the most accurate available (see above), Williams (1994b) suggested that the critical LOD for annual resonance most likely occurred during the late Neoproterozoic or early Palaeozoic. Other important resonances probably occurred during the Archaean, but their timing is clouded by the lack of reliable palaeorotational data for that eon. Greff-Lefftz and Legros (1999b) confirmed the occurrence of core resonances during Earth history. However, their postulated times of resonance in the late Palaeozoic and Precambrian are not meaningful because they are based on a discredited palaeorotational model, namely a mean rate of despin for the Phanerozoic that equates with the present astronomical value (see above; Lambeck, 1980) and which is nearly twice the mean deceleration rate implied by the Elatina-Reynella datum (Williams, 1989a, b, 2000). Resonances of the fluid core may have led to instability and heat release near the coremantle boundary (Hinderer and Legros, 1988). Williams (1994b) postulated that widespread tectonothermal reworking events in late Neoproterozoic-early Palaeozoic times may have resulted from plumes generated at the core-mantle boundary (sections 3.2-3.4) in response to the annual resonance of the fluid core. The possible correlation of earlier plume-generated tectonothermal events with important resonances of the fluid core may be

5. 9. Earth's Precambrian Rotation and the Evolving Lunar Orbit

481

tested by the acquisition of high quality, validated palaeorotational data for the Palaeoproterozoic and Archaean.

History of the Moon's Orbit The mean rate of lunar recession of 2.24 + 0.32 cm yr -1 since 600 Ma indicated by the Elatina datum is only c. 59% of the present rate of lunar recession of 3.82 +0.07 cm yr-1 obtained by lunar laser ranging (Dickey et al., 1994). The observed high rate of lunar recession may reflect the near-resonance of oceanic free modes and tidal frequencies (Lambeck, 1980; Stindermann, 1982). The mean Earth-Moon distance aT at an earlier time T is approximated by

aT--ao

I

13 __(h~ 2 ao

)2/13 ,

(4)

where a0 is the present mean Earth-Moon distance and (h0) is the mean rate of lunar recession (Walker and Zahnle, 1986). Projecting into the past a rate of tidal energy dissipation consistent with the observed rate of lunar recession indicates a close approach of the Moon at c. 1.5 Ga (Fig. 5.9-3, curve a). A tidal dissipation rate suggested by Phanerozoic palaeontological data (Lambeck, 1980) implies a close approach at c. 1.9 Ga (Fig. 5.9-3, curve b), as does the Big Cottonwood datum of Sonett and Chan (1998). Employing an average rate of tidal energy dissipation consistent with the Elatina datum (a/ao = 0.965) pushes back a possible close approach of the Moon to 3 Ga (Fig. 5.9-3, curve c). However, the geological records of the Earth and the Moon provide no evidence of such a cataclysm, and an even lower rate of tidal dissipation seems required for times prior to the Neoproterozoic. Additional validated palaeotidal data for an epoch before 600 Ma would permit the early history of the Moon's orbit to be traced. Following Walker and Zahnle (1986), the mean lunar distance prior to the Neoproterozoic is approximated by

aT=al

13 ( h i ) ) 2/13 1---z-(T-TI)z

(5)

al

where a l is the mean Earth-Moon distance at time Tl (600 Ma) and (hi) represents a dissipation rate consistent with the Elatina datum and a datum for an earlier epoch. A dissipation rate, albeit not validated, inferred from the study of cyclic banding in iron-formation from the 2.45 Ga Weeli Wolli Formation (Table 5.9-1; Williams, 2000) can be used with the Elatina datum to demonstrate the tracing of the lunar orbit prior to 3 Ga (Fig. 5.9-3, curve d). According to this scenario, a close approach of the Moon did not occur during geological history. The findings, while unverified, are consistent with other tidal histories (e.g., Hansen, 1982; Webb, 1982). The uncertainties in available palaeotidal data for the Palaeoproterozoic and Archaean (section 7.5) are too great to permit a reasonable estimate of the mean Earth-Moon distance at 4.5 Ga. Validated palaeotidal data that demonstrably are internally self-consistent and in

Chapter 5: Evolution of the Hydrosphere and Atmosphere

482

Fig. 5.9-3. Change in mean Earth-Moon distance with time, as suggested by different average rates of tidal energy dissipation. Curve a: employing a dissipation rate consistent with the present rate of lunar recession (Dickey et al., 1994), in equation (4). Curve b: employing a dissipation rate consistent with a mean rate of lunar recession since c. 500 Ma suggested by palaeontological data (Lambeck, 1980), in equation (4). The Big Cottonwood datum is from Sonett and Chan (1998). Curve c: employing a dissipation rate consistent with the Elatina datum, in equation (4). Curve d: employing a dissipation rate consistent with the Elatina datum, and the Weeli Wolli datum where the cyclicity is viewed as fortnightly bands grouped in annual cycles (Table 5.9-1), in equation (5); shaded area shows the error for the Weeli Wolli datum. Reproduced from Williams (2000) with the permission of the American Geophysical Union.

accordance with the laws of celestial mechanics are required for those eons to allow the early LOD and lunar orbit to be determined at a high level of confidence.

5.10.

ANCIENT CLIMATIC AND TECTONIC SETTINGS INFERRED FROM PALAEOSOLS D E V E L O P E D ON IGNEOUS ROCKS

H.W. NESBITT AND G.M. YOUNG

Introduction Weathering (see also section 5.11) of igneous rocks is strongly influenced by climate, primarily temperature and humidity. Past climatic conditions may be evaluated from the The Precambrian Earth: Temposand Events l-Mited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. MueIler and O. Catuneanu

5.10. Ancient Climatic and Tectonic Settings

483

composition of weathering profiles, palaeoprofiles (palaeosols), sediments and sedimentary rocks, thus providing insight into the history of the Earth which cannot be obtained easily by other means. The physical and compositional state of a weathering profile is determined primarily by two competing influences, chemical weathering (driven by climatic conditions) and mechanical erosion (driven by tectonism and eustasy). By careful petrographic and geochemical documentation of the state of a palaeosol, past climatic and tectonic (or eustatic) conditions affecting it can be evaluated. The approach relies primarily upon evaluation of the rate of chemical weathering of source rocks (Nesbitt and Young, 1984) relative to the rate of mechanical erosion of weathered materials from profiles. The weathering of igneous rocks is simplest to treat and is the focus of this section, but metamorphic and sedimentary rocks may be treated in a similar fashion. The approach proposed is not meant to be exhaustive with regard to the range (or detail) of approaches available, but is a summary of the bulk compositional-mineralogical methodology proposed by the authors; the literature cited is likewise restricted.

Overview of Chemical Weathering of lgneous Rocks Chemical weathering of alkali and alkaline earth Al-bearing silicates may be viewed as an acid-base reaction where the acids are those of the natural aqueous environment and the bases are the primary alkali and alkaline earth Al-silicate minerals. The acids include carbonic acid (most abundant), natural organic, sulphuric and hydrochloric acid whereas the bases are the feldspars (most abundant), phyllosilicates, amphiboles, pyroxenes, orthosilicates and others. As emphasised by Garrels (1967), the reaction of these acids and bases produces dissolved salts (predominantly dilute Na- and Ca-bicarbonate solutions observed in soils, shallow springs, streams and rivers) and A1-Si-bearing residues (mainly of alkaliand alkaline earth-depleted clay minerals). Of the primary minerals, feldspars are by far the most abundant, representing about 60-70% of the labile (readily weathered) silicate minerals of the upper continental crust (Wedepohl, 1969; Nesbitt and Young, 1984). Quartz and some minor and trace minerals are much less labile than feldspars, hence generally much less susceptible to chemical weathering. Amphiboles, micas and pyroxenes constitute less than 20% of the upper crust (Nesbitt and Young, 1984). Study of the weathering of common plutonic rocks is therefore largely concerned with weathering of feldspars. Consideration of weathering of natural silicate glasses expands the study to include all common igneous rocks.

Compositions of Common Igneous Rocks and Their Minerals The preponderance of feldspars in igneous rocks greatly simplifies the treatment of weathering. Silica is present in all the common minerals of igneous rocks and is therefore an insensitive indicator of changes resulting from weathering. Compositional differences

484

Chapter 5: Evolution of the Hydrosphere and Atmosphere

among the feldspars relate primarily to the relative abundances of alkalis, alkaline earths and aluminium, as well as transition elements if mafic minerals are of concern. The common minerals of igneous rocks can be portrayed in an A1203-CaO + Na20-K20 plot (A-CN-K triangle) using molar proportions (Fig. 5.10-1a, crosses). These include plagioclase (P1.), K-feldspar (K-Sp.), aluminous biotite (A1-Bi.), hornblende (Hb.) and diopside (Di.). Augite (Au) plots between Di. and Hb. with its location determined by its A1 content (bar on the CN-A boundary of Fig. 5.10-1 a represents typical range of composition). Average igneous rock compositions (Nockolds, 1954) are plotted on Figures 5.10-1 a, b. Granites (sensu stricto) contain subequal amounts of quartz, plagioclase and K-feldspar. Quartz abundances are immaterial to this treatment because SiO2 is not a component of the ternary diagram. True granites plot close to the feldspar join (dashed line joining plagioclase and K-feldspar in Fig. 5.10-1 a) at the position indicating about equal amounts of each feldspar. Granodiorites plot near the feldspar join but closer to the plagioclase composition, indicating about three quarters plagioclase and one quarter K-feldspar in average granodiorite. Quartz diorites plot very close to the plagioclase (P1.) composition indicating that they contain little K-feldspar. Average diorites and gabbros plot close to the plagioclase composition and below the feldspar join, indicating abundant plagioclase and appreciable amounts of mafic minerals (e.g., hornblende, pyroxene). All major primary minerals of the igneous rocks plot close to, or below, the feldspar join (Fig. 5.10-1a). The common clay minerals smectite (Sm.), illite, kaolinite (Ka.), gibbsite (Gib.) and chlorite (Cht.) are aluminous and in contrast to the primary minerals, plot well above the join. As a result of the compositional separation of primary minerals and their secondary weathering products on the A-CN-K diagram, weathering of granitic rocks should evolve up the diagram, as primary minerals are destroyed and secondary minerals are produced. Samples from weathering profiles, and shales derived therefrom should plot much higher on the diagram than the granitic rock averages. The proportions of feldspars to various mafic minerals in the igneous rocks may be evaluated from a "mafics triangle" (Fig. 5.10-1 b, constructed by plotting molar proportions of A1203-CaO + Na20 + K20-Fe203(T) + MgO, the A-CNK-FM triangle). In addition to the minerals plotted on Fig. 5.10-la, the oxides (Ox.), orthopyroxenes (Opx.) and calcite (Cc.) are included in Fig. 5.10-lb. The diagram may be used to determine the proportion of kaolinite (or gibbsite) and chlorite in profiles and in siliciclastics rocks (note that these cannot be distinguished in Fig. 5.10-1a because both minerals plot at the A apex). Weathered samples from profiles and siliciclastic sediments generally plot above the line joining feldspars (Fel.) and orthopyroxene in Fig. 5.10-1b due to presence of clay minerals in these materials (e.g., shales or sandstone matrix). Fresh igneous rocks plot close to or below the line joining feldspars and orthopyroxenes, indicating the abundance of primary minerals. True granites are highly feldspathic and consequently plot close to the feldspar composition (Fel.), whereas gabbros/basalts plot closer to the mafic mineral compositions, indicating that these rocks contain a higher proportion of melanocratic minerals than do granites.

485

5.10. Ancient Climatic and Tectonic Settings

100i

,~

A

(a)

Gib.,Cht.

i ~Granodiorite ~AI.Bi.

60 -

.~

50 40

" Hb.

CN

10 20 30 40 50 60 70 80 90 M ole P erce nt

, ~ A~ , Gib.,

(b)

Gran h o :b ro Co/ as:t' CNK

10

20

30

40 50 60 70 M ole P e rce nt

80

90

FM

Fig. 5.10-1. Illustrates the relationship between the average bulk chemical compositions of some common igneous rocks and the compositions of primary minerals of these rocks, and secondary minerals produced by weathering of these rocks. (a) An A - C N - K diagram where A = A1203, CN = CaO* + Na20, and K -- K 2 0 (molar basis). The asterisk indicates CaO associated with silicates (see Fedo et al., 1995). (b) An A - C N K - F M diagram where A = A1203, CNK = CaO* + N a 2 0 + K20, and FM = FeO(total) + MgO. Dots represent plutonic rocks and circles their volcanic equivalents; crosses represent common minerals. Ka. = kaolinite, Gib. = gibbsite, Cht. = chlorite, Sm. = smectite, Ms. = muscovite, AI-Bi. = aluminous biotite, P1. = plagioclase, K-sp. = K-feldspar, Hb. = hornblende, Au. = augite, Di. = diopside, Ox. = Fe-oxides, Cc. = calcite, Opx. = orthopyroxene.

486

Chapter 5: Evolution of the Hydrosphere and Atmosphere

Chemical Weathering of Igneous Rocks As emphasised by Nesbitt and Young (1984), the residues of weathered feldspathic rocks define distinctive compositional trends on the A-CN-K and A-CNK-FM triangles (Figs. 5.10-la, b), due primarily to the preferential leaching of alkalis and alkaline earths (relative to A1) from the feldspars (Fig. 5.10-2). Furthermore, these trends are readily and accurately predicted for common igneous rocks in that all trends are effectively parallel to the A-CN boundary of Figure 5.10-2a, and parallel to the A-CNK boundary of Figure 5.10-2b. The extent to which the parental material has been weathered can be quantified by using either the A1203 content of Figure 5.10-2a, or the Chemical Index of Alteration (CIA; see Nesbitt and Young, 1982; Fedo et al., 1995). CIA values equal the A1203 values of Figures 5.10-1a and 5.10-2a if the components are plotted as percentages rather than fractions (as indicated by the vertical scale on the left side of these diagrams). Chemical weathering trends parallel to the A-CN boundary inevitably approach the A-K boundary, indicating that effectively all Na and Ca have been removed from the materials plotting at the boundary, and also indicating that effectively all plagioclase has been weathered to clay minerals. K-feldspar remaining in these weathered materials continues to be leached, and in the process K is removed from the profile in preference to A1203. There is a consequent abrupt change in weathering trend along the A-K boundary, towards the A apex (Fig. 5.10-2a, short arrow; compositions of the residues lose K-feldspar and migrate towards the A apex). There are, as a result, two segments to the resulting weathering trend of a granitic rock, the two trends coincident at the A-K boundary. The arguments are equally applicable to the weathering of other igneous rocks and weathering trends can be constructed in like manner. The compositions of residues change systematically with their compositional trace parallel to the A-CN boundary, and then towards the A apex. As consequence, the CIA value (or the A1203 percentage of Fig. 5.10-2a) is a direct and sensitive indicator of the amount of weathering suffered by any bulk sample collected from a weathering profile.

Mineralogical Zones in Profiles The proportion of Na + K + Ca to AI in the smectitic and illitic clay mineral groups is intermediate between those of the parental igneous rocks and the kaolinite or gibbsite mineral groups, so that the smectites and illites plot at intermediate values of CIA (or A1203%) on Figures 5.10-1 a and 5.10-2a. The smectites and illites are therefore produced in abundance during the initial and intermediate stages of weathering (Meunier et al., 1976; Eggleton, 1986) whereas kaolinite and ultimately gibbsite will be produced in abundance during the advanced weathering stage (Altschuler et al., 1963). As a result of weathering, mature profiles exhibit mineralogical zonation (Fig. 5.10-3a), with the complex clay minerals (e.g., smectites, illites, vermiculites) most abundant in zones intermediate between fresh parental rock and the upper, extensively weathered zone where kaolinite group minerals (and ulti-

487

5.10. Ancient Climatic and Tectonic Settings

(a)

A

100I90 80

~,Ka.,Gib.,Cht. Sm.

~

Illite Ms.

70

9

~,~

4o

9

.

.

.

.

E

.

.

.

.

.

.

.

.

.

.

-

~

n

3O

.

.

.

.

2O 10 C N

10

20

30

40

50

60

70

80

90

Mole Percent

A ,

(b)

"~ite,~ Granites~/

theringtrends

Fel:~,-,.,..~

Ox.+ px.

C CNK

10

20

30

40

50

60

70

80

90

FM

M ole Pe rcent Fig. 5.10-2. Illustrates the compositional trends defined by the residues of the various igneous rocks during their weathering. Symbols and abbreviations are as in Fig. 5.10-1.

mately gibbsite) may dominate (Craig and Loughnan, 1964; Nesbitt and Wilson, 1992; Nesbitt and Markovics, 1997). Thus, zonations in both bulk chemical and mineralogical compositions are observed in weathering profiles.

488

Chapter 5: Evolution of the Hydrosphere and Atmosphere

Fig. 5.10-3. Hypothetical, idealised weathering profile developed on feldspathic granitic rock where chemical weathering is intense and mechanical erosion is minimal. (a) Changes to the intensity of shading corresponds to intensity of chemical weathering (greatest intensity at the top of the profile). The major mineralogical changes are identified beside the diagram. Controls on Rates of Chemical Weathering and Mechanical Erosion

As previously noted, chemical weathering may be viewed as an acid-base reaction whereby acids introduced to bedrock react with the constituent minerals to produce dissolved salts and secondary minerals. The rate of introduction of acids to profiles is determined, in turn, by annual rainfall and rate of production and breakdown of litter (e.g., leaves and grasses in the later Phanerozoic). Briefly, bacterial breakdown of litter to CO2 and to low molecular weight organic acids, and their subsequent acquisition by rainwater percolating through O horizons of soils, introduce the acids to the weathering profile. The rate of chemical weathering therefore is controlled primarily by climatic conditions, principally annual rainfall and temperature (climate as in Fig. 5.10-3b). The CO2 content of the atmosphere may have varied substantially over geological time (sections 5.2 and 5.11). Increased CO2 partial pressure of the atmosphere results in increased acidity in rainwater, thus promoting mineral weathering by increasing the acid content of each volume of soil water contacting solid bases (e.g., feldspars).

5.10. Ancient Climatic and Tectonic Settings

489

Fig. 5.10-3 (continued). (b) A (thermometer-like) diagram illustrating the relative intensities of chemical weathering and mechanical erosion required to produce a profile similar to that shown in (a). Climate and tectonism are the ultimate forces affecting the profile, with the former acting through chemical weathering and the latter through mechanical erosion.

Tectonic activity typically produces relief thus promoting mechanical erosion of disaggregated portions of profiles (through runoff). Similarly rise or fall of sea level promotes erosion. The rate of mechanical erosion is therefore controlled primarily by tectonism and eustasy (Fig. 5.10-3b), although climate also has an effect in that a medium (e.g., air, water, ice) usually is required to sustain mass wasting of weathering profiles (see also section 7.1). Evidence related to erosion includes truncation, or reworking of the upper portions of profiles. Careful field study of palaeosol development below an unconformity (or hiatus) may reveal variations in thickness related to palaeotopography and evidence for erosion of the original profile. Variable thickness of a palaeosol (or its absence) coupled with truncation of the secondary mineral assemblages of the palaeosol provides evidence for erosion. Modem weathering profiles develop distinctive textural features with a massive zone at the top, beneath which an incompetent zone with primary textures preserved is found (saprolite). Below these zones is the transition to fresh rock (Nesbitt and Markovics, 1997). Study of these properties may provide information on tectonism or eustasy and their rates relative to the rate of chemical weathering.

490

Chapter 5: Evolution of the Hydrosphere and Atmosphere

Because sediments represent detritus derived primarily from soils developed on source rocks, combined study of sedimentary rocks and palaeosols derived from the same source rocks may provide appreciably more insight into the relative rates of chemical weathering and erosion (tectonism/eustasy) affecting the source lands of the basin (Nesbitt and Young, 1997; Fedo et al., 1996).

Relative Rates of Chemical Weathering and Mechanical Erosion Where chemical weathering rates are much more rapid than rates of erosion, profiles are well developed, with the exposed surface being composed primarily of aluminous, secondary clay minerals and quartz (Figs. 5.10-3a, 5.10-4a). If mechanical erosion is commensurate with the rate of chemical weathering (Fig. 5.10-4b), the weathered mantle capping source rock will be less well developed and the exposed surface will be composed of both clay minerals (with complex clays, smectites, illites, vermiculites dominating kaolinite and gibbsite) and primary minerals. Where the rate of mechanical erosion greatly dominates over the rate of chemical weathering (Fig. 5.10-4c) a disaggregated weathering profile may be very thin or absent and primary bedrock may constitute much of the exposed surface. Any detritus derived from the exposed surfaces is composed mostly of primary minerals (with minor complex clays). Although derived from the same bedrock source, the composition of detritus derived from each of these hypothetical profiles would differ greatly, the differences being due solely to the relative rates of chemical weathering and mechanical erosion.

Example: Palaeosols beneath the Palaeoproterozoic Huronian Supergroup (2.2-2.4 Ga) The thickness of a palaeosol is an indicator of the difference in the rates of chemical weathering and mechanical erosion (Fig. 5.10-4). Palaeosols occur at several localities beneath the classical lower Palaeoproterozoic Huronian Supergroup (section 5.6) on the north shore of Lake Huron, Ontario, Canada, and one is discussed in some detail here. The Ville Marie palaeosol (Rainbird et al., 1990) was developed on a Late Archaean (Algoman) granite (c. 2.5 Ga) and is overlain by sandstone of the Huronian Lorrain Formation (older than 2.2 Ga). There is no distinctive A or B horizon developed, although these are unlikely to have developed without production of a substantial litter zone or O horizon, the development of which requires the presence of land plants. Without appreciable organic acids to extract and transport Fe (and A1) from the A horizon, and to deposit it in the B horizon (its rusty colour produced by formation of ferric oxyhydroxides), these horizons do not form. The Ville Marie profile is about 10 m thick where sampled. Although the palaeosurface displays appreciable relief (25 m elevation over about 200 m distance) there is no obvious evidence of erosion of the profile at the sample site. Its thickness (where sampled) and the preservation of primary textures suggest that chemical weathering dominated mechanical erosion during its formation and that little of the profile was eroded before deposition of the overlying sandstones.

5.10. Ancient Climatic and Tectonic Settings

491

Fig. 5.10-4. Hypothetical, idealised profiles with illustration of the relative intensities of chemical weathering and mechanical erosion required to produce a profile with the properties shown. The sequence (a)-(c) illustrates the effect of increasing intensity of mechanical erosion relative to chemical weathering.

Bulk chemical and mineralogical compositions indicate that the profile was mature (i.e., that chemical weathering dominated). As noted by Rainbird et al. (1990), kaolinite group minerals (kandites) and quartz are abundant whereas feldspars are scarce in the uppermost zone of the profile. The bulk chemical composition of the upper profile indicates K-metasomatism resulting in the replacement of kaolinite by sericite (discussed subsequently). The base of the profile is composed primarily of quartz and feldspars with minor quantities of clay minerals (now sericite). The proportion of feldspar decreases and the clay mineral content (now mostly sericite) increases from base to top of the profile, as expected with increased weathering intensity towards the top of the profile (Fig. 5.10-3). These considerations led Rainbird et al. (1990) and Fedo et al. (1995) to conclude that weathering of the Ville Marie granite during the early Proterozoic proceeded under a hu-

492

Chapter 5: Evolution of the Hydrosphere and Atmosphere

mid, warm temperate climatic regime akin to that of present-day conditions in the southern Appalachians; rainfall was appreciable and temperatures were moderately warm. There is evidence, however, that rainfall may have been seasonal, as discussed below. Although there is no evidence for removal of the top of the Ville Marie palaeosol, they commonly have been removed (e.g., during transgression) and care must be taken to evaluate the amount removed.

Early Diagenesis of Profiles and Climatic Implications The Murrurundi profile is developed on a basalt, located on the central east coast of Australia (New South Wales) where the climate is subtropical and seasonally humid. A perched water table located about 0.5-1.5 m below the surface is recharged during the humid season and is subject to evapotranspiration in the dry season, during which period carbonates precipitate and montmorillonite forms in the zone of the perched water table, producing bulk chemical and mineralogical weathering trends with anomalously high concentrations of SIO2, CaO, MgO and carbonate (measured as CO2) within the zone associated with the perched water table (Craig and Loughnan, 1964; Nesbitt and Young, 1989). Nesbitt and Young (1989) argued that the carbonate and montmorillonite accumulations were superimposed on the weathering profile and reflect an early diagenetic event. These diagenetic minerals would not be produced if the climate were continually humid, thus their presence in a laterally continuous but highly restricted vertical zone, where CIA values would be otherwise high, provides insight into the seasonal nature of the climate. Similar CaO, MgO, CO2 and SiO2 anomalies, and the presence of carbonates are observed in the early Proterozoic Ville Marie palaeosol (Rainbird et al., 1990), the Palaeozoic Pre-Fountain granodioritic palaeosol (Wahlstrom, 1948) and the early Proterozoic Hekpoort basaltic palaeosol (Button, 1973). These accumulations likely mark the position of the water table or a perched water table, and strongly suggest that the climate was distinctly seasonal (alternating humid and dry).

Metasomatism and Metamorphism of Profiles After burial, profiles are likely to be conduits through which ground waters and formation waters migrate, due primarily to their enhanced porosity and permeability (compared with the source rocks to the profiles). Profiles are also subjected to elevated temperatures as burial proceeds, thus clay minerals and partially degraded primary minerals are subject to reconstitution at elevated temperatures through acquisition of K, Na or Ca from subsurface waters (Nesbitt and Young, 1989). Kaolinite typically is altered to K-micas (sericite) and partially degraded feldspars may acquire K or Na to form K-feldspar or albite, depending on whether the waters are continental or marine in origin. Continental waters typically have lower Na/K values than sea water-derived formation waters (Nesbitt and Young, 1989). The Pre-Fountain, Palaeozoic granodioritic palaeosol (Wahlstrom, 1948), Hekpoort basaltic palaeosol (Button, 1973) and the Ville Marie palaeosol are examples where K-metasomatism of kaolinite and feldspars has occurred. The Butler Hill granitic

5.10. Ancient Climatic and Tectonic Settings

493

palaeosol provides an example where Na-metasomatism of feldspars has occurred (Blaxland, 1974; Nesbitt and Young, 1989). In this profile, the samples from the intermediate portion of the profile plot close to the feldspar join, but well to the sodic (CN) side of the predicted weathering trend (trend emanating from fresh parent rock and oriented parallel to the A-CN boundary) (Fig. 5.10-2a). Apparently, albite has formed from degraded feldspars during diagenesis, suggesting the involvement of a marine-derived solution. Mg-metasomatism may accompany K and Na metasomatism. Provided temperatures and solution compositions are appropriate, Mg may be acquired by clay minerals to produce montmorillonites and 7 E chlorites (bertherine). The basaltic Hekpoort palaeosol is an example where both K- and Mg-metasomatism occurred with burial and metamorphism. Samples analysed by Button (1973) and plotted on an A-CNK-FM diagram (Fig. 5.10-2b) plot on a straight line joining muscovite and chlorite, and span this entire compositional range (Nesbitt and Young, 1989). Those plotting closest to the muscovite composition were most weathered (contained the most kaolinite) and suffered the greatest amount of K-metasomatism. The other samples were subjected to variable amounts of both K- and Mg-metasomatism. Diagenetic, metamorphic and metasomatic reactions must be identified if climatic conditions are to be extracted from palaeosols. Conclusions

Chemical weathering of the most common igneous rocks involves primarily the destruction of primary labile minerals (mainly feldspars), the leaching of alkali and alkali earth elements from weathering profiles, and the production of secondary clay minerals within profiles. These processes can be quantified by calculation of a Chemical Index of Alteration and have predictable and consistent trends in A1203-CaO + Na20-K20 compositional space (ternary diagrams). The study of weathering profiles developed on igneous rocks provides a means of investigating the relative importance of chemical weathering and mechanical erosion affecting the profile, and hence provides a window into climate and tectonism or eustasy during formation of the profile. Palaeoclimatic conditions may be inferred from such observations as depth of profile development and intensity of mineral alteration within a profile. In some weathering profiles and palaeosols with impeded drainage, diagenetic development of carbonate- and smectite-rich (SiO2-rich) horizons may indicate seasonal variations in humidity. Many palaeosols show evidence of significant diagenetic-metasomatic alteration, particularly of K-, Na- and Mg-metasomatism. Comparison of modern and ancient profiles is, however, possible by using a combination of geochemical and mineralogical techniques which make allowance for diagenetic (metasomatic) changes (Fedo et al., 1995). Tectonic and eustatic processes lead to mechanical erosion of profiles and their effects on source lands may be deduced by careful field, geochemical and mineralogical study of both palaeosols and sedimentary rocks. Such type of study is one of the few avenues available to obtain evidence for palaeoclimatic and tectonic (or eustatic) conditions prevalent during the early evolution of the Earth. Although palaeoclimates and

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palaeo-tectonism/eustasy (section 7.1) may not be deduced in total from geochemicalmineralogical investigation of palaeosols, additional evidence bearing on these may be obtained in combination with stratigraphic and sedimentological studies of associated supracrustal sequences. The study of palaeosols may provide valuable information regarding climatic and tectonic conditions in the past. Most interpretations are based on comparison with modern profiles (the uniformitarian principle; section 7.1) but it should be kept in mind that presentday conditions such as the relative areas of continents and oceans, global temperature distribution and atmospheric composition may not provide a good analogue for those in the Precambrian (e.g., sections 5.2 and 3.6). Corcoran and Mueller (section 5.11) examine Archaean weathering and discuss several examples from the rock record.

5.11.

AGGRESSIVE ARCHAEAN WEATHERING

EL. CORCORAN AND W.U. MUELLER Introduction

Archaean sedimentary deposits and their source rocks were highly affected by weatheringaggressive climatic conditions. Increased levels of heat (Kasting, 1993), humidity (Des Marais, 1994b), and greenhouse gases, such as carbon dioxide (Young, 1991a; Kasting, 1.993) and methane (Pavlov et al., 2001b; Kasting and Siefert, 2002) inferred for the early Precambrian (section 5.2), led to intense chemical weathering of labile minerals and unstable rock fragments (Fig. 5.11-1). Clay minerals such as kaolinite, primarily produced from chemical alteration of feldspars (section 5.10), combined with smectites derived from erosion of abundant volcanic rocks, would have been the first products washed out of depositional systems, leaving the more stable minerals, such as quartz, to predominate the framework mineralogy. Notwithstanding, the phenomenon of quartz enrichment as a result of intense chemical weathering is not always evident in the rock record in cases where the source rocks were silica-poor. Determining the intensity of chemical weathering has been significantly facilitated since Nesbitt and Young (1982) introduced the concept of the CIA (Chemical Index of Alteration), a palaeoweathering index (section 5.10). Intense weathering of fine-grained sedimentary rocks should be indicated by high CIA values (> 85). Condie et al. (2001) averaged CIA values from shales of Late Archaean to Neoproterozoic age and found that there is a decrease in the CIA over time from 80 in the Archaean, to 75 in the Proterozoic, and finally to 70 in the Phanerozoic, further supporting increased chemical weathering during the Archaean. Evidence for a CO2- and CH4-rich reducing atmosphere can be found in high CIA values, quartz-rich deposits, and detrital heavy minerals that are consistent with low atmospheric oxygen levels (see, however, detailed discussions in section 5.2). The Precambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

5.11. Aggressive Archaean Weathering

495

ArchaeanAtmosphere breakdown of feldspars into clays= luartz enrichment, higherCIA values /High002 (between 100-1000 PALat 3.0 Ga)

increasedatmosphericheat

, increasedatmosphericheat I and acidityin meteoricwaters

High CH4 Probable low 02 higher evaporation rates resulting in greater rainfall

AI~ ~

High heat (up to 85 ~ . ~ No

'~ higher degreesof chemical weathering

no intakeof CO2, ,~.~higher levelsof run-off

promoted increasein CH4

!

greenhousegases= more heat increased atmosphericheat

under conditionsof intensechemicalweathering, the role of sediment residencetime would be significantlyreduced

Fig. 5.11-1. Flow chart illustrating the main atmospheric controls responsible for aggressive chemical weathering during the Archaean. Note how the interaction of high carbon dioxide and methane levels with low oxygen levels and no vegetation contributed to an increase in atmospheric heat and greater rainfall, which lead to high degrees of chemical weathering.

Review of the Archaean Atmosphere Oxygen (02) Kasting and Siefert (2002) stated that the Earth's atmosphere has been directly linked with the biota (e.g., section 6.2) on the evolving planet. For example, the rise in atmospheric 02 at approximately 2.3 Ga (cf., however, with section 5.2) has been attributed to the action of cyanobacteria (Holland, 1984; Farquhar et al., 2000a). A minor amount of oxygenproducing organisms during the Early Archaean may have contributed to an atmosphere in which oxygen was present in only trace quantities (Sleep, 2002). In addition, low oxygen levels have been attributed to volcanoes emanating CO, H2 and CH4, which effectively acted as a sink for the oxygen derived from photosynthetic bacteria (Kump et al., 2001; Sleep, 2002). Sulphur isotope signatures (sections 5.2 and 5.5) derived from sulphate and sulphide in Precambrian sedimentary rocks support photochemical oxidation of volcanogenic sulphur in the atmosphere as the cause of oceanic sulphate rather than oxidative weathering of continental sulphides (Farquhar et al., 2000a). Photochemical oxidation is inferred to have diminished due to increased oxygen production between 2350 and 2100 Ma, based on the change from pyritic to oxide-bearing conglomerates, decrease of banded iron-formation development, and deposition of red beds (Krupp et al., 1994) (see Ohmoto, section 5.2, for divergent arguments). The Archaean oxygen-poor state has also been inferred from chemical profiles of palaeosols (section 5.10), which show that the level of atmospheric 02 increased at about

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2.75-2.0 Ga (Rye and Holland, 1998). The presence of red beds immediately overlying some of these palaeosols reflects a significant increase in atmospheric oxygen at approximately 2.2 Ga, and is consistent with carbon isotope data (see also section 5.3) determined from 2.5-1.9 Ga rocks from the Fennoscandian Shield (Karhu and Holland, 1996). A study by Murakami et al. (2001) of a late Archaean to early Proterozoic palaeosol developed on an Archaean granite at Pronto Mine, Canada, also supports a relatively anoxic atmosphere at 2.6-2.45 Ga. The mineral rhabdophane, found in the weathering profile, contains high Ce concentrations with unoxidised Ce 3+, favouring an anoxic weathering solution. Rasmussen and Buick (1999) support a relatively anoxic Archaean atmosphere based on the abundance of siderite and siderite overgrowths in 3250-2750 Ma fluvial sandstones from the Pilbara Craton. These views of an oxygen-poor early Precambrian atmosphere have been based largely on chemical evidence from palaeosols and are refuted by Ohmoto (1996b, 1999) who maintains that reduced palaeosols can develop in an oxic atmosphere (see Ohmoto, section 5.2). Beukes et al. (2002) noted that the c. 2.2 Ga Hekpoort palaeosol of the Transvaal Supergroup, Kaapvaal craton, contains a lower iron-poor and an upper iron-rich zone, and maintained that several iron-poor palaeosols used to infer an ancient reduced atmosphere could represent only the lower parts of a complete weathering profile (see also section 5.10; interplay of weathering and erosion). Although the debate concerning the levels of atmospheric oxygen on the early Earth continues (summary in section 5.2), assessment of chemical weathering during the Archaean is also dependent on greenhouse gases such as CO2 and CH4, which are directly related to global warming.

Carbon dioxide (C02) A principal factor in the process of chemical weathering related to climate is atmospheric CO2. As suggested by Walker et al. (1981), abundant CO2 on the early Earth increased surface temperatures, and higher surface temperatures resulted in elevated evaporation of H20 from a primitive ocean, which in turn produced heavier rainfall on continents. High surface temperatures and precipitation are optimal climatic conditions for intense chemical weathering. The inference of 20-35% lower solar luminosity in the early Archaean with minor evidence of glaciations (sections 5.6 and 5.7), implies that there must have been significant quantities of one or more greenhouse gases in the early atmosphere (HendersonSellers, 1988). These would have included CO or CO2 based on the present abundance of carbon on the Earth and in chondrite meteorites (Tajika and Matsui, 1993). In the atmosphere, CO would have been oxidised to CO2 by OH radicals (Kasting and Ackerman, 1986), thus significantly increasing the amount of CO2 during the Archaean (see also section 5.2). According to Kasting (1993), the pCO2 of the 3.0 Ga atmosphere was approximately 100-1000 times PAL (present atmospheric level), resulting in more acidic surface waters than those of today (Ohmoto, 1999), which further promoted chemical weathering (see detailed discussion, section 5.2). Methane (CH4) Although CO2 played a pivotal role in warming the Earth's atmosphere during the Archaean, CH4 may have also contributed to an ancient greenhouse effect (Pavlov et al.,

5. I 1. Aggressive Archaean Weathering

497

2001 b). Methanogenic bacteria would have produced substantial levels of CH4 in an anoxic atmosphere (Kasting et al., 2001 ). The carbon isotope record (section 5.3) between 3.5 and 2.8 Ga is commensurate with a significant production of methane (Pavlov et al., 2001b). Rye and Holland (1996), using chemical analyses of basalt-derived palaeosols, determined that the amount of CO2 needed in the atmosphere to counterbalance the solar luminosity at 2.75 Ga would have been enhanced by the abundance of methane (see also Kasting and Siefert, 2002) (detailed discussion is provided by Ohmoto, section 5.2).

Effects of Chemical Weathering on the Composition of Archaean Sedimentary Rocks The chemical composition and mineral proportions of sedimentary rocks vary considerably when subjected to intense chemical weathering (Table 5.11-1). The primary control governing aggressive chemical alteration is climate (Johnsson, 1993) (section 5.10). A hot, humid, CO2- and CH4-rich atmosphere, in addition to an absence of vegetation, as are postulated for the Archaean, would have been extremely favourable for intense chemical weathering on the early Earth. Chemical alteration is at its optimum in humid climates where feldspars, micas, and rock fragments are preferentially broken down relative to more resistant minerals such as quartz and zircon (Potter, 1986; Fedo et al., 1995, 1997). This breakdown of labile minerals and rock fragments can occur at the sediment source, and/or during or following sediment dispersal. The intensity of chemical weathering is subsequently enhanced in low relief settings where sediment residence times are high, inhibiting rapid burial (Grantham and Velbel, 1988; Johnsson and Stallard, 1989). However, a lack of vegetation during the Archaean translates into shorter residence times, so that on average, the intensity of chemical weathering must have been significantly higher than today to produce similar deposits, as is supported by the geochemical data set of Condie et al. (2001). The effects of chemical weathering are best recorded in geochemical data such as the CIA (section 5.10) and PIA (plagioclase index of alteration: based on relative proportions of A1203, K20, CaO and Na20, where a value of 100 indicates that all primary plagioclase has been altered to clay minerals), specific oxide ratios, and in some cases, the behaviour of the rare earth elements (Fedo et al., 1996; Sugitani et al., 1996; Bhat and Gosh, 2001; Condie et al., 2001). Petrographically, the results of chemical alteration can be represented by quartz enrichment and low proportions of labile minerals, provided that parameters concerning relief and depositional setting are known (Corcoran et al., 1998; Donaldson and de Kemp, 1998; Corcoran and Mueller, 2002). In addition, the presence of specific heavy minerals, such as uraninite, in Archaean sedimentary sequences have been used to deduce chemical weathering conditions at the time of deposition (Maynard et al., 1.991) (see also section 5.2).

Examples of aggressive Archaean weathering Corcoran et al. (1998) showed how wave- and tide-influenced quartz arenites and quartzrich sandstones of the Keskarrah Formation, Slave Province, Northwest Territories, Canada, developed in a high-relief, tectonically-controlled setting primarily as a result

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Table 5.11-1. Potential effects of chemical weathering on the composition of Archaean sedimentary rocks Geochemical

Petrological

Enriched in K, Ba, Rb compared with average Archaean upper crust

Enriched in quartz, which need not be rounded

Depleted in Ca, Na, Sr compared with average Archaean upper crust

Low feldspar and rock fragment component

High A1203/Na20 ratios

Presence of detrital uraninite

Low A1203/TiO2 ratios

Detrital minerals enriched in Ti (e.g., rutile)

High concentration of HREE relative to LREE High CIA and PIA values (> 85)

of chemical weathering. The quartz-rich nature of the sandstones (Fig. 5.11-2) supports in situ weathering of a granodioritic source, although mafic volcanic rocks were also an important contributor of detritus, as shown by abundant mafic volcanic clasts in conglomerates (36% of total). Feldspars, which dominate the mineralogy of the volcanic sequence, were altered to clays in a weathering-aggressive environment and were transported through suspension into deeper water. The CO2-rich, humid Archaean atmosphere would have been favourable for the breakdown of labile feldspars, and further diminution probably occurred during hydraulic sorting and shoreline reworking. Because the Keskarrah Formation developed in a high relief setting, as indicated by up to 4 m sized boulders along remnant basin margins, the rate of chemical alteration would have minimised the role of sediment residence times during deposition. Donaldson and de Kemp (1998), in their study of Archaean quartz arenite sequences from the Superior and Churchhill Provinces, Canada, suggested a fluvial to shallow marine setting for the Woodburn Lake Group. Because a fluvial setting is not normally conducive to the development of quartz-rich deposits, the authors concluded that the combination of a granitoid source and intense chemical weathering were responsible for the quartz enrichment. Corcoran and Mueller (2002) compare the relative influence of chemical weathering during development of the 2.6 Ga Keskarrah Formation with two other Slave Province clastic sequences of similar age, based on geochemistry and clasts in sandstones and conglomerates. The Keskarrah, Beaulieu Rapids and Jackson Lake Formations developed in tectonically active, fault-bound basins and are characterised by one or more fining-

Opposite: Fig. 5.11-2. Characteristics of the quartz-rich sandstones in the Keskarrah formation. (a) Buff white weathered, laminated quartz arenite divided into planar bedsets (dashed lines). Arrow points to top. (b) Photomicrograph of sandstone containing 94% quartz. Note the matrix containing white mica (m). Field of view is 2.5 mm. (c) Photomicrograph of quartz arenite containing 95% quartz. Note the subangular to subrounded grains, inconsistent with recycling. Field of view is 2.5 mm.

5.11. Aggressive Archaean Weathering

499

500

Chapter 5: Evolution of the Hydrosphere and Atmosphere

(a)

(b)

: _

_

....a ......... ~176 ...... . .... ,o "~176176176176

02

[ 1

-.o......:::::::::::::::::::::::::::::::::

0.1

1

1 ....

I .... 55

I ....

I .... 65

I ....

1 .... 75

I ....

I .... 85

1

1

1

I

i

wl

I

10

30

Sc (ppm)

I .... 95

Si02% o []

Keskarrah Formation Beaulieu Rapids Formation

Fig. 5.11-2. Harker variation diagrams of (a) MgO wt.% versus SiO2 wt.%, and (b) Th/Sc (ppm) versus Sc (ppm), displaying clear distinctions between the Keskarrah and Beaulieu Rapids Formation samples. The clusters support a mixed mafic volcanic/plutonic source for the Beaulieu Rapids Formation and a more plutonic-dominated source for the Keskarrah Formation, in agreement with point and clast count results.

upwards sequences. The three successions, which formed in high relief settings, are relatively quartz-rich with respect to their source rocks, thus providing the initial indication that chemical weathering may have influenced strongly the mineralogy of the sedimentary deposits. A comparison of point and clast counts from two of the basins, the Keskarrah and Beaulieu Rapids Formations, supports a mainly granitic source for the former and a mixed mafic volcanic-plutonic source for the latter. These findings are commensurate with the geochemical results, which show higher Sc and MgO values for the Beaulieu Rapids Formation and higher Th and SiO2 values for the Keskarrah Formation (Fig. 5.11-3). CIA (section 5.10) values were calculated for selected fine-grained sandstones and siltstones, whose averages were then plotted on an A - C N - K ternary diagram (Fig. 5.11-4a). Average CIA values of 66 and 71 for the Keskarrah and Beaulieu Rapids sandstones, respectively, indicate high degrees of chemical weathering, with the siltstone sample from the Keskarrah Formation, having a very high value of 95, suggestive of intense weathering conditions. Because the sandstones lie to the right of the chemical alteration trend for average upper Archaean crust, a correction for potassium metasomatism was conducted, as is described by Fedo et al. (1995) (see also section 5.10). The resultant CIA values of 78 and 83 for the Keskarrah and Beaulieu Rapids Formations, respectively, further substantiate the theory of a weathering-aggressive Archaean atmosphere. In addition to the CIA, the PIA (Fedo et al., 1995) is a reliable tool in determining chemical weathering intensity and resolves the common problem of K abundance due to metasomatism (see also, discussion in section 5.10). When plotted on an A - C - N diagram, the average PIA values for the Keskarrah

5.11. Aggressive Archaean Weathering

(a)

501

projected pre-metasomatized values

J

100 - -

A1203

90--

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

82

\

80--

70-~

60--

m

r

5040-30-20-10-O--

CaO*+ Na20

K20 AI203

(b)

100

"

K20

--

90--

80-i

13. 70--

60--

50--

CaO*

Na20

Fig. 5.11-3. Ternary diagrams illustrating the extent of chemical weathering on Archaean sedimentary rocks. (a) A-CN-K diagram for Keskarrah and Beaulieu Rapids sandstones, Keskarrah siltstone (data from Corcoran and Mueller, 2002), Buhwa shales (data from Fedo et al., 1996), and Rampur Group pelites (data from Bhat and Ghosh, 2001). Note the increase in CIA values (small numbers) with correction for K metasomatism (dashed and solid lines). (b) A-C-N diagram for Keskarrah and Beaulieu Rapids sandstones and Keskarrah siltstone. High PIA values (small numbers) support intense chemical weathering. Note the diagram is cut at 50% A1203-K20. Ternary diagrams from Nesbitt and Young (1982) and Fedo et al. (1995).

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

and Beaulieu Rapids sandstones are 76 and 82, respectively, with a very high value of 98 for the Keskarrah siltstone (Fig. 5.11-4b). Although the CIA and PIA values support aggressive chemical weathering, the Beaulieu Rapids Formation does not contain quartz arenites (rocks containing > 95% quartz), as is the case for the Keskarrah Formation. This discrepancy is attributed to differences in depositional environment and source composition. Because the Beaulieu Rapids Formation developed in a mainly alluvial/fluvial setting, it was not subject to wave reworking and sorting, processes which the Keskarrah sediments underwent. In addition, the source rocks for the Beaulieu Rapids Formation had a stronger mafic volcanic component than the Keskarrah Formation sources. Consequently, in order to develop quartz arenites in high-relief, tectonically active settings, a combination of a plutonic-dominated source, high degrees of sorting and significant chemical weathering are crucial. All of these parameters were met during evolution of the Keskarrah Formation (Fig. 5.11-5). Fedo et al. (1996) studied shales from the Archaean Buhwa greenstone belt, Zimbabwe, and found that the samples underwent intense chemical weathering, possibly as a result of the high CO2 content of the Archaean atmosphere. Large ion lithophile elements (LILE), such as Rb and Ba had high values compared with smaller LILE, for example, Na, Ca, and Sr, which tend to be selectively leached from the alteration profile. The average CIA value for the shales is approximately 75, but Fedo et al. (1996) believe that the rocks experienced K-metasomatism. The majority of the samples fall on the AlzO3-K20 join on the A - C N - K ternary diagram (Fig. 5.11-4), which reflects a weathered granitic source. However, immobile element behaviour is more consistent with a tonalitic source, thus supporting K-metasomatism. The pre-metasomatism average CIA value was determined to be approximately 95, in accordance with intense chemical weathering of the source rocks (Figure 5.11-4a). The authors also recognised a paucity of feldspars in associated siltstones and sandstones, in accordance with alteration to aluminous clays. An average PIA value of 98 provides further evidence for conversion of primary plagioclase to clay minerals. Similarly, Bhat and Ghosh (2001) inferred high degrees of source chemical weathering for late Archaean (2.51 Ga) pelites from the Rampur Group, western Himalayas, based on several lines of evidence determined from geochemical results. Relatively high A1203/Na20 ratios and PIA values were consistent with intense chemical weathering. Although the CIA values were average (c. 68-77), the authors believe this to be a relic of K-metasomatism. The correction of the Rampur CIA values, following the method of Fedo et al. (1996), showed that the pre-metasomatism values were between 74 and 92, reflecting high to intense source rock chemical weathering (Fig. 5.11-4a). Sugitani et al. (1996), through chemical analyses of shallow water Archaean cherts from the Pilbara block, Western Australia, found that AlzO3/TiO2 values were fundamentally lower (some < 1.0) than the lowest value anticipated for terrestrially derived rocks. AIzO3/TiO2 ratios in sediments and sedimentary rocks are directly controlled by source rock composition (acidity) and degree of mixing with other sources. Considering values < 4.0 as anomalously low, the authors believe that Archaean atmospheric conditions were responsible for the Ti enrichment. The elevated CO2 content in the Archaean atmosphere and the resultant acidity of meteoric waters could have produced AI-Ti fractionation if the

5.11. Aggressive Archaean Weathering

503

Fig. 5.11-4. Palaeogeographic model of the Keskarrah Formation illustrating the ideal factors that produced quartz arenites in a high relief setting. The main source rocks were granodiorites with a secondary mafic volcanic source component. The source area experienced in situ chemical weathering under humid, hot, CO2-rich atmospheric conditions. Source material was eroded and the resultant sediments were subjected to chemical weathering during transport. Coalescing alluvial fans and fan-deltas formed along the slopes of the high relief basin and prograded onto a marine shelf. Finer-grained sediments were affected by longshore drift, swash~ackwash processes, storm transport, and subtidal waves, which further contributed to quartz enrichment. Modified from Corcoran and Mueller (2002).

pH values were < 4. In addition, high concentrations of Zr and HREE (heavy rare earth elements) relative to the LREE (light REE) is commensurate with chemical weathering of source rocks because the LREE are more mobile under low pH conditions (Sugitani et al., 1996). An alternative or additional cause of source rock chemical weathering could have been the build-up of acidic waters during geothermal activity, which would have been pervasive during the Archaean. Maynard et al. (1991) supported an oxygen-poor atmosphere conducive to aggressive chemical weathering during deposition in the c. 3.1-2.8 Ga Witwatersrand basin, South Africa. Abundant detrital uraninite and pyrite indicates deposition occurred in a lowoxygen atmosphere, based on the fact that these minerals are considered susceptible to oxidation under present-day conditions. Because the Witwatersrand quartz arenites and modern sands from the Indus River were both deposited in a foreland basin setting and

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Chapter 5: Evolution of the Hydrosphere and Atmosphere

both sequences contain detrital uraninite, these authors chose to examine whether they could be compared and used as a case against the climatic significance of uraninite in sedimentary deposits. Maynard et al. (1991) noted high quartz and low feldspar and rock fragment abundances in the Witwatersrand deposits, in addition to high CIA values, supporting high degrees of chemical weathering. In contrast, CIA values and quartz content in the Indus sediments are low and are more consistent with minor chemical weathering. Based on this finding, Maynard et al. (1991) argued that the two depositional systems could not be analogous and that the process involved in quartz enrichment (chemical weathering) for the Witswatersrand deposits is the same process responsible for uraninite stability (see Ohmoto, section 5.2, for detailed discussion of Witwatersrand uraninite and implications for early Precambrian palaeo-atmospheric composition).

Implications Archaean sedimentary sequences are prime candidates for evaluation of ancient atmospheric and Earth-surface conditions. Examples worldwide support the inference of aggressive chemical weathering, as they resemble the geochemical and petrographic imprints of present-day deposits from humid climatic settings. The main question that arises is: to what extent can these imprints be attributed to the effects of climate? (see also section 5.10). Although chemical weathering results in quartz enrichment, so too do recycling, sorting, and reworking in a shoreline environment. What role did sedimentary recycling have in the formation of these early deposits? Although post-Archaean recycling is estimated at approximately 90% cannibalistic (Veizer and Jansen, 1985b) (see also section 7.11), is it safe to assume that most Archaean sedimentary sequences are first-cycle deposits? Clues could be derived potentially from examining clast roundness. A quartz arenite developed through successive sedimentary cycles will contain rounded clasts, whereas intense chemical weathering could result in quartz-rich deposits with subangular to subrounded grains, as is common, for example, in the Keskarrah Formation quartz arenites (Fig. 5.11-2c). Archaean deposits are significant to our understanding of the early Earth's evolution. Geochemical and petrographic results not only corroborate inferences of a hot, humid, CO2- and CH4-rich atmosphere conducive to intense chemical weathering, but also show that in some cases, climate was the overriding factor. Under present-day surface conditions, an absence of vegetation in a high-relief setting promotes rapid burial, low residence times and poorly sorted deposits. The lack of vegetation in the Archaean should have promoted run-off, thereby decreasing sediment residence times, but recognition of Archaean quartz arenites that developed in tectonically-active settings indicates that the role of relief was subordinate compared with climate. Similarly, the importance of depositional setting may have been secondary to chemical weathering, based on the presence of alluvial-fluvial quartz arenites. The distinct possibility of climate as the overriding factor during deposition of Archaean sediments supports an early Earth atmosphere that was much more weathering-aggressive than that of the present day.

5.12. C o m m e n t a r y

5.12.

505

COMMENTARY

EG. ERIKSSON AND W. ALTERMANN There is an intimate relationship among life, atmospheric and ocean chemistry on Earth (Ohmoto, section 5.2); biogeochemical signatures within the sedimentary record thus provide a way of studying the Earth's early biosphere and atmosphere (Knoll and Canfield, 1998). The oxygen content of the atmosphere influences the geochemical cycles of sulphur and carbon, amongst many other elements (Ohmoto, section 5.2); carbon (section 5.3) and sulphur (section 5.5) isotopes thus provide possible proxies for Precambrian palaeoredox and biologic inventory. Although essentially all free oxygen is biological in origin, oxygenic photosynthetic organisms such as algae and Phanerozoic-modern plants are themselves aerobes (Ohmoto, section 5.2); cyanobacteria, however, are facultative aerobes because they are capable of using molecular oxygen for their metabolism, but do not necessarily have to do so (section 6.3). The time-point of the emergence of cyanobacteria and photosynthesis in Earth history is not certain, but appears to be well bracketed by c. 2.7 and 2.5 Ga (Brocks et al., 1999; Kazmierczak and Altermann, 2002). Morphological evidence in microfossils and isotopic evidence in bulk rock compositions point to a possibly much earlier appearance for both. The organic matter resulting from the well known photosynthesis reaction is almost all decomposed (in the long term) when exposed to the atmosphere and surface water, through a variety of pathways, and all atmospheric oxygen is thus renewed every c. 3000 years (section 5.2). The production of organic matter and the negative feedback mechanism when this matter decomposes, apart from partially controlling atmospheric oxygen, also influences concentrations of atmospheric greenhouse gases such as CO2 and CH4 (Ohmoto, section 5.2). The other major control on atmospheric oxygen is the long term changes in various oxygen fluxes. As a small portion (0.14% at present) of organic matter in marine sediments (terrestrial equivalents are insignificant) is buried and thus removed from the decomposition part of the Corg-O2cycle, accumulation of oxygen over time scales > 3000 years occurs (Lasaga and Ohmoto, 2002). Long-term 02 consumption occurs through oxidation of reduced volcanic gases such as H2, HzS, SO2, CH4 and CO, and by oxidation of fossil carbon (kerogen) in sediments during pedogenesis (Holland, 1978; section 5.2). The present-day inferred steady-state atmospheric 02 value thus reflects the fact that long term oxygen production and consumption fluxes are in balance; when they are not, long term accumulation or loss of atmospheric O2 occurs (Ohmoto, section 5.2). Two major models exist for atmospheric oxygen evolution: (1) the " C - W - H - K model" (e.g., Cloud, 1968; Walker, 1977; Holland, 2002; Kasting and Siefert, 2002), based on the assumption that life originated under reducing atmospheric conditions; an early, minor rise of O2 at 3 or 2.8 Ga (Rye and Holland, 2000; Kasting and Siefert, 2002) was followed by a more oxic atmosphere at c. 2 Ga, which enabled emergence of the eukarya. This model suggests biogenic methane as the primary greenhouse gas (with subordinate CO2) counteracting the "faint young sun" before the rise of O2 at The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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c. 2.2 Ga. Carbon dioxide levels of 300 PAL (present atmospheric level) and < 0.1 ppm O2 at c. 3 Ga are inferred; furthermore, this model implies anoxic oceans until c. 600 Ma, except for the photic zone (less than c. 100 m). Low SO42 levels (except in local evaporitic basins) are inferred until c. 2.2 Ga, with a gradual increase until c. 0.8 Ga when a second step-wise increase to modern levels occurred. Before c. 1.8 Ga, global oceans likely had high Fe z+ and low HzS, with the reverse pertaining from c. 1.8 till c. 0.8 Ga, and at c. 0.6 Ga the oceans became SOl--rich and HzS-poor (Walker and Brimblecomb, 1985; Bjerrum and Canfield, 2002). (2) the "D-O" model (e.g., Dimroth and Kimberley, 1976; Lasaga and Ohmoto, 2002) infers emergence of oxygenic photosynthesis and possibly of cyanobacteria soon after the first differentiation of oceanic and continental lithosphere at c. 4 Ga, followed by a single rise in atmospheric O2 from < 10-3 to c. 1 PAL soon after 4 Ga, and thereafter relatively constant pO2 levels within 50% of PAL. This model further suggests low CH4 levels, CO2 as the primary greenhouse gas (countering the faint young Sun), as well as oxygenated oceans with essentially constant SO42 levels (apart from semiclosed local basins) since c. 4 Ga; Fe z+ and HzS in normal oceans (i.e., except for local anoxic basins) remained low from c. 4 Ga. Different models (such as the "C-W-H-K" and the "D-O" models above) for biogeological evolution and the evolution of atmospheric and oceanic chemistry rest largely on differences in the interpretation of the same geological, palaeontological, and biogeochemical data: (1) the Precambrian fossil record; (2) the presence, in the early Precambrian record, of minerals such as uraninite, pyrite and siderite, which are unstable in modern fluvial sedimentary environments; (3) the known and inferred behaviour of Fe in subaerial environments (particularly palaeosols and red beds) and, (4) in marine environments (especially banded iron-formations); (5) the geochemical sulphur cycle and evolution of sulphur-utilising bacteria, and (6) the geochemical cycle of carbon. The data themselves are equivocal in that they apparently support essentially mutually exclusive models. The major reason for this is the uncertainty whether the data indeed indicate original conditions or whether they have been altered by later influences; incorrect stratigraphic interpretations are also pertinent in this regard. As an example, the problems inherent in using geochemical (isotopic) tracers for estimating palaeo-oceanic chemistry in Precambrian basins are illustrated by the Neoproterozoic Gariep basin, Namibia, where C and Sr isotopic compositions exhibit large temporal variations within a depository which had a fully open ocean exchange (Frimmel, section 5.8). Carbon isotopic data from carbonate rocks in Australian Precambrian basins (613Ccarb-time curve) resemble closely global trends; the curve is flat in the Neoarchaean (at c. 2.6 Ga) and into the early Palaeoproterozoic (Lindsay and Brasier, section 5.3). Although an early Archaean evolution of sulphate reducing bacteria is possible, an oxygen-deficient atmosphere at that time (following the C - W - H - K model) would have resulted only in local examples of signature isotopic fractionations (Lyons et al., section 5.5). As oxygen levels possibly increased in the Palaeoproterozoic (in contrast to the "D-O" model), associated with continental weathering, sulphate concentrations probably

5.12. Commentary

507

increased to levels where bacterial sulphate reduction was well expressed isotopically, as is evident from abundant 34S-enriched pyrites and relatively marked sulphur isotope variation in marine sulphates and sulphides (section 5.5). Two major 613Ccarb oscillatory excursions occur at c. 2.2-2.3 Ga and at c. 0.65 Ga, separated by essentially flat patterns (Lindsay and Brasier, section 5.3). It is inferred that the 613Ccarb patterns reflect the release of endogenic planetary energy through hydrothermal activity, plate tectonics and the supercontinent cycle (Lindsay and Brasier, 2002; Brasier et al., 2002) rather than atmospheric compositional changes on their own. Major supercontinental assembly events at c. 2.8, 2.0 and 1.0 Ga were associated with mantle instability (sections 3.2-3.4; see also last paragraph below) and concomitant large intracratonic sag basins accumulated sediments over 200-500 My periods, including major carbonate platform successions (section 5.3). Karhu and Holland (1996) interpret the large positive 313C excursion they identify in carbonate rocks at c. 2.2-1.9 Ga as indicating a significant increase in the burial flux of organic carbon (with concomitant increase of the O2 production flux; see section 5.2) to result in a "Great Oxidation Event (GOE)" of c. 2.3 Ga. The problem with this interpretation is that the "event" predates deposition of the relevant carbonates by c. 100 My (Ohmoto, section 5.2). Ohmoto emphasises that application of the C isotopic mass balance approach to constrain the burial flux of Corg will provide information on the ratio of the burial fluxes of organic carbon to carbonate carbon, rather than constraining the burial flux of organic carbon itself (section 5.2). Ohmoto (section 5.2) suggests that models supporting an anoxic atmosphere prior to a c. 2.3 Ga "GOE" (e.g., Holland, 1994; Rye and Holland, 1998; Kasting and Brown, 1998; Kump et al., 2001) do not take account adequately of the coupling of two significant negative feedback mechanisms: (1) the burial flux of Corg depends on atmospheric pO2, and as a result a decrease in the latter from 1 PAL to 0.5 PAL would result in a seven-fold increase in the O2 production flux; (2) the O2 consumption flux by soil oxidation provides the other mechanism. Ohmoto (section 5.2) maintains that coupling of these two negative feedback mechanisms rather suggests that pO2 was maintained at c. 0.5-2 PAL since the emergence of the oxygenic photosynthetic organisms at > 2.8 Ga, or even as far back as > 3.7 Ga. In Australia, the first major Ccarb isotopic excursions followed development of the Hamersley basin, with striking coincidence of isotopic oscillations with major BIF (peaking at c. 2.5 Ga; Trendall and Blockley, section 5.4) deposition (as Fe precipitated with circulation of oxygen through the oceans) and with evidence for the earliest global glaciation thereafter (Lindsay and Brasier, section 5.3), at c. 2.45m2.4-2.2 Ga (Young, section 5.6). However, genetic modelling of IF (assuming it to be a lithified chemical precipitate with close correspondence of primary composition and that preserved today) and its temporal and global distribution, suggests that an association with ocean-atmosphere oxygenation may be invalid (Trendall and Blockley, section 5.4). Palaeoproterozoic glaciogenic deposits occur in passive margin tectonic settings in North America, and partly in foreland basin settings in Western Australia (Young, section 5.6). In South Africa they may occur within a large epeiric (e.g., section 7.7) basin.

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In contrast to the more widespread and stratigraphically complex Neoproterozoic glacial deposits, the c. 2.4-2.2 Ga deposits were possibly of shorter duration and show no temporal association with BIF or cap carbonates. The "snowball earth hypothesis" (SEH) (e.g., Kirschvink, 1992; Hoffman et al., 1998b) is often applied to both Proterozoic glacial intervals, but field data do not support this idea (see discussions in sections 5.6-5.8). The global nature of glaciation is inherent in the SEH, but contemporaneity and global distribution of the Neoproterozoic glaciogenic products of discrete events are poorly constrained, with a current temporal distribution over almost 300 My (Young, section 5.6; see also section 5.8). Evidence for low palaeolatitudes (e.g., Harland and Bidgood, 1959; Harland, 1964; Embleton and Williams, 1986; Schmidt et al., 1991; Schmidt and Williams, 1995; Sohl et al., 1999) and glaciation at and near sea level is commonly used to support SEH (section 5.7). Strong arguments against snowball Earth (SE) include: thick diamictite successions; associated waterlain deposits; evidence for glacial cycles and eustatic oscillations; strong evidence for unfrozen seas from preserved tidalites, tidal rhythmites and thick glaciomarine successions; evidence for gradual climatic changes at the onset and amelioration of Proterozoic glacial events; and indications of strong seasonality in glaciogenic deposits (e.g., Williams 1975, 1986, 1998) (Williams, section 5.7; Young, section 5.6). There is also a lack of unequivocal high latitude Proterozoic glacial deposits (Evans, 2000). Although proponents of SE suggest that Neoproterozoic iron-formations reflect re-oxygenation of the oceans (in which hydrothermal iron built up during SE) at the end of glacial events (Kirschvink, 1992), their stratigraphic distribution, facies associations and geochemistry may also be explained by a depositional model where glaciers flowed into developing Red Sea-type rift basins (Young, section 5.6). There is little doubt that Earth experienced major climatic perturbations in the Palaeoand Neoproterozoic (see also section 5.3), but these two, as yet poorly understood widespread glaciations are preceded by development of supercontinents (e.g., sections 3.2 and 3.9) (Young, section 5.6). The possible Neoproterozoic rift basins may then relate to supercontinent attenuation and breakup (section 5.6; see also section 5.8, for an example of the relationship between the supercontinent cycle and Neoproterozoic glacial deposits). It is postulated that high continental freeboard (section 7.1) concomitant with supercontinent assembly enhanced weathering regimes (sections 5.10 and 5.11) and thus CO2 drawdown also, and low palaeolatitudinal location would have increased albedo, together leading to global cooling (Young, section 5.6). The large obliquity hypothesis, although not providing a prime cause of global glaciation, offers a viable mechanism for the distribution and nature of glacial environments and can explain many of the features observed in the Neoproterozoic glacial deposits (e.g., Williams, 1975, 1993; section 5.7). Frimmel (section 5.8) uses geochronology and chemostratigraphy to assess sedimentation rates (see also section 7.11) of glacial and non-glacial deposits from the 770-540 Ma period in the Gariep basin, Namibia. Two glaciogenic units identified there are correlated with the global Neoproterozoic Sturtian (750-740 Ma) and Marinoan (590-580 Ma) (or possibly with the 560 Ma Moelv) glaciations (section 5.8). Positive 613C excursions are not reliable stratigraphic markers due to their host carbonate rocks having been deposited in either restricted basin conditions or under very shallow water (Frimmel, section 5.8).

5.12. Commentary

509

Integrated sedimentation rates for the interval between the two identified glacial events are anomalously low, suggesting either a c. 100 My period characterised by sea level lowstand (section 8.2) and cold palaeoclimate, or low subsidence rates due to enhanced midocean ridge spreading rates. Frimmel (section 5.8) concludes that the sedimentary record, and C and Sr isotopic chemostratigraphy (cf. sea water proxies) for the post-Sturtianpre-Marinoan period reflect the interaction of tectonic (and concomitant hydrothermal influences), eustatic and palaeoclimatic factors. As weathering of igneous rocks is influenced mainly by climate (largely temperature and humidity), ancient weathering profiles, palaeosols and sedimentary rocks enable palaeoclimatic reconstruction (Nesbitt and Young, section 5.10). Although bulk chemical and mineralogical zonation can be well preserved in such Precambrian successions, the interplay between chemical weathering and mechanical erosion (influenced mainly by tectonism and eustasy) is important when interpreting palaeosols, and diagenesis and metasomatism are complicating factors. Certainly, raised greenhouse gases such as CO2 (section 5.2) during the earlier Precambrian would have enhanced weathering (section 5.10); in the Archaean, in addition, enhanced geothermal activity would have promoted acidic waters and led to generally aggressive weathering conditions (Corcoran and Mueller, section 5.11). The lack of vegetation in Precambrian palaeoenvironments would have translated into faster erosion and abbreviated detritus residence periods before burial, thereby reducing the effects of the enhanced weathering regimes (Condie et al., 2001; section 5.11). Iron-formation (IF) is essentially a Precambrian rock type, with the oldest known banded iron-formations being c. 3.8 Ga, and succeeding early Precambrian IFs are generally relatively thin, laterally restricted and commonly associated with greenstone belt volcanism (Trendall and Blockley, section 5.4). A peak in (banded) IF-time distribution is seen at c. 2.5 Ga, largely due to deposits in the Hamersley (Pilbara craton, Australia) and Transvaal (Kaapvaal craton, South Africa) basins, with a later period of abundant (granular) IF at c. 1.8 Ga; a long hiatus preceded local development of varied small IFs of Neoproterozoic age (section 5.4). It is thought that iron and silica precipitated (the former by photosynthetic activity) together annually in large, gently subsiding, deep (> 500 m) offshore shelf basins, from an ocean water source with a concentration of 10-20 ppm dissolved ferrous iron, derived from primary seafloor volcanism and retained in solution for long periods (Trendall and Blockley, 1970; Trendall and Blockley, section 5.4). Archaean sea water is thought to have been enriched in iron only below the pycnocline (Eriksson et al., 1997). Upwelling of deep, iron-rich water (Holland 1973, 1984) into the shallow shelf settings of the Palaeoproterozoic Transvaal and Hamersley basins (Klein and Beukes, 1989) or a pycnocline close to the level of these basin floors enabled major deposition of IF (Trendall and Blockley, section 5.4). This iron- and Eh-stratified ocean model obviates the use of IF occurrence in the Precambrian record as a proxy for atmospheric oxygen content, as IF deposition could have occurred without direct atmospheric influence. As in most research on iron-formations, the contribution by Trendall and Blockley (section 5.4) omits the silica problem. The amounts of SiO2 in IF are enormous and its source and mode of precipitation remain uncertain. Silica precipitation without involvement of silica-precipitating organisms in uncertain chemi-

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cal environments is as yet poorly understood. This problem is reviewed by Klemm (2000), who discusses a model of SiO2 precipitation from saturated (15-20 ppm SIO2) sea water, at pH 7.5-8 and Eh of 0.1-0.3. Applying the stratified ocean model, the lack of major IF in the mid-Precambrian may be explained by increasing oceanic oxidation as organisms and photosynthesis became more abundant and efficient (section 5.4). Episodic growth of the global reduced carbon reservoir supports the possibility of stepwise oxygenation of the atmosphere consequent upon episodic burial of carbon during large scale tectonic cycles (Des Marais, 1994a, 1997; des Marais et al., 1992) (section 5.3). The basically flat portions of the ~13Ccarb curve, prior to c. 2.2-2.3 Ga and in the mid-Proterozoic (section 5.3), are related to low tectonic activity and CO2 in the ocean-atmosphere system being in near-equilibrium with the mass balance of the carbon cycle (Brasier and Lindsay, 1998; Lindsay and Brasier, 2000). Although sulphate availability in the early and mid-Proterozoic ocean remained low (relative to the Phanerozoic), concentrations and associated BSR (bacterial sulphate reduction) were possibly able to support globally euxinic conditions in the anoxic deep ocean waters (section 5.5). Oxygen levels may have risen significantly by the Neoproterozoic, thereby raising oceanic sulphate, to allow enhanced sulphur isotope fractionation between sulphate and sulphide to achieve values typical for the Phanerozoic; the increasing fractionation probably reflects bacterial evolution and related isotopic effects within the oxidative part of the biogeochemical sulphur cycle (Lyons et al., section 5.5). Palaeoclimatic modelling based on Earth's inferred palaeorotation and applying sophisticated software packages, suggests significant changes, including equatorwards latitudinal shift of the planetary-scale circulation belts and reduced wind speeds throughout the atmosphere (e.g., Rautenbach, 2001). The theory of lunar tidal friction proposes that attraction of the Moon on Earth's tidal bulge results in the transfer of angular momentum from the spin of the Earth to the Moon's orbit, thereby slowing Earth's rotation and leading to recession of the Moon (Lambeck, 1980); the Sun has a smaller effect on Earth's rotation rate (Brosche and Wtinsch, 1990) (Williams, section 5.9). The implications of this theory, when applied to projection of the present rates of tidal energy dissipation back in geological time, is a catastrophic near-approach of the Moon to Earth at c. 1.5 Ga, for which there is no evidence; less reliable data from older rocks suggest such an approach at earlier times, up to c. 3 Ga, but again, this seems an unlikely event (Williams, section 5.9). Cyclic tidal rhythmites provide a means for estimating the Precambrian Earth's rotation and the orbit of the Moon (section 5.9). Data from the c. 600 Ma Reynella-Elatina Formations and from the c. 750 Ma Pualco Tillite, both from South Australia, indicate a length of day (LOD) of 21.9 and 21.4 hours per day, respectively (section 5.9). Data from older Precambrian rocks do not permit reliable estimations of the change in Earth's rotation rate to be made. These palaeotidal and palaeorotational data also allow a test to be made of Earth's expansion, with no evidence for significant change, at least from the late Neoproterozoic (Carey, 1976; Williams, 1998b, 2000). As Earth's rotation slowed down due to tidal friction, its forced nutations (periodic tipping) under the combination of lunar and solar torques were subject to resonances with the free nutation fluid of the core, at various frequencies (Toomre, 1974; Hinderer and Legros,

5.12. Commentary

511

1988). LOD can be used to estimate that annual resonances probably took place during the late Neoproterozoic-early Palaeozoic, and there were likely important resonances during the Archaean as well (Williams, section 5.9). Resonance at the core-mantle boundary and concomitant instability would have resulted in heat release by mantle (super)plumes (sections 3.2-3.4) and the relationship of celestial mechanics (see also chapter 1) with temporal distribution of superplumes is thus most likely an important one. As is obvious from many of the principles discussed above, the primary controls on Earth's (Precambrian) geological evolution lie in the interaction of plate tectonics with mantle superplumes and related thermal processes (chapters 2-4); interacting secondary influences, palaeoclimate, eustasy (chapter 8), ocean-atmosphere chemistry (chapter 5), geo-biology (chapter 6), and sedimentation (chapter 7) are directly dependent on these primary controls (e.g., Eriksson et al., 2001 a, b).

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The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) 9 2004 Elsevier B.V. All rights reserved

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Chapter 6

EVOLUTION

6.1.

OF LIFE AND PRECAMBRIAN

BIO-GEOLOGY

INTRODUCTION

W. ALTERMANN Having discussed the earliest history of the Earth and the solar system in chapter 1, and the endogenic forces acting during the formation of continental crust and driving plate tectonics, superplumes and volcanism (chapters 2, 3 and 4, resp.), attention here focuses on the exogenic forces shaping the Earth' surface and shallow interior. This was begun already in chapter 5, where the evolution of the atmosphere and hydrosphere and their impact on sedimentary environments, climate and weathering conditions were elucidated. However, the composition of the atmosphere and hydrosphere is largely dependent on the existence of life and on its metabolic activity, and Earth's unique, oxygen-rich atmosphere is the result of this activity. The origin of life on Earth lies in the Hadaean darkness and the question of how and where life began remains unresolved. It is generally accepted that life arose early in Earth history, as soon as conditions became favourable for preservation of simple organic compounds, through the formation of a stable crust on Earth, and the atmosphere and oceans. This was probably achieved some 100-300 My after the accumulation of cosmic dust to form the Earth (section 1.2). The energy necessary for the synthesis of complex organic molecules was provided by the Sun's luminosity, by atmospheric electric discharge and by geothermal activity. Light is the most readily available energy source, and therefore photochemical reactions probably played a major role in the pre-biotic synthesis of organic compounds. The carbon participating in these reactions could have been remanent from the primary accretion phase of the Earth, or could have originated from extra-terrestrial reactions and carbonaceous chondrites (McClendon, 1999). Because the early Earth offered a cornucopia of various environments for the abiotic synthesis of life, and because we do not know of any extant or extinct life in the universe, despite the existence of complicated organic molecules, there seems to be no need for explaining the emergence of life on Earth by germination from space. Such fruitless theories add little to our knowledge on the origin of life. Terrestrial environments like deep submarine vents, continental thermal springs, volcanic vents, impact craters and hypersaline lagoons are usually regarded as hostile to life (cf. sections 1.2 and 6.6). Nevertheless, such environments with readily available energy were widespread on the early Earth, and it is likely that some of them may have served as the cradle of life (Schopf, 1999). Extremophilic organisms inhabiting such environments (i.e., bacteria, archaea), are among the most primitive in the evolutionary di-

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versification path. Thermophilic, acidophilic and sulphur-based metabolism were probably present in the earliest prokaryotes. The earliest living cells must have been self-assembled complex molecular systems of polymers, with a self-replicating mechanism, encapsulated by a membrane of lipids and proteins (Deamer, 1993). It is a matter of controversy whether the first organisms were heterotrophic (using abiologically produced organic compounds from the surrounding environment as source of energy) or autotrophic. Oparin (1938) was the first to propose that the earliest organisms were heterotrophic. Later workers (e.g., de Duve, 1991; W/~chtersh~iuser, 1994) believe that the first organisms were chemoautotrophic (e.g., using carbon dioxide to produce their own energy carriers like glucose and sugar). The chemical reaction to form the first living cell and the beginning of biological evolution remains a mystery. Four stages of development were essential in the origin of life (Schopf, 1992d, 1999). (1) Specific environmental conditions must have enforced self organisation of the major elements of life (C, H, O, N and P) to form complex molecules and simple organic compounds. (2) These compounds became stabilised by using energy from the environment and by building more complex compounds (nutrition and growth) and protecting themselves from hydrolysis (by membranes). (3) Self-replication was invented through unknown processes. (4) Mutation in the replication process allowed for evolutionary processes and for the development of new species. The timing and environment relative to the appearance of early life on Earth are important for understanding of past and present ecosystems, and for exploring the possibility of extraterrestrial life. Geochemical isotopic evidence reported from Archaean cratons suggests that life has existed on Earth since at least 3.8 Ga. These reports, however, have been criticised repeatedly as reflecting detection of traces of younger colonisation of rocks by (endolithic) bacteria (contaminants), or as reflecting misinterpretations during mapping and geological interpretation of the host-rock formations (e.g., cf. sections 2.2 and 6.2). Our understanding of the genesis of these rocks from the ancient Archaean (> 3.5 Ga) is still inadequate. Recent studies have also cast doubt on the genesis of some of the Earth's oldest fossils, implying that structures mimicking bacterial fossils formed abiologically in hydrothermal environments. The controversy and the public debate on these findings have attracted much attention in the scientific community. Any reports of the record of early life, therefore, must hold up to the most rigorous inspection, before they can be accepted as genuine and authentic. In section 6.2 the pioneer of Archaean palaeobiology, J.W. Schopf, defends passionately his research and elucidates the "rules of the hunt" for Archaean microfossils. The criteria for recognition and acceptance of microfossils as authentic Archaean occurrences are explained together with the various methods of investigation. Later in section 6.2 the Archaean fossil record as known today, is discussed in the light of these explanations. In the subsequent section 6.3, the effects of life on Precambrian sedimentation and sedimentary environments are discussed by W. Altermann. In the Mesoarchaean, this influence was still restricted to small areas and its effects remain vague. As long as life did not significantly alter the composition of the atmosphere and hydrosphere, no particular sedimentary deposits can be ascribed directly to organic metabolism. Only at the end of the

6.1. Introduction

515

Archaean, from c. 2.9 to 2.6 Ga onwards, did biomineralisation (sections 6.4 and 6.5) and the buildup of huge carbonate platforms became an apparent result of life in the sedimentary record. Another result of emerging life might have been the stabilisation of siliciclastic sediments by microbial mats. In the Proterozoic, however, life became a significant influence on sedimentary environments, ranging from sedimentation of giant BIF deposits, that can be attributed to the metabolic activity of organisms (cf. also sections 5.2 and 5.4), to a strong influence on the weathering patterns of rocks via altered atmospheric composition. In the Neoproterozoic, sedimentation patterns were altered again, with the development of organisms that were able to burrow and graze the sediments or produce skeletal reefs, restructuring the architecture of continental shelves. In section 6.4, the microbial origin of Precambrian carbonates is discussed with an actualistic approach, by J. Kazmierczak, S. Kempe and W. Altermann. Modern environments sustaining in situ calcifying benthic cyanobacterial mats and producing carbonate deposits analogous to Precambrian sedimentary carbonates are discussed, and high alkalinity and sodium predominance are proposed to serve as a model for Precambrian seas. Structural, textural and mineralogical similarity of Precambrian and modern samples support the assumption that Precambrian and modern sedimentary carbonate rocks are products of very similar microbiota and calcification processes. The arguments for a highly alkaline Precambrian ocean are debated and chemical reactions governing water alkalinity, Ca-saturation and bio-calcification processes are illuminated in analogy to modern examples. Problems in definition, classification and stratigraphy of Precambrian stromatolites are discussed in section 6.5, by W. Altermann. Next to a stromatolite classification scheme taken from Hofmann (2000), the usefulness of non-biologic versus biologic definitions of stromatolites, and the influence of biological versus non-biological sedimentary factors on stromatolite morphology are debated. Photographic plates showing Archaean stromatolites from different environments are presented and the stratigraphic usefulness of stromatolites is questioned. Stromatolites are viewed in this outline of current Precambrian stromatolite research, as being mainly environmental indicators. The question of possible extraterrestrial life and the lessons we can learn from Precambrian geology in respect to exobiology are reviewed in section 6.6. Here, E Westall explains how Precambrian geology and palaeobiology are interwoven with the search for extraterrestrial life and with the investigations of extraterrestrial conditions, where life could possibly have evolved, or become adapted to such conditions. In contemporary times many earth science departments at universities world-wide have founded new institutions or devoted their palaeobiological sections to exobiology or astrobiology. Their research concentrates on studies of the possible origin, spread and adaptation of life in the universe. This novel discipline uses biological, biochemical, molecular and geological methods and is becoming increasingly popular and interdisciplinary. Although an authentic subject for study has yet to be found (e.g., an extraterrestrial living or fossil organism), it is revealed in section 6.6 of this chapter, that conditions on several planets were in the past, or even are today, comparable to many environments of the Precambrian Earth. Therefore, life could have originated, theroetically, also on such planets by similar processes and in similar environmets to those of the Precambrian Earth. Because the level of conservation of old rocks on such plan-

Chapter 6: Evolution of Life and Precambrian Bio-Geology

516

ets is probably much higher than on Earth, due to the lack of plate tectonic activity, we can perhaps learn even more on the origin of life from rocks sampled by extraterrestrial missions.

6.2.

EARTH'S EARLIEST BIOSPHERE: STATUS OF THE HUNT

J.W. SCHOPF

Nature of the Problem The Precambrian segment of geological time encompasses some 85% of the history of life, a record dominated by pr'okaryotic (bacterial and cyanobacterial) microbes. Such microorganisms are minute, incompletely preserved in geological materials, and exhibit simple morphologies that can be mimicked by non-biologic mineralic microstructures; discrimination between true microbial fossils and microscopic pseudofossil look-alikes can therefore be difficult. In the younger (Proterozoic) Precambrian, differentiation of true fossils from pseudofossil mimics usually presents only minor problemsmfossils are abundant, often well preserved, and not uncommonly so similar in morphology to modern microbial taxa that both their biogenicity and biological affinities can be established readily (Mendelson and Schopf, 1992). But interpretation of the older (Archaean) Precambrian fossil record is fraught with difficulties. Because of geological recycling, preserved rock units become increasingly rare with increasing geological age (cf. crustal growth rates, section 2.8); few units dating from the Archaean (> 2,500 Ma) have survived to the present. Moreover, because of tectonism and related geologic processes, most Archaean rocks have been metamorphosed, often severely, with older units generally being more altered than younger ones. Taken together, these two factorsmthe lack of survivability and the metamorphic alteration of the ancient rock record~play havoc with attempts to trace back the earliest records of life. Indeed, in comparison with that of the Proterozoic, the Archaean fossil record is miniscule, comprised of few bonafide fossils scattered over an enormous segment of geological time (Schopf and Walter, 1983; Schopf, 1992a). Yet the rewards to science of tracing the roots of life ever farther into the remote recesses of the geologic past, of uncovering the record of Earth's earliest fossils, is similarly enormous. Molecular phylogenies, such as the rRNA-based "Tree of Life", can provide firm evidence of the relative relations of the various microbial lineages, of "earlier" relative to "later" times of emergence. But at the current status of the science, only the known fossil record can provide a clear-cut understanding of the antiquity of life on Earth (and, of course, even then can provide only a minimum age for life's existence). And the solution to this problem holds the key to not only unravelling the knotty problem of when life emerged on Earth, but to the broader question of whether life once did or now exists elsewhere in our solar system, since the same questions will be asked and similar techniques applied in the search for evidence of past life in rock samples from other planetary bodies (cf. section 6.6). The PrecambrianEarth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Because of the difficulties inherent to the hunt for records of life's earliest history, in years past mistakes have been made (Schopf and Walter, 1983) as researchers have sought to extend the limits of the field, only to discover that their reach exceeded their grasp. Indeed, accepted evidence of early life has proven elusive as this area of science has developed into one of the most contentious in all of palaeontology (Awramik et al., 1983, 1988; Buick, 1984, 1988; Brasier et al., 2002; Dalton, 2002; Van Zuilen et al., 2002). To sharpen such discussions, this essay focuses only on the oldest evidence of life now known~that dating from between 3200 and 3500 Ma ago--and centres particularly on the crucial problem of differentiating true fossils from pseudofossil look-alikes, of separating the bonafide from the bogus.

Rules of the Hunt There are fine lines in science between what is known, guessed, and hoped-for, lines that are sometimes crossed. But sound science is not a guessing game. The goal is to know. "Possibly ... perhaps ... might be" are not firm answers and feel-good solutions do not count. For example, evidence of past life either exists, or it does not, in the famous meteorite from Mars, ALH84001 (McKay et al., 1996). Critical evaluation of the known facts will sort the matter out. But the controversy about the putative fossil evidence in ALH84001--still simmering in some quarters--might have been avoided had three lessons brought to light by this episode been heeded from the outset: (1) First, the search for evidence of past life, whether in Earth rocks or extraterrestrial samples, must be multidisciplinarymat a minimum, based on the techniques and findings of biologists, palaeontologists, geologists and geochemists--or, better still, interdisciplinary, carried out by researchers schooled in relevant aspects of both the life and physical sciences. (2) Second, the evidence sought should be positive, data that affirm the biological origin of the features detected. Evidence that is neutral, equally consistent with either biology or non-biology, is by its nature inadequate to prove the existence of past life. And interpretations based on the flawed logic of negative reasoning~inference by default, such as a claim that because a feature is not obviously mineralic it "must" be biogenic~are likely to be in error. (3) Third, to be acceptable, evidence of past life must meet five specific tests (Schopf and Walter, 1983; Schopf, 1992a, 1993), detailed below. (i) Provenancemis the source of the rock sample that contains the putative biologic features established firmly? In particular, the stratigraphic and geographic provenance should be known precisely, as demonstrated, for example, by replicate sampling by different workers. (ii) Age--is the age of the rock sample known with appropriate precision? For example, is the age tightly constrained by multiple measurements and other lines of evidence? (iii) Indigenousnessmare the putative biologic features indigenous to the rock? Specifically, are the features embedded in the rock matrix rather than being surficial contami-

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nants?--a test that can be met for minerals and fossils by studies of petrographic thin sections and for chemical signatures by analyses in situ of particulate, immobile, organic matter. (iv) Syngenicity--are the features syngenetic with a primary mineral phase of the rock? Or, on the contrary, are they of later origin?--for example introduced into pores or fractures and lithified by secondary or later-generation minerals, a test that can usually be met by detailed studies of petrographic thin sections. (v) Biogenicity--are the features assuredly biological? This test, almost always the most difficult to satisfy, provides the focus for much of the remainder of this essay.

Evidence of Ancient Life Four independent but potentially mutually reinforcing lines of evidence have been used to trace the earliest records of life: minerals, fossils, carbonaceous matter, and isotopic signatures. Of these, minerals have proven to be the least compelling, primarily because no known minerals are uniquely of microbial origin; biologic systems can promote deposition of diverse mineralic species (silica, carbonates, iron oxides, sulphides), but the same minerals can be formed abundantly by entirely non-biologic means. However, the other three biologic indicators--fossils, organic matter and isotopic signatures---especially if taken together, can provide unambiguous evidence of life. Of these, the strongest is the presence of unquestionable fossils: a mineral-replaced dinosaur bone or fossil log constitutes assured evidence of past life despite its non-carbonaceous composition and whether its isotopes fit biology or not. As a first approximation, therefore, the problem of tracing back life's history can be viewed as centring on a single critical question: how old are Earth's earliest fossils? To answer this question, two types of fossils, both present in the early rock record, need to be considered: stromatolites and cellularly preserved microorganisms.

Fossil stromatolites Formally defined, a stromatolite is an accretionary organosedimentary structure, commonly thinly layered, megascopic, and calcareous, produced by the activities of mat-building communities of mucilage-secreting microorganisms, filamentous and coccoid photoautotrophic prokaryotes such as cyanobacteria. Usage of the term, however, is somewhat variable. Some workers prefer to restrict "stromatolite" to geologic specimens and use "microbial mat" for their modern counterparts. Others term structures stromatolites only if they have fairly high relief above the neighbouring substrate and refer to flat-lying examples as mats or sheets. But common to all appropriate definitions is the concept that it is the biologic origin of the layering in stromatolites that makes them distinctiveuthe term stromatolite is properly applied only to such structures that are known or inferred to be biogenic (for contrasting points of view, cf. section 6.5). Because most stromatolites are calcareous, most do not contain structurally identifiable remnants of the microbes that built them, the organic cells having been crushed between growing carbonate grains during diagenesis and lithification. Hence, like a track, trail, or burrow preserved in an ancient sediment, stromatolites are classed as trace fossils,

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organosedimentary structures that evidence biologic activity yet are themselves not fossilised organisms. For a more thorough discussion of stromatolites and their environmental and evolutionary implications, see section 6.5.

Cellular fossils Optical microscopic studies of two types of preparations have been used to detect remnants of the minute body fossils of cellular microbes: acid macerations and petrographic thin sections. Transmission and scanning electron microscopy have also been used to investigate such fossils, but though both have provided useful information about microfossils previously detected in maceration or thin section preparations, neither has proven reliable as an initial detection technique. Optical microscopy of macerations Maceration, the easier and faster of the two principal detection techniques, is carried out by dissolving a rock in mineral acid (hydrochloric acid for limestones, hydrofluoric for cherts and siltstones). Because of their kerogenous composition, organic-walled microfossils pass through the process essentially unaltered. Abundant fossils are concentrated in the resulting sludge-like acid-resistant residue, which can be slurried onto a microscope slide for study. Such macerations, however, are notoriously subject to contamination: microbes adhering to even carefully cleaned rock surfaces; living contaminants (bacteria, cyanobacteria, microscopic algae and fungi) in laboratory water and commercially available mineral acids; and an almost limitless array of fossil-like objects introduced during transfer of the residue onto microscope slides-spores, pollen grains, lint or paper fibers, even bits of small spiders trapped in water pipes. From the 1950s into the 1970s, as active studies of the Precambrian fossil record were beginning in earnest, all of these maceration-borne contaminants were misinterpreted as fossils by one worker or another (Schopf and Walter, 1983; Mendelson and Schopf, 1992). Optical microscopy of petrographic thin sections In petrographic thin sections, the other type of preparation used for optical microscopic detection of ancient microfossils, permineralised (petrified) fossils are identified in situ, thoroughly encased within the rock, so indigenousness can be demonstrated and the possibility of laboratory contamination ruled out. Similarly, because thin sections can provide clear evidence of the relationship between fossil-like objects and their encompassing mineral matrix, their study can answer the question of whether the objects are coeval with a primary mineral phase rather than being of later origin. However, despite the obvious benefits of thin section studies, some workers have preferred to focus their hunt for ancient fossils on more easily prepared acid-resistant macerations. For the Proterozoic Precambrian, where the fossil record is well enough known that misidentification of contaminants and fossil-like artefacts can be avoided, the fast, simple, maceration technique can be highly useful. But to avoid mistakes in the Archaean, where the fossil record is not nearly so well known, use of the more rigorous thin section technique is essential.

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Electron microscopy as a detection technique

Although both transmission and scanning electron microscopy (TEM and SEM, resp.) have been used to characterise the cellular morphology of Precambrian microfossils, neither has proven to be a reliable detection technique. The TEM studies have involved examination either of ultra-thin sections of epoxy-embedded organic-walled fossils or of fossil-like structures detected in plastic ("formvar") surface replicas of polished and etched petrographic thin sections. Studies of ultra-thin sections of Precambrian microfossils have yielded useful insight into the fine structure of their cell walls, membranes, and internal organic contents (Schopf and Oehler, 1976), but because such studies are based on contamination-prone macerations, this approach should be used only to characterise fossils known also from thin sections. Contamination similarly presents a problem for TEM studies of surface replicas, as does the introduction of diverse non-biogenic artefacts (blisters, bubbles, strands of formvar, and so forth), that in the 1960s and early 1970s were repeatedly misidentified as "ancient fossils" (Schopf and Walter, 1983). Preparation of samples for study by SEM is simpler than for TEM and the images acquired are generally easier to interpret. Nevertheless, experience has shown that establishing by SEM that the objects detected are both indigenous to the rock examined and syngenetic with first-phase minerals can be virtually impossible. Moreover, as in studies by TEM, diverse non-biologic artefacts have been misinterpreted as fossils by SEM, especially those occurring in rock samples in which weathering (or acid-etching during sample preparation) has altered and smoothed mineralic morphologies into biologic-like shapes (Schopf and Walter, 1983). In sum, neither TEM nor SEM has proven to be a reliable technique for the detection of microfossils in the Precambrian rock record. In the absence of confirming data from studies of petrographic thin sections, such reports should be regarded with scepticism. Biogenicity

Though optical studies of petrographic thin sections can answer the questions of indigenousness and syngenicity, the problem of establishing biogenicity often remains. Here, too, mistakes have been made--"life-like" dust particles, ball-shaped mineral grains, clumps and shreds of compressed coaly kerogen, solid opaque globules, and a wide variety of other microscopic objects have all been claimed to be Precambrian fossils, often on the basis of only one or a few specimens and despite the absence of identifiable cells or other features distinctive of living systems (Schopf and Walter, 1983; Mendelson and Schopf, 1992). Too often such reports have been based on the notion that because an object is not obviously mineralic, it "must" be a fossil. Such reasoning is flawed. A similarly spurious argument is the suggestion that putative Archaean fossils "should not be accepted until all possibilities of their non-biological origin have been exhausted" (Brasier et al., 2002). Here, the problem stems not so much from the concept itself (though it makes little sense to include "all possibilities", whether they are truly plausible or not), but from the way it has been applied, a strategy of casting doubt on the biological origin of the objects in question without showing how they actually formed (Brasier et al., 2002). To propose plausible

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alternative hypotheses is a hallmark of science; but to raise doubt and then beg the question furthers understanding not at all. A constructive and far better approach to the problem of establishing biogenicity is to insist that claims of ancient fossils be backed by data that show what the fossil-like objects actually aremrather than what they seemingly might be or apparently are not--backed by evidence sufficiently strong to rule out plausible non-biologic sources. The key lies in positive data that affirm a biological origin: a suite of traits unique to life that are not shared by inanimate matter. And it is crucial to consider not one, not another, but a suite of such traits, since considered by themselves, any single trait may prove misleading. For example, because organic matter can be produced by non-biologic processes (as when life originated, or at present in interstellar space), the mere presence of coaly particles in an Archaean sediment is insufficient to establish the existence of life. And because unicelllike organic spheroids can form in the absence of biological systems (from clumping of organic particles in seawater, or by organic matter coating ball-shaped mineral grains), tiny kerogenous spheroids cannot in and of themselves be regarded as assured fossil cells. Probably the best way to avoid being misled by non-biologic structures is to accept as true fossils only those of fairly complex form. This may seem an unreasonably stringent rule for truly ancient fossils since the earliest cellular life must certainly have been quite simple, probably individual tiny sac-like spheroids. But until we have a sounder base of knowledge and better rules to separate non-fossils from pseudofossil mimics, we should err on the side of caution. Special care of this sort may no longer be necessary for reports of fossils from the Proterozoic where evidence of life is overwhelming, but for the Archaean, where so little yet is known, demanding rules must still apply. At present, it is probably best to accept as assuredly biologic only those putative Archaean fossils that are morphologically sufficiently complicated to be unquestionably of biological form, for example, colonies of coccoidal cells embedded in a surrounding, distinct, organic envelope, or sinuous threadlike filaments made up of chains of numerous more or less identical cells.

Morphology of spheroids Unlike minerals, the cells of biologic systems, composed of pliable organic substances, tend to eschew sharp angles. Yet because mineral grains are not uncommonly of microbial size, and because the originally angular edges and corners of such grains are routinely rounded off by erosion and subsequent transport, spheroidal "fossil-like" mineralic granules can be confused easily with true coccoidal fossils. Criteria in addition to simple spheroidal form must therefore be used to discriminate true fossils from such non-biologic mimics. Foremost among such criteria are those derived from the genetics of biologic systems, for spheroidal microbes the genetically determined and therefore limited range of characteristic morphologies they exhibit which define a discrete (and for modern microorganisms, well established) region of morphospace. Thus, a diverse assemblage of true coccoidal microbial fossils can be expected to be composed of microscopic objects that: (1) are spheroidal, or prior to preservation were initially spheroidal, their shape being determined by a robust and commonly easily discernible cell wall;

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(2) have a finite range of cell diameters (which if like that of modern prokaryotes, extends from < 0.5 ~tm to c. 60 ~tm, most commonly 5.0 ~tm or less) among members of morphologically defined groups of specimens that exhibit the same or similar limited range of a given suite of morphological characters (morphospecies), encompassing a definable, limited, size range (Schopf, 1992b); (3) may exhibit "lima bean-shaped" cells in pairs in which the flatter sides of the two cells are adjacent, an organisation reflecting their derivation from the division of a common parent cell; (4) if planktonic and unicellular, are preserved consistently in the sedimentary record as scattered, isolated, individual spheroids; (5) if benthic and colonial, are commonly enclosed by an originally mucilagenous (diffuse or diaphanous to well-defined, thin to thick, single-layered or laminated) organic envelope and usually exhibit a relatively flat lower surface (where the colony propagated across a substrate) and a more bulbose upper surface (where expansion of the colony was physically less constrained); and (6) include multiple examples of the various morphospecies (if one example of a particular taxon can be preserved, others should be, too) in varying states of preservation, ranging from specimens that are disrupted and degraded, to those that are flattened, folded, or otherwise partially altered, to those that more closely resemble their original morphologies.

Morphology offilaments In principle, suites of non-biological mineralic structures that mimic convincingly the morphological characteristics of a diverse assemblage of septate microbial filaments must be exceedingly rare (if they exist at all). Microbial filaments are cylindrical, not planar, so they should not be confused easily with mineralic veinlets (probably the most common "filament-like" morphology present in petrographic thin sections). And, like spheroidal microorganisms, because filamentous microbes exhibit a genetically determined limited range of characteristic morphologies, they, too, occupy a discrete and definable region of morphospace that can provide a key to their identification. In particular, a diverse assemblage of septate microbial fossil filaments can be expected to be composed of microscopic objects that: (1) are cylindrical, or prior to preservation were initially cylindrical, their shape being defined by robust lateral cell walls; (2) have a finite range of widths (which if like that of modern microbes extends from < 0.5 ~tm to c. 100 ~tm, most commonly 10 ~m or less) and a finite range of lengths, typically ranging from a few tens to a few hundred microns (Schopf, 1992b); (3) are divided into segments of more or less uniform size by rather regularly spaced cross-walls (transverse septa) that define their cellular subunits and may exhibit cells in pairs or bifurcated (partially divided) cells that reflect the original presence of partial septations formed as a prelude to cell division;

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(4) have a rather uniform diameter throughout their lengths, but may systematically taper towards (especially relatively close to) their apices, and terminate in morphologically distinctive (e.g., flat, rounded, hemispherical, conical, pillow-shaped) end cells; (5) among members of groups of specimens having similar morphologies (morphospecies), show a systematic correlation of cell length and width that spans a limited and definable size range; (6) due to their original flexibility, exhibit varying degrees of sinuosity that tend to be correlated with filament length--shorter filaments tending to be less sinuous and to have fewer flexures; longer filaments commonly being more sinuous and having more flexures (just as a longer string can be wound into a greater number of coils than a shorter one); (7) due to original variations in flexibility, exhibit varying degrees of sinuosity that tend to be correlated with filament widthmbroader filaments, especially those having equant cells, tending to be less sinuous and to have fewer flexures; narrower filaments tending to be more sinuous and to have more flexures (just as a length of string can be coiled into a tighter packet than a similar length of rope); (8) include multiple examples of individuals that exhibit the same or similarly limited range of a definable suite of morphological characters (if one example of a particular taxon can be preserved, others should be also); and (9) include multiple examples of members of such morphologically defined groups (morphospecies) in varying states of preservation, ranging from specimens (or parts of specimens) that are degraded to narrow strands or "ballooned" into elongate sac-like shapes, to those that are disrupted, broken into partially preserved fragments that show only hints of cellularity, to better preserved specimens that more closely resemble their original morphologies. Carbonaceous matter

Chemical extraction, isolation, and identification of organic biomarkers, particularly of various types of hydrocarbons, has provided useful insight into the evolutionary history of Proterozoic life (Summons, 1992). For example, identification of the protozoan biomarker tetrahymenol in c. 930 Ma sediments (Summons, 1992), supported by the detection of fossil testate amoebae in the same sedimentary sequence (Schopf, 1992c; Porter and Knoll, 2000), has established a minimum age for the Proterozoic emergence of protozoan protists. With but few exceptions, however, biomarker studies have not proven applicable to the Archaean. The single most promising such report to date is that of the detection of steranes (hydrogenated derivatives of sterols like those typical of eukaryotic cellular membranes) in c. 2700 Ma sediments, interpreted as evidencing the Archaean existence of cellular eukaryotes (Brocks et al., 1999). But even this report is subject to question since all organic compounds are soluble to some extent in groundwater, and can therefore be introduced into a sedimentary unit long after its deposition, and since no means have been devised to directly date the geological age of extracted organics. The basic problem is that in the Archaean, unlike the Proterozoic, richly petroliferous deposits are unknown, sediments in which indigenous hydrocarbons are sufficiently abundant to "swamp out" backgroundlevel contamination.

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In contrast to extractable biomarkers, the particulate organic constituents of ancient sediments, finely divided particles of carbonaceous kerogen, are immobile, locked in place within their surrounding mineral matrix. Because such kerogens are thus demonstrably syngenetic with the encompassing minerals, and because they comprise the bulk of the carbonaceous constituents present in sedimentary rocks, even those that are richly petroliferous, most analyses of Precambrian organic mattermand virtually all studies of that in Archaean-age depositsmhave focused on the chemistry of kerogen. Two types of analyses have proven especially useful: laser-Raman spectroscopy of kerogenous materials and mass spectrometric measurements of carbon isotopic compositions. A third analytical technique, atomic force microscopy, has recently provided means to elucidate the submicronscale structure of Precambrian carbonaceous matter.

Raman spectroscopy of individual microfossils Since the first discoveries of permineralised (petrified) three-dimensionally preserved Precambrian microorganisms (Tyler and Barghoorn, 1954; Barghoorn and Tyler, 1965; Cloud, 1965; Barghoorn and Schopf, 1965) and recognition of the problems involved in distinguishing such fossils from mineralic pseudofossil look-alikes (e.g., Tyler and Barghoorn, 1963), it has been evident that a need exists to directly correlate the chemical makeup of individual fossil-like objects with their optically discernible morphology. In particular, evidence consistent with a biogenic interpretation would be provided were chemical data to show that such putative microfossils were composed of carbonaceous matter, as would be expected of permineralised organic-walled microorganisms, and would be virtually unequivocal were the analyses to demonstrate that such carbonaceous matter constituted "kerogen", geochemically altered organic matter of biologic origin. This need has recently been met by laser-Raman imagery, a technique new to palaeobiology that permits a one-to-one correlation of carbonaceous molecular-structural makeup and discernible cellular morphology in fossil specimens as minute as 1 ~tm in size (Kudryavtsev et al., 2001; Schopf et al., 2002). Non-intrusive and non-destructive, this technique can be used to analyse fossils situated within petrographic thin sections (to depths > 60 ~tm), whether the sections have been polished or are unpolished and covered by a thin veneer of microscopy immersion oil (Schopf et al., 2002), as well as fossils freed from a rock matrix by acid maceration (Schopf et al., in preparation). However, laser-Raman chemical analysis of ancient organic matter is not in and of itself a panacea. As noted above, because organic matter can be produced by non-biologic processes, as it unquestionably was before life emerged, the mere presence of coaly particles in an Archaean sediment is insufficient to establish the existence of life. Moreover, geochemical maturation of almost all naturally occurring organic substancesmwhether they are of biological origin or are products of solely non-biological processes---can be expected to give rise to the same set of derivative products, polycyclic aromatic hydrocarbons that under geological conditions can become increasingly graphitised and, ultimately, converted to crystalline graphite (Pasteris and Wopenka, 1991; Wopenka and Pasteris, 1993). That such carbonaceous matter is in fact kerogenous (i.e., of biological origin) can, of course, be established if it is shown to constitute the cell walls or other identifiable struc-

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tures of well-preserved fossils (in which instance it is the mode of occurrence of the organic matter, not its chemically distinctive characteristics, that evidences its biogenicity). And kerogen can be readily identified in little-metamorphosed fossiliferous Proterozoic units in which the biogenicity of the carbonaceous matter is established, not only by its occurrence both in the fossils and the associated sapropelic detritus, but by its distinctively kerogenous laser-Raman signature (Kudryavtsev et al., 2001). But microfossils that could be called well preserved are unknown from older Archaean-age deposits, and all such units analysed to date (Schopf et al., 2002; Schopf et al., in preparation) have been metamorphosed well beyond a stage at which the biogencity of their carbonaceous components could be established on the basis of laser-Raman chemistry, alone. By itself, therefore, laser-Raman spectroscopy of geochemically altered organic matter can establish only its carbonaceous makeup, not its biogenicity. Nevertheless, combined with evidence from optically discernible cellular morphology, or with that showing that the organic matter analysed is a part of a continuum that extends from the carbonaceous constituents of relatively "well preserved" to "poorly preserved" microbial fossils (Schopf et al., in preparation), laser-Raman studies can provide convincing evidence of a biological origin. Here, laser-Raman imagery, rather than Raman spectrometry per se, has proven particularly powerful (Kudryavtsev et al., 2001; Schopf et al., 2002). In essence, this technique produces high resolution virtual "chemical maps" that show the spatial distribution of the polycyclic aromatic kerogenous components of the fossils analysed (Figs. 6.2-1 and 6.2-2) and of any nearby sapropelic organic materials. By doing so, it not only provides direct evidence of the carbonaceous makeup of such fossils, but permits comparison of the chemistry of the kerogen of which they are composed with that present in the encompassing matrix. Because such comparison can show whether the fossils and the associated particulate kerogens are coeval (having had identical geochemical histories and, thus, the same or very similar molecular compositions) or are of differing origins and histories (as, e.g., would be evidenced by comparison of the chemistry of kerogenous sapropelic matter emplaced during sedimentation with that of the organic components of non-syngenetic endolithic microorganisms), laser-Raman imagery provides means to address the indigenousness and syngenicity of putatively ancient fossil-like objects. Similarly, such analyses can yield definitive data regarding the extent of geochemical alteration of such kerogens, providing evidence by which to investigate the products and processes of the geochemical maturation of ancient organic matter (Schopf et al., 2002; Schopf et al., in preparation). Atomic force microscopy of individual microfossils Although atomic force microscopy (AFM) is a technique used to elucidate the submicronscale fine structure rather than the chemical composition of materials analysed, it is here considered together with chemical analytical techniques because of the insight it can provide into the structural makeup of the kerogenous constituents of petrified Precambrian microorganisms. Unlike laser-Raman imagery, by which fossils situated within petrographic thin sections can be analysed, AFM like SEM, is a technique that permits analyses only of petrified fossils exposed at the upper surfaces of polished thin sections. To carry out such studies, the fossils must first be exposed in bas-relief by gently etching the sections

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in a dilute solution of mineral acid. The surfaces are then systematically scanned by AFM at high resolution to obtain virtual images that map the topography of the exposed cell walls and their component subunits. Such analyses typically take several hours to complete and at the present state-of-the-art produce a relatively low (c. 20%) yield of usable results (Kempe et al., 2002), a yield that can be expected to increase, perhaps markedly, as more experience is gained with the use of this new technique. To date, AFM results, coupled with laser-Raman analyses, have shown that the kerogenous walls of Precambrian petrified microfossils are composed of stacked, platy, angular polyaromatic hydrocarbon subunits, c. 200 nm in size, oriented perpendicular to the cell walls (Fig. 6.2-3; Kempe et al., 2002). Because of this evidently characteristically biologic organisation, AFM, backed by laser-Raman imagery, holds promise for differentiating true fossils from non-biologic pseudofossils (Kempe et al., 2002).

Isotopic signatures Stemming from the pioneering studies of Park and Epstein (1963) and Hoering (1967), an impressive amount of data has been accumulated on the carbon isotopic compositions of coexisting inorganic carbonate minerals and organic kerogens in Precambrian fossiliferous sediments (Strauss and Moore, 1992). For kerogens, most such data have been obtained on the total organic carbon fraction of selected rock specimens and on carbonaceous residues concentrated by acid maceration of large (kilogram-sized) rock samples, most notably as a result of two extensive studies carried out by the "Precambrian Paleobiology Research Group" (Hayes et al., 1983; Strauss et al., 1992b). Recently, House et al. (2000) introduced a new technique, measurement of the isotopic composition of the kerogenous matter that comprises the cell walls of individual petrified Precambrian microscopic fossils by means of ion microprobe spectrometry, an approach that has now also been used by Ueno et al. (2001) to determine the carbon isotopic makeup of filamentous microfossils c. 3490 Ma in age. Analyses of bulk samples and of individual microfossils To date, hundreds of measurements have been made of the isotopic compositions of the kerogenous components of bulk samples of diverse fossil-bearing shales and cherts (Strauss and Moore, 1992). These analyses, that in comparison with values measured on

Fig. 6.2-1. Optical photomicrographs, interpretative drawings, Raman images, and Raman spectra of four permineralised specimens of Primaevifilum amoenum (a-d), a filamentous cyanobacterium-like prokaryote, in petrographic thin sections of the c. 3,465 Ma old Apex chert of northwestern Western Australia (Schopf, 1992a, 1993). The carbonaceous composition of these fossils is established by the two prominent vibrational bands in the Raman spectra (at c. 1,350 cm-! and c. 1,600 cm -1 ), acquired from kerogen having the spatial distribution shown for each specimen in the accompanying Raman images, as measured by the techniques summarised by Kudryavtsev et al. (2001) and Schopf et al. (2002). (a) The Natural History Museum (NHM), London, specimen No. V.6316415]. (b) NHM specimen No. V.631 64[6]. (c) NHM specimen No. V.6372812]. (d) NHM specimen No. V.63166[ 1].

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coexisting carbonates yield a consistent signature of isotopic discrimination between the two carbon reservoirs of 25 + 10%0 (interpreted as evidencing biological photoautotrophic carbon fixation; Hayes et al., 1983; Strauss et al., 1992b), can be traced to at least as early as c. 3500 Ma (Fig. 6.2-4). Such analyses, however, are time-consuming and labour intensive, and by their nature can yield only an average value of the components measured, data providing strong evidence of the existence of photoautotrophic primary producers, but no means to either differentiate between the contributions of oxygenic and anoxygenic photosynthesisers or link carbon isotopic composition to particular biologic species of fossil microbial communities. Although this latter deficiency, the inability to link isotopic makeup to particular taxa, is obviated by the new technique of ion microprobe spectrometry of individual microscopic fossils, this approach, too, is tedious and technically difficult, l.,ike atomic force microscopy, isotopic measurements by use of an ion microprobe require that the specimens analysed be exposed at the upper surfaces of polished petrographic thin sections. For specimens densely packed in richly fossiliferous deposits (House et al., 2000), especially those that are relatively large (> 20 ktm in size), this requirement can be met rather easily. But in sparsely fossiliferous assemblages, particularly those composed chiefly of narrow (c. 1 ktm diameter), sinuous, threadlike filaments, identification of the fossils at the section surface by use of the low resolution optical microscopic apparatus typical of ion microprobes can be difficult. Moreover, because of instrument-caused systematic drift of measurements obtained, attainment of the + 1%o precision desired of such analyses can require repeated comparative measurements of the sample and an appropriate standard of known isotopic composition, making such analyses exceedingly time-consuming. Nevertheless, the consistency between the measurements acquired by ion microprobe spectrometry and those obtained by traditional bulk sample analyses (Fig. 6.2-5) well demonstrate the efficacy of this new technique. The Early Archaean fossil record

Although the documented Archaean fossil record is minuscule in comparison with that of the Proterozoic, it is substantially better known and more diverse than is generally appreciated. Eight particularly ancient fossil-bearing deposits, c. 3200-3500 Ma in age, have been

Fig. 6.2-2. Optical photomicrographs (and an interpretive drawing), Raman images, and Raman spectra of especially ancient permineralised kerogenous microfossils in petrographic thin sections of Archaean cherts. The carbonaceous composition of these fossils is established by the two prominent vibrational bands in the Raman spectra (at c. 1,350 cm-! and c. 1,600 cm-I), acquired from kerogen having the spatial distribution shown for each specimen in the accompanying Raman images, as measured by the techniques summarised by Kudryavtsev et al. (2001) and Schopf et al. (2002). (a) Unnamed sheath-enclosed colonies of coccoidal microbial cells from the c. 3,420 Ma old Strelley Pool Chert of Western Australia (Schopf and Packer, 1987; Schopf et al., in prep.). (b) Unnamed filamentous microbial fossil from the c. 3,375 Ma old Kromberg Formation of South Africa (Walsh and Lowe, 1985; Schopf et al., 2002).

530

Chapter 6: Evolution of Life and Precambrian Bio-Geology

Fig. 6.2-3. Optical photomicrographs (a, b), atomic force images (c-g), and computer-generated sketch (h) of parts of a permineralised spheroidal Proterozoic microfossil, exposed at the upper surface of a petrographic thin section of stromatolitic chert, from the c. 650 Ma old Chichkan Formation of southern Kazakhstan (Kempe et al., 2002). (a) Specimen transected at its equatorial plane by the surface of the thin section, superimposed by a plot of an electron microprobe measurement showing that carbon is concentrated in and near its cell wall. (b) The lower half of the specimen, embedded in the chert matrix, showing the characteristic surface texture of this taxon of Proterozoic microfossil. (c) Atomic force micrograph showing the exposed cell wall analysed. (d) The portion of the wall imaged in e-g and illustrated in h, a segment of the wall shown in the lowermost part of c.

described, containing in toto both stromatolites and spheroidal and filamentous microfossils supported by both laser-Raman and carbon isotopic analyses of their kerogenous components (Fig. 6.2-6). Other true fossils have been reported from the Archaean (Schopf and Walter, 1983; Lanier, 1986; Klein et al., 1987; Altermann and Schopf, 1995; Kazmierczak and Altermann, 2002) as have numerous fossil-like objects of possible but uncertain biogenicity (Schopf and Walter, 1983; Naqvi et al., 1987; Venkatachala et al., 1990; Walsh, 1992), but the following review focuses only on these eight oldest known units that, when taken together, set a minimum age for the known antiquity of life.

6.2. Earth's Earliest Biosphere

531

Fig. 6.2-3 (continued). (e) That portion of the wall indicated by the white rectangle in d, shown here at higher magnification. (f) That portion of the wall indicated by the white rectangle in e, shown here at higher magnification. (g) That portion of the wall shown in f with a superimposed computer-generated outline of its component parts. (h) A computer-generated sketch showing the c. 200 nm-sized stacked angular platelets of polycyclic aromatic hydrocarbons that make up the organically preserved cell wall.

Table 6.2-1 summarises the salient data now available regarding these eight especially ancient deposits. As indicated in this table, diverse evidence of life has been reported from both of the only two thick stratigraphic successions known from this segment of the Archaean, the Pilbara Craton of northwestern Western Australia and the Barberton Mountain Land in and near Swaziland in southeastern South Africa. Fossils from each of the eight units pass the tests required of bona fide evidence of Archaean life--being of assured provenance, age, indigenousness, syngenicity and biogenicity--although the exact stratigraphic source of one of the two reports of microfossils from the Dresser Formation of Western Australia (Awramik et al., 1983; Awramik, 1992a) remains uncertain, a fossilbearing horizon that has yet to be relocated (Awramik et al., 1988; Schopf, 1999); both the biogenicity and syngenicity of certain of the fossil-like objects described in one of the three reports from the Kromberg Formation of South Africa (Westall et al., 2001 a) are open to question (Altermann, 2001).

532

Chapter 6: Evolution of Life and Precambrian Bio-Geology

Fig. 6.2-4. Carbon isotopic values of coexisting carbonate and organic carbon measured in bulk samples of Phanerozoic and Precambrian sedimentary rocks; for the Precambrian, represented by data from 100 fossiliferous cherts and shales, shown as average values for groups of samples from 50 My long intervals (Strauss and Moore, 1992).

In environmental setting, the eight fossiliferous units (Table 6.2-1) range from shallow marine to assuredly (Sulphur Springs deposit) or possibly hydrothermal (Apex chert: Van Kranendonk, in press; chert veins of the Dresser Formation: Ueno et al., 2001). Stromatolites from the two terrains include fiat-lying, domical, columnar, and conical forms, and the microfossils detected include isolated single cells; paired, evidently dividing cells; ensheathed colonies of coccoidal cells (Fig. 6.2-2a); and both cylindrical and cylindricalseptate filaments that range from < 1 to c. 20 lam in diameter (e.g., Figs. 6.2-1, 6.2-2b). For fossils in three of the depositsmthe Apex chert (Fig. 6.2-1) and the Strelley Pool chert and Kromberg Formation (Fig. 6.2-2)mlaser-Raman imagery has established a one-to-one correlation of kerogenous composition and cellular morphology, and Raman spectra have also been acquired for those in one other deposit, narrow branched(?) filaments in chert veins of the Dresser Formation (Ueno et al., 2001). Carbon isotopic values ranging from - 2 7 to -32%0, well within the range typical of Precambrian biogenic kerogens (Fig. 6.2-4), have been measured on the carbonaceous components of seven of the deposits (Fig. 6.2-7), with comparable values having been measured by ion microprobe spectrometry on individual fossils of the c. 3490 Ma Dresser Formation (Fig. 6.2-5), the oldest of the eight fossil-bearing units. The most thoroughly studied fossils of these various assemblages are the filamentous microbes reported from the second oldest of the eight units, the c. 3465 Ma Apex chert of Western Australia (Schopf and Packer, 1987; Schopf, 1992a, 1993): (1) Morphometric studies of c. 2,000 cells in c. 180 specimens has established a systematic correlation between cell shape, filament diameter, and taxon-specific terminal cell morphology for each of the 11 filamentous species described from the unit (Schopf, 1993).

Table 6.2-1. Oldest evidence of life Geology A ge (Ma) Unit

Locality

hl P

Palaeontology Stromatolites Shape Environment

Microfossils Morphology

Environment

- -geochemistry Organic .Ram:," AVE. - s ' . ' c P.~.-R.. ('%c) Laser-],.-...-.. Bulk Fossils Spectra Images

c. 3,490

Dresser Formation

Western Australia

Domicall

Shallow marine1

Diverse filaments' Narrow filaments3

Shallow marine2 J4 ~ ~ d r o t h e r m a l ~ -31.5

c. 3,465

Apex chert

Western Australia

-

-

Diverse

~~drothermal~

(11

c. 3,458

c. 3,430

J3 -36.4 = 18) (n = 9)

J4

J3

-

R J

~'8

2

-27.7 (n = 10)

Hooggenoeg South Africa Formation

-

Strelley Pool Western chert Australia

Conical1

Kromberg Formation

South Africa

Swartkoppie Formation

South Africa

Fig Tree Group

South Africa

Sulphur springs deposit

Western Australia

-

Rods and coccoidsln Shallow marinelo

J4

-

-

-

-

-31.8 (n = 16) Shallow marine1I

Colonial coccoids" l 2

shallow marine5

Wavy-flat laminated''

Shallow marine1'

Diverse filaments8, l 3 Shallow marine'' Colonial c o c c o i d ~ ' ~ Shallow marine14 Colonial c o c c o i d ~ ' ~ )Shallow marine''

-

-

Dividing c o c c ~ i d s ~Shallow ~ marinels

J4

-32.3 (n = 5)

c. 3,375

c. 3,340

c. 3.300

c. 3,235

J4

-29.6 (n = 29) J4

-27.2 (n = 15) ~olumnarl~ Shallow marine16

Solitary coccoids17

-

Narrow

-

Shallow marinei7

J4

-28.8 (n = 31) Hydrothermal18

f Packer, 1987; 'schopf, 1992a; 7 ~ c h o p f 1993; , 8~chopf I ~ a l t e et r al., 1980; '~wramik et al., 1983; ' ~ e n oet al., 2001; 'strauss and Moore, 1992; s ~ c h o p and et al., 2002; ')van Kranendonk, in press; 1 0 ~ e s t a let l al., 2001a; "Hofmann et al., 1999; IZschopf et al., in preparation; I3walsh and Lowe, 1985; I4h.luirand Grant, 1976; l s ~ n o land l Barghoorn, 1977; 1 6 ~ y e r l et y al., 1986; I7~arghoornand Schopf, 1966; IB~asmussen, 2000.

UI W

534

Chapter 6: Evolution of Life and Precambrian Bio-Geology

Fig. 6.2-5. Carbon isotopic values of individual Precambrian microfossils measured by ion microprobe spectrometry compared with those of the carbonate and total organic carbon measured in bulk samples of the same geologic units. Values plotted for carbonate and total organic carbon are from Strauss and Moore (1992); for microfossils from the Bitter Springs and Gunflint Formations, from House et al. (2000); and those for microfossils from the Dresser Formation, from Ueno et al. (2001).

(2) The sinuous three-dimensional morphology of the cylindrical filaments and the correlation among the taxa of filament width and degree of sinuousity establish their original flexibility (Schopf, 1992a, 1993).

6.2. Earth's Earliest Biosphere

535

OLDEST EVIDENCE OF LIFE

STROMAT- MICROOLITES FOSSILS (~13CpD B RAMAN

3500tl bDresserFm i IhApex chert

~

! HooclqenoegFm / bStrelleyPoolChert V

'~

~

'~

~~/

'V~ V/ V

'~

V

/

V

3400 Age

| bKrombemFm

(Ma) |_ bSwartkoppieFm 3300"-~ ~Fig Tree Grp

v v v

v

i~ SulphurSprings 3200

--

Fig. 6.2-6. Temporal distribution of the principal evidence for early life known from geologic units 3,200-3,500 Ma in age. (See Table 6.2-l for references.)

(3) Both the uniseriate, chain-like cellular organisation of the fossils and the presence of paired and bifurcated cells, reflecting the original presence of transverse partial septations, demonstrate that the filaments grew by a process of cell division like that typical of living filamentous microbes (Schopf, 1993). (4) The organismal organisation exhibited by all members of the assemblage is comparable to that of fossils known from other deposits of similar age (Table 6.2-1). (5) And not only has laser-Raman imagery established the carbonaceous, kerogenous composition of numerous individual specimens (Schopf et al., 2002), evidencing their original organic makeup, but isotopic measurements of the Apex kerogen have demonstrated its distinctive biological signature (Fig. 6.2-7). Despite this evidence, it has been suggested that rather than being fossilised filamentous microbes, the Apex fossils might be the degraded remnants of ensheathed colonies of coccoidal cyanobacteria (Kazmierczak and Kremer, 2002) or even be non-biologic artefacts of entirely abiotic origin (Brasier et al., 2002). Though such proposals merit consideration, they not only do not fit with, but are contraindicated by the known evidence, and are suggestions that in large measure stem from a problem common to all of the natural sciences--one well known in radioastronomy, for example, where it is referred to as "separating the signal from the noise". Filaments claimed by Brasier et al. (2002) to be branching, rather than unbranched as originally described (Schopf and Packer, 1987; Schopf, 1992a, 1993), are misinterpreted instrumental artefacts (Fig. 6.2-8)mnot branching filamentsmwhereas other supposedly branching filaments reported by Dalton (2002)

536

Chapter 6: Evolution of Life and Precambrian Bio-Geology

Fig. 6.2-7. Carbon isotopic values of carbonate and organic carbon measured in bulk samples of the seven oldest microfossiliferous units now known (Strauss and Moore, 1992; Brasier et al., 2002).

to have been observed by Packer are chalcedonic mineral aggregates of demonstrably nonbiologic origin (Fig. 6.2-9). For the Apex assemblage, as is always true in analyses of fossil microbial communities (e.g., Knoll et al., 1988), the crucial problem has been to differentiate between true fossils--cellular organic remnants of assured biologic origin--and the relatively nondescript charred bits and pieces of such fossils and/or aggregated clots of degraded more or less amorphous kerogenous organic matter that are also present in the deposit. Though this "signal-to-noise" problem is particularly acute for the Apex fossils, because of the decidedly poor preservation of the assemblage (Schopf, 1993, 1999), it has been solved effectively by the detailed analyses of fossil morphologies summarised above, coupled with Raman imagery, establishing the carbonaceous makeup of the preserved microbes (Schopf et al., 2002), and by isotopic analyses of the Apex kerogen, showing that it, like the sapropelic carbonaceous components of all other fossil-bearing deposits of about the same age (Fig. 6.2-7), has the distinctive signature of biology. In short, both the biogcnicity of the Apex fossils and theh" identitication a~ cellular remnants of filamentous microbes are firmly established. This assemblage, together with those of the seven other especially ancient fossil-bearing units now known (Fig. 6.2-6, Table 6.2-1), sets a minimum age for the antiquity of life on Earth.

6.2. Earth's Earliest Biosphere

537

Fig. 6.2-8. Optical photomicrographs, interpretative drawings, Raman images, and Raman spectra of unbranched filamentous fossil microbes in petrographic thin sections of the c. 3,465 Ma old Apex chert of northwestern Western Australia (a, c) compared with the digitally produced automontage images of the same specimens (b, d) misinterpreted by Braiser et al. (2002, their Figs. 2d, g) as illustrating "branching". (a, b) Archaeoscillatoriopsis disciformis (Natural History Museum No. V.6316518]); "optical sectioning" at three successive focal depths (a) demonstrates that the fossil filament is recumbently folded rather being "branched" as it is depicted in the digital image (b, arrow), a computer-generated montage of images acquired at varying depths which misrepresents the morphology of the specimen. (c, d) Pimaevifilum delicatulum (Natural History Museum No. V.6316512]); the optical photomicrograph in c shows the specimen to be uniseriate and unbranched, indicating that the arcuate "side branch" depicted in the digital image (d, arrow) is not contiguous with the fossil, but is an object situated deeper within the thin section that has been brought into focus by use of the automontage system.

538

Chapter 6: Evolution of Life and Precambrian Bio-Geology

F~g. 6.2-9. Optical photomicrographs ot supposedly toSSll-llKe oojects composed of arcuate cusps formed by crystallisation of radiating fibres of chalcedonic silica, photographed at UCLA (l 6 January 1986) by B.M. Packer and erroneously reported by Dalton (2002) as representing branching fossil filaments, in petrographic thin sections of the c. 3,465 Ma old Apex chert of northwestern Western Australia. (a) Mineralic object recorded in Packer's laboratory notes as "branching thing" (Film LL39, negatives 3 and 4). (b) Mineralic object recorded in Packer's laboratory notes as "branching stuff" (Film LL39, negatives 20-23). Earth's Earliest Biosphere

On the basis of lessons learned from the past four decades of trial, error, and ultimate success in the search for evidence of Precambrian life, the biological origin of putative Archaean fossils can be accepted if they are: (1) composed of kerogenous organic matter (or are shown to be mineral-replaced); (2) sufficiently complex in organisational and cellular structure to rule out plausible non-biologic origins; (3) represented by numerous specimens (if one example of a taxon is preserved, others should be, too); and are (4) mem-

6.3. Evolving Life and Its Effect

539

bers of multicomponent assemblages (rather than comprising monospecific biocoenoses, a situation rarely if ever encountered in nature) that (5) exhibit a range of taphonomic degradation consistent with their mode of preservation. In accordance with younger Precambrian fossils and modern microorganisms, the objects also should be shown to (6) exhibit (genebased) morphological variability; (7) have inhabited a plausibly liveable environment; (8) exhibit organismal form interpretable as having resulted from biologic cell division; and (9) have a biogenic carbon isotopic signature. Taken as a whole, the eight microfossiliferous units now known from Archaean deposits 3200-3500 Ma in age satisfy these nine criteria for biogenicity, as well as the four other tests (provenance, age, indigenousness and syngenicity) that must be met of bona fide ancient fossils. The evidence is clear: as early as "c. 3500 Ma ago, microbial life was flourishing and presumably widespread" (Schopf et al., 2002).

6.3.

EVOLVING LIFE AND ITS EFFECT ON PRECAMBRIAN SEDIMENTATION

W. ALTERMANN As has been impressively demonstrated in the foregoing section 6.2, the detection of traces of ancient life in Archaean rocks is extremely difficult and often controversial. Reports of biologic activity in Earth's earliest sediments are equivocal because of the metamorphic and diagenetic alteration such rocks experienced. The c. 3.8 Ga metasediments of the Isua greenstone belt (see also section 2.3), Greenland are usually regarded as the oldest sedimentary rocks on Earth, but as has been discussed by Myers in section 2.2, many of these interpretations are equivocal. Schidlowski (1988) reported a whole rock carbon isotopic signature typical of biogenic 1 2 C / 1 3 C ratios from the metasediments of Isua. Mojzsis et al. (1996) reported similar results, obtained by an ion microprobe analysis on apatite mineral grains from inferred Isua metasediments which are thought to have survived the metamorphism. Rosing (1999) studied graphite globules in the interpreted pelagic Isua metasediments with an isotopic composition of 6C 13 of about - 1 9 per mil (vs. PDB). All these findings must, however, now be regarded as untenable because of likely misinterpretations of the rocks hosting these biological signatures (see discussion in sections 2.2, 5.2 and 5.3). The earliest signs of life in the sedimentary record are thus about 3.5 Ga and already include preserved microbial fossils. These earliest optically recognisable microfossils were reported from the 3.49 Ga Dresser Formation, Pilbara craton, Western Australia (cf. section 6.2). The appearance of microfossils in the geological record is thus very unexpected and is not preceded by any known biogeochemical announcement. Evolving life has gradually changed Earth's sedimentary environments, via weathering and depositional processes influenced by the chemistry of the atmosphere and oceans, and by the direct biochemical precipitation of sediments. As a consequence of atmospheric change (e.g., sections 5.2-5.5), life also may have influenced the ability of water to transport chemical complexes at different oxidation stages and with it the genesis of various The Precambrian Farth: Temposand Events Edited by P.G. Eriksson, W. Altermann, I).R. Nelson, W.U. Mueller and O. Catuneanu

540

Chapter 6: Evolution of Life and Precambrian Bio-Geology

mineral deposits. Diagenetic and hydrothermal fluids precipitate their metal ion load when entering different oxygenation and pH states, and metal ions can, vice versa, be mobilised from their chemical complexes by microbial mediation. Biota and decaying organic matter in the sedimentary system thus control the geochemistry of the resulting sediment. Stabilisation of siliciclastic sediments by microbial mats (see also sections 7.9 and 7.10) appears to have been an important process in the Proterozoic, but is difficult to prove because of organic degradation. Some rare examples of sediment stabilisation by microbial mats can also be recognised from the Neoarchaean (Altermann, 2002). Only at the terminal Neoproterozoic did the direct influence on sediments by grazing and burrowing organisms become apparent in the rock record. The Proterozoic Fossil Record

The microfossil record of the Proterozoic is incomparably greater than that of the Archaean (section 6.2). More than 2800 authentic Proterozoic microfossil occurrences contrast with approximately 30 findings for the Archaean. This discrepancy is the result of the evolution of life, but also of the much better preservation of Proterozoic rocks. On the one hand, the appearance of eukaryotic cells and, most significantly, of sexual reproduction in the Proterozoic, led to a rapid diversification of life that occupied subsequently all hydrospheric environments and there flourished, leaving its various fossilised remains in great abundance. The better preservation state of Proterozoic rocks, on the other hand, has contributed greatly to the overwhelmingly better record of Proterozoic microfossils. From the Proterozoic, over 600 stromatolitic units are known world-wide. Schopf and Klein (1992, their Table 22.3) presented a list of authentic microfossils from the Proterozoic, described from a total number of 328 units, by approximately 200 authors, in hundreds of publications. Fossilised remains of prokaryotic organisms from the Archaean and early Proterozoic are difficult to interpret, because their metabolism is poorly understood. Isotopic signatures preserved in old fossiliferous rocks are often equivocal, as their ranges overlap and do not strictly separate the different metabolic pathways, and aerobic, anaerobic or sulphur-based metabolism (section 5.5) are not easily distinguished (Mojzsis et al., 1996). Moreover, some purely geochemical reactions can also produce similar 12C/13C ratios and contamination by microbes colonising percolating groundwaters and cracks and pores in the rock presents an uncontrolled problem (see discussion by Schopf, section 6.2). Some higher, evolved Proterozoic microfossils are, however, significant in the interpretation of hydrospheric and atmospheric conditions of the Precambrian Earth (detailed discussion by Ohmoto, section 5.2). After almost 1500 My of the long exclusivity of prokaryots, eukaryotes first appeared in the geologic record, as suddenly as the emergence of prokaryotes 3.5 Ga ago. Spiralshaped, megascopic fossils (Grypania) were found in the 2.1 Ga Neguanee banded ironformation (BIF) (section 5.4), in Michigan, by Han and Runnegar (1992) and were classified as probable eukaryotic algae because of their large size, of c. 1 mm width and up to 90 mm length of the filaments. This classification, relying mainly on morphometric arguments, can however be questioned. Walter et al. (1990) classified similar, but younger

6.3. Evolving Life and Its Effect

541

Grypania fossils as probable multicellular algae. It is also possible that the Grypania from the Neguanee BIF represents a large cyanobacteria or even belongs to an extinct biological group (Knoll, 1996). Coccoid microfossils larger than 60 lam in diameter are also usually regarded as eukaryots because of their size. Such large coccoids only became abundant in the Mesoproterozoic. Eukaryotic organisms probably arose from prokaryotes, which lived in symbiosis with other eubacteria (Margulis, 1981). If the interpretation of Grypania by Han and Runnegar (1992) is correct, then the division of the phylogenetic tree separating prokaryotic eubacteria from archaea and eukaryots must have occurred long before 2.1 Ga. Because all eukaryotes are strict aerobes and require a relatively high oxygen concentration in the atmosphere to maintain respiration, and could thus not have emerged before the oxygen levels reached 1-2% PAL (Chapman and Schopf, 1983), this interpretation seems to support Towe (1990) and Ohmoto (1999) who argue that such high levels could have existed already during the Archaean (compare section 5.2). Additionally, in the 2.7 Ga prehnite-pumpellyite-bearing metamorphosed shales of the Fortescue Group from the Pilbara Craton, Western Australia, molecular evidence (biological lipids like methylhopanes or steranes) was found by Brocks et al. (1999), for oxygenic photosynthesis and for the existence of eukaryotes. These findings may, however, be explained in a different way than by their derivation from Archaean eukaryotic cells (e.g., Proterozoic or younger contamination and fossilisation of contaminants in the rock). Acritarchs first appear in 1.75 Ga rocks and became the most widespread fossils in Meso- and Neoproterozoic rocks, reaching their maximum diversity at around 600 Ma, after the Varanger ice age. These eukaryotic algae of unknown biological affinity, allow for a good biostratigraphic subdivision of the Neoproterozoic. Cyst-like structures, interpreted as reproductive bodies and evidence of meiotic cell division, first appeared in the fossil record at around 1.1 Ga, when a rapid diversification of eukaryotic phytoplankton took place. This diversification reached a maximum at about 900 Ma and was followed by a major decline at 800 to 700 Ma. This period is characterised by a major decrease in atmospheric CO2 and an increase of O2 (Holland, 1984), and by widespread glaciations (sections 5.6-5.8), conditions that were unfavourable for photosynthetic activity of algae and which were followed by a major decline in stromatolite abundance and diversity. Metazoan fossils appeared in the geologic record in the Neoproterozoic, first announced by trace fossils in sedimentary rocks of the terminal Mesoproterozoic, but unannounced by any possible links to their ancestors. Seilacher et al. (1998) reported trace fossils from a 1.1 Ga sandstone formation from India (see also section 7.10). The oldest trace fossils interpreted as burrows produced by bilateralian animals (Treptichnuspedum), occur in the 548-545 Ma shallow marine siliciclastic rocks of the Nama Group, Namibia (Jensen et al., 2000) and are age equivalent to the Ediacara fauna, that range from about 565 Ma into the Cambrian. The relationships of the Vendian Ediacara fauna to living organisms are highly disputed and their classification is uncertain, hence even their incorporation in the animal kingdom might be equivocal (Seilacher, 1989). Somewhat earlier than the Vendozoa, multicellular fossils appear preserved in phosphorites of the Neoproterozoic (570 Ma) Doushantuo Formation of southern China (Li et al., 1998; Xiao et al., 1998). If the interpre-

542

Chapter 6: Evolution of Life and Precambrian Bio-Geology

tation by Xiao et al. (1998) of these findings is correct, fossil animal embryos preserved in early cleavage stages indicate that the divergence of lineages leading to bilateralians may have taken place long before the macroscopic remains of body fossils appear in the rock record. Clearly, the evolution of sediment grazing and colonising metazoa and the development of mineralised skeletons have altered the sedimentary milieu significantly and have added a whole array of rock-forming sedimentary deposits to the environment. Colonisation by sediment burrowers and grazers must have led to widespread oxygenation of the sediment subsurface and to changes in diagenetic conditions (Fedonkin, 1996). The emergence of filterers and suspension feeders probably cleaned the ocean water and thereby increased the depth of the photic zone. Better light penetration in turn gave way to an extended colonisation of the sediment surface. The strong diversification of eukaryotic life in the Neoproterozoic also had far reaching consequences for biostratigraphy, as it has become possible to distinguish between facies-typical palaeoecological communities and, for the first time, the Neoproterozoic fossil record offered a biostratigraphic resolution that permits inter-basinal correlations (e.g., section 5.8). The Oldest Bio-Sedimentary Rocks Stromatolites are the first directly visible impact of life on sediments (cf. section 6.5). The oldest known, 3.5 Ga stromatolites from the Warrawoona Group, Pilbara craton, Australia, bear evidence of shallow lagoonal and evaporitic conditions (Barley et al., 1979; Dunlop, 1978) or shallow hydrothermal basins (cf. sections 2.7 and 6.2). They are closely associated and interbedded with the oldest little-altered and clearly recognisable sedimentary rocks, like cherts, conglomerates, sandstones and silicified carbonates. From the Neoarchaean rock record, however, microbial mats thriving in deep, aphotic conditions have been reported (section 6.5) and simple anaerobic mats can thus be expected to have covered the sediment surface in the deep subtidal realms of the Precambrian (see also section 7.9). The Warrawoona stromatolites form small structures of limited lateral extent. Large continuous carbonate platformal buildups appear first in the Wit Mfolozi Formation of the Pongola Supergroup, South Africa (c. 3.0 Ga). These laterally extensive, stromatolitic reefs, up to 30 m thick, interfinger with peritidal dolarenites, and, for the first time in Earth's history, demonstrate the ability of microbial communities to form large bioherms and to influence and reconstruct locally the architecture of sedimentary basins (Walter, 1983). This capability, however, might well have been acquired much earlier, and the occurrence of the Wit Mfolozi platform may indicate rather stable tectonic conditions within the rigid craton margin, than biologic evolution of microbial communities, as suggested by Grotzinger (1994). Grotzinger (1989) found numerous analogies of Precambrian carbonate platforms to Phanerozoic ramps and rimmed shelf environments, signifying that the absence of eukaryots and metazoans is a less important factor in the development of carbonate shelf morphology than microbially induced carbonate precipitation, eustacy (section 7.1) and tectonics. At 2.6 Ga, giant carbonate platforms emerged in intracratonic basins, where biogenic activity governed sedimentation. Equally large carbonate platforms existed on the Pilbara craton

6.3. Evolving Life and Its Effect

543

of Western Australia and on the Kaapvaal craton, South Africa (Nelson et al., 1999). On these carbonate platforms, extending for hundreds of thousands of square kilometres, sediment accumulation rates (see also section 7.11) and the amount of organic production were comparable to those on modern carbonate platforms and in microbial mats (Lanier, 1986; Altermann and Nelson, 1998). These giant intracratonic basins (cf. epeiric sea basins, section 7.7) embraced a huge variety of facies, ranging from supratidal to deep subtidal, below-storm-wave-base, reflected by different assemblages of stromatolites and by different morphologies of bioherms and biostromes (cf. with section 6.5 and Figs. 6.5-1-6.5-5). Microbial biostromes and bioherms governed the internal basin architecture, water depth and facies distribution in these depositories.

Precambrian Life in Relation to Sedimentary Calcite, Dolomite, Sodium, Sulphur, Iron and Silica As discussed in sections 5.2 and 6.4, the oxygenation state and the Ca content of Precambrian oceans are strongly disputed. Direct inorganic precipitation of calcite from Casaturated sea water on the sea floor, was proposed by Sumner and Grotzinger (1996a), who argued that Archaean carbonates differ from Proterozoic carbonate deposits by the occurrences of carbonate cements precipitated directly, as thick crusts, on the sediment surfaces in the Archaean. However, from the middle Proterozoic on, directly precipitated sea floor Ca-cements become rare, contemporaneously with the decrease in taxonomic diversity and abundance of stromatolites. Grotzinger (1994) suggested that the switch from a Na-(carbonate)-rich ocean to a NaC1 (halite) ocean occurred when the calcitic sea-floor precipitates gave way to sulphate evaporites, at about 1.8 Ga and concluded, that the middle Proterozoic seawater probably suffered from a substantial decrease in the calcium carbonate content and in the total HCO 3 to Ca~- ratio. This may have led to the subsequent stromatolite decline, as recorded in carbonate deposits at about 1.8 Ga. Grotzinger and Kasting (1993) related the lack of sulphate evaporites in rocks older than 1.8 Ga to low levels of oceanic sulphate, as a consequence of a low oxygen concentration in the atmosphere (see also section 5.5). Contrary to Grotzinger (1994), Kempe and Degens (1985) suggested that the Precambrian oceans may have been bicarbonate and soda dominated, with low chloride concentrations, up to c. 1.0 Ga. Slow accumulation of NaC1 from hydrothermal leaching of the ocean floor and gradual removal of Na-carbonates to the crust led to the demise of the "soda ocean". High Na/C1 ratio favours Na-carbonate instead of Na-chloride precipitation and allows for high concentrations of phosphate, due to low Ca 2+ content. The high pH would allow for high concentrations of organic complexes, making them available as nutrients for primitive life. The gradually increasing calcium concentration could eventually have led to the appearance of biocalcification by increasing the Ca-stress on organisms (Kempe and Degens, 1985). Kazmierczak and Degens (1986) argued that the increasing Ca 2+ pressure in the oceans could have made the aggregation of cells possible, resulting finally in the development of multicellular life.

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Microbial biomineralisation, however, is much earlier than the development of shells and skeletons discussed above. Trapping and binding of sediment particles is not evident in many early Precambrian stromatolites, which are extremely finely laminated. Such fine laminae are probably the result of direct precipitation either by inorganic, or by organically mediated or organic processes. Because of diagenetic recrystallisation, direct biomineralisation is difficult to demonstrate in Precambrian cyanobacterial mats. However, Kazmierczak and Altermann (2002) were recently able to demonstrate that Neoarchaean coccoid cyanobacteria must have calcified in a similar way to modern cyanobacteria, in alkaline or flesh water environments (cf. section 6.4). Additionally, similar authigenic phyllosilicates have been found in Neoarchaean fossilised cyanobacterial colonies, as are found in recent, degrading cyanobacterial mats. Because most of the Precambrian stromatolitic platforms are thoroughly dolomitised, dolomitisation is yet another important and not fully solved problem. In the 2.6 Ga Campbellrand Subgroup (Ghaap Group, Transvaal Supergroup) dolomites of South Africa, Wright and Altermann (2000) observed microbially mediated nucleation of dolmicrite along the outer margins of cyanobacterial sheaths, and progressive dolomitisation with increasing anoxic degradation of the sheaths. They concluded that the degree and type of organic degradation was a major controlling factor on carbonate mineralogy and dolomite formation. Active bacterial sulphate reduction (cf. section 5.5) probably removed iron as pyrite from the environment and calcite was precipitated, whereas dolomite formed below the sulphate reduction zone. Banded iron-formations (BIF) are most conspicuous Precambrian sedimentary deposits, genetically often related to the evolution of the atmosphere and biosphere (viz. section 5.4). Microbial mediation in precipitation of BIF is a widely discussed concept, where the Fe is usually thought to be precipitated non-biologically, but under the consumption of 02 from photosynthetic source reservoirs (Cloud, 1988; cf. sections 5.2 and 5.4). BIFs, however, contain much more SiO2 than Fe and the precipitation of SiO2 is an equally puzzling problem as the widely discussed precipitation of Fe (section 5.12). Konhauser and Ferris (1996) suggested that oversaturation of the ocean water with silica resulted in almost continuous precipitation of chert, while iron was deposited episodically, triggered by planktonic 02 productivity. Anoxygenic bacterial photosynthesis, however, probably played a more important role in this process than oxygenic photosynthesis (Widdel et al., 1993). Biogenic silica precipitation in BIF was proposed by several authors, among them LaBerge (1973) and Klemm (1979), although silica-precipitating organisms are unknown in the Precambrian; therefore some of the authors advocating biological silica precipitation later revised their position and proposed purely physico-chemical models for BIF precipitation (e.g., Klemm, 2000). The stabilisation of sediment by microbial mats (section 7.9) and the construction of carbonate buildups by burial of carbon and calcium in carbonate rocks are the most apparent examples of the direct influence of life on sediments, sedimentary basins and the environment in general. These processes are recorded in the earliest sedimentary rocks. The successive precipitation of large iron deposits (B IF), the disappearance of uraninite and pyrite conglomerates from the sedimentary record and the emergence of the first red beds

6.4. Microbial Origin o f Precambrian Carbonates

545

and oxygenated palaeosols in the Palaeoproterozoic (e.g., Eriksson et al., 1998b), are further overwhelming examples of the influence life had on the atmosphere and hydrosphere (cf. Ohmoto, section 5.2). The subsequent development of eukaryots and later metazoans altered not only the biotopes of microbial mats, but also introduced pelletal and skeletal sediments as new types to the sedimentary record. It is remarkable that the earliest trace fossils, metazoan fossils and eukaryotic phytoplankton with mineralised skeletons appear nearly contemporaneously with the first sedimentary phosphate deposits in the rock record (Runnegar, 1992). The increased O2 and reduced CO2 content of the atmosphere at the end of the Precambrian, the widespread glaciations, the decline in stromatolite diversity and the evolutionary changes documented in the fossil record at the end of the Precambrian, appear to be processes related to and most probably amplifying each other (cf. sections 5.2, 5.5-5.8, 6.4 and 6.5, and Altermann (2002) for a summary).

6.4.

MICROBIAL ORIGIN OF PRECAMBRIAN CARBONATES: LESSONS FROM MODERN ANALOGUES

J. KAZMIERCZAK, S. KEMPE AND W. ALTERMANN The Riddle of Precambrian Carbonates The origin of Precambrian sedimentary carbonate rocks is generally not well understood because ancient carbonates are, as a rule, diagenetically recrystallised, dolomitised, dedolomitised, often chertified and tectonised. Because of such alterations the role of microorganisms in the precipitation of Precambrian carbonate minerals is difficult to assess and is poorly documented (e.g., Fairchild, 1991; Riding, 2000). The kinetics of carbonate precipitation are often related to oxygen or iron concentration in the seawater (Sumner and Grotzinger, 1996b; Sumner, 2002). Opinions proposing that most, if not all, Archaean carbonate deposits originated non-biologically directly on the sea-floor, or were accumulated from massive whitings are popular (Grotzinger, 1994; Grotzinger and Knoll, 1995). Another point of view is that Archaean stromatolites, usually accepted as classical examples of microbially generated bio-sedimentary structures, might have originated from purely inorganic processes (Lowe, 1994b; Grotzinger and Rothman, 1996; cf. section 6.5). It has also been suggested that, only from the beginning of the Mesoproterozoic, and particularly in the Neoproterozoic, did the role of microorganisms in the formation of marine carbonate sediments became important (Grotzinger, 1994; Bartley et al., 2000). Two approaches to the dilemma of the role of microorganisms in the formation of Precambrian carbonates appear sagacious. Firstly, the detection of specimens representing close association of well defined microbiota with carbonate minerals demonstrably produced under their influence (i.e., Wright and Altermann, 2000; Kazmierczak and Altermann, 2002). The other way is to compare Precambrian carbonates, hypothetically produced by microbial communities, with modern carbonates displaying structural, textural and biotic parallelism. The only living group of common prokaryotic organisms with The Precambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. ('atuneanu

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well-documented biocalcification potential are mat-forming cyanobacteria. Consequently, knowledge of sedimentary processes controlled by these microorganisms might elucidate the origin of Precambrian carbonate deposits. Following the second approach, we discuss in this contribution two modern environments sustaining in situ calcifying benthic cyanobacterial mats and producing carbonate deposits highly analogous to Precambrian sedimentary carbonates. Structural, textural and mineralogical similarity of Precambrian and modern samples supports the assumption that, in both cases, the sedimentary rocks are products of very similar microbiota and calcification processes. The modern analogues are associated exclusively with hydrochemical conditions prevalent in the environments occupied by the cyanobacterial mats.

Hydrochemical control of in situ calcification of cyanobacterial mats Hydrochemical analyses of environments sustaining in situ calcification of cyanobacterial mats show that the common factor controlling this process is the high calcium carbonate supersaturation in the mat ambience (Kempe and Kazmierczak, 1990a, 1994; Merz-PreiB and Riding, 1999; Arp et al., 2001; Ktihl et al., in press). Theoretical geochemical considerations (Kempe and Degens, 1985) supported by field hydrochemical data (Kempe and Kazmierczak, 1990b, 1993) indicate that, to achieve the level of supersaturation necessary for carbonate mineralisation of cyanobacterial mats, the early ocean may have been highly alkaline. In order to assess the role of alkalinity in the past it is mandatory to study the processes that govern the production and consumption of alkalinity in recent water bodies. Present-Day Oceanic Processes The total alkalinity is defined as the sum of charges of the anions of weak acids, i.e., predominantly carbonic acid: TA -- [HCO 3 ] + 2[CO~-] + [OH-] + other anions of weak acids.

(1)

Present oceanic values range roughly from 2.20 to 2.26 meq kg -l. The total inorganic carbon available in seawater is defined as the molar sum of: TCO2 or Sum C O 2 -- [HCO 3 ] + [CO~-] + [CO2aq.] + [ H 2 C O 3 ]

(2)

with present values ranging from 2.01 to 2.23 mmol kg -l . The state of seawater with regard to the saturation of carbonate minerals is best characterised by their saturation index (SI): Slaragonitc or calcitc - -

log[Ca2+l * [ C O ~ - ] / K a r a g o n i t c

or calcitc,

(3)

K is the equilibrium constant of the respective mineral. SI becomes zero at saturation, positive if the water is supersaturated with regard to the respective mineral and negative if it is undersaturated. Present oceanic values range roughly between - 0 . 3 and +0.7.

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6.4. MicrobialOrigin of Precambrian Carbonates

Table 6.4-1. Current oceanic processes and carbonate parameter changes Process pCO2 in air increases pCO2 in air decreases Photosynthesis Respiration CaCO3 precipitation CaCO3 solution

Change in alkalinity No change No change Slight increase by NO3 release Slight decrease by NO3 uptake Decrease Increase

Change in TCO2 Increase Decrease Decrease

Changein pH

Changein SI

Decrease Increase Increase

Decrease Increase Increase

Increase

Decrease

Decrease

Decrease Increase

Decrease Increase

Decrease Increase

In today's ocean, the carbonate system is governed by the atmospheric CO2 level (manmade or glacial/interglacial cycles), the biological CO2 uptake and release (photosynthesis and respiration) and the calcium carbonate precipitation and dissolution. The direction of modification of the carbonate system parameters by these processes is given in Table 6.4-1. Ocean surface water is (due to low CO2 pressure) supersaturated with regard to CaCO3 minerals. Even deep oceanic water, if brought up to the surface, would degas its excess CO2, warm up and become supersaturated with regard to CaCO3. This oceanic supersaturation is, however, not governed by thermodynamics because precipitation is strongly inhibited kinetically, just as the dissolution of natural calcite is kinetically inhibited (Dreybrodt, 1999; Svensson, 1992). Present oceanic CaCO3 precipitation occurs almost exclusively by enzymatic intracellular biomineralisation. The rate calculated from sediment trap experiments amounts to 33.8 teramols C/a, in comparison to the total exported organic carbon flux of 36.3 teramol C/a (Honjo et al., 2002). It is suggested that the open marine production of CaCO3 is the principal mechanism that removes organic carbon from the euphotic zone. The ocean system seems to be tuned to this 1 : 1 Corg/Cinorg export ratio, reducing both the YCO2 and the alkalinity at a ratio of 2 mol C :2 meq. A decrease in the biomineralisation rate would apparently also cause a decrease of the Corg flux, thus slowing the total export. Replenishment of the system occurs by upwelling of deeper, older water, richer in TCO2 and of higher alkalinity. This carbon cycle maintains the ocean within narrow ranges of the carbonate system parameters (Table 6.4-1). Other forms of CaCO3 deposition, i.e. permineralisation of extracellular polymeric substances (EPS) or plain inorganic precipitation play only an insignificant role. They are, however, observed in other modern environments, where the SIaragonite and SIcalcite are > 0.8 (Kempe and Kazmierczak, 1993, 1994, 1997). Processes that Influenced the Ocean in the Past Silicate weathering and the concept of the Early Soda Ocean When estimating the ancient ocean composition, two more processes need to be considered: silicate weathering and sulphate reduction; both influence alkalinity, TCO2 and the

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Table 6.4-2. Changes of carbonate system parameters by processes important in the geologic past Process Silicate weathering Silica deposition Sulphate reduction Sulphide oxidation

Change in alkalinity Increase No change Increase Decrease

Change in TCO2 Increase No change Increase Decrease

Change in SI Increase No change Decrease Increase

saturation index (SI), according to Table 6.4-2. Silicate weathering has a long-term influence on ocean alkalinity (Urey-Reaction). It is described here as the decomposition of average continental crust plagioclase: Na0.62Ca0.38All.38Si2.6208 + H2CO3 -+ kaolinite + SiO2 + 0.6Na + + 0.4Ca 2+ + 0.6HCO 3 + 0.4CO~-.

(4)

The weathering produces two sorts of alkalinity: calcium-bound (or Mg-bound, if we consider, for example, olivine weathering) and sodium-bound (or potassium-bound, in case of K-feldspar weathering). The Ca-bound alkalinity is removed via CaCO3 deposition. The Na- and K-bound alkalinity, however, cannot be removed easily. Conceptually, this alkalinity may become removed in the long term by reverse weathering in sea floor hydrothermal reactions, which consume O H - and return the CO2 to the system. But in the short term, it will add to the alkalinity pool of the ocean. Silicate weathering is a rapid process (Kempe and Degens, 1985). The present day continental carbon sink (see also section 5.3), achieved through silicate weathering, has been calculated as 0.1 Gt C/a (109 t of carbon/year; Kempe, 1979). Thus, an amount equivalent to the present atmospheric volume of CO2 (c. 700 Gt C) is processed this way every 7000 years. Silicate weathering was an important process in the Archaean. After the loss of the primordial atmosphere by the impact of a Mars-sized planet and the formation of the Moon, very early in Earth's history (e.g., Melosh et al., 1993), a new atmosphere was formed by mantle outgassing and through cometary impacts. Thick layers of fine-grained regoliths composed of volcanic glasses and impact breccias (sections 1.3 and 1.4) must have been interacting with water and CO2, promptly weathering and delivering solutions to primordial oceans. From the total mass balance of the inorganic acids H2CO3, HC1 and H2SO4, consumed during weathering throughout Earth history (i.e., c. 65.5 x 1021 g of C or5.5 • 1021 mol, 52 • 1021 g C l o r 1.6x 1021 mol, and 5.2 x 1021 gS or0.16 • 1021 mol; Kempe and Degens, 1985), it can be calculated that there was 3.7 times more carbonic acid than hydrochloric acid and 34 times more carbonic acid than sulphuric acid available. The overall composition of the primordial crust (section 1.2) was most likely that of a komatiite (section 4.3) which had a [Ca + Mg]/[Na + K] ratio of 1.6. Hence, not enough C1 and SO4 were available to balance Na and K, that is some of the alkaline metals must have been balanced by carbonate. Early weathering solutions must therefore have had a predominance of carbonate ions rather than those of chloride. Garrels and Mackenzie (1967) have shown that solutions of various compositions, when subjected to evaporation, will become alkaline if

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there is a small surplus of carbonates over Ca + Mg or a surplus of Na + K over C1 + SO4. CaCO3 will be removed from the solution and Na and K will increase in concentration. For this reason the early ocean could have been highly alkaline, a composition dubbed "Soda Ocean" (Kempe and Degens, 1985; Kempe et al., 1989; Kempe and Kazmierczak, 1994). In fact, for the same reasons, even early oceans on Mars could have been alkaline and the hypocryotic ocean on Europa may still be alkaline (Kempe and Kazmierczak, 1997, 2002; cf. section 6.6). Soda-dominated environments today are located almost exclusively along volcanic active plate boundaries, where the chemical composition of water, the microbiology and the formation of sediments can be studied.

Sulphate reduction and the alkalinity pump Sulphate reduction (see also section 5.5) is, apart from a few deep anaerobic basins and the Black Sea, insignificant in the present ocean. It is, however, a very powerful process to alter the alkalinity of the water according to the following bulk equation, which describes the decomposition of a Redfield compound (a fictitious molecule with an average elemental composition of living matter in the ocean): C106H2630110NI6PI + 53SO 2- + 14H20

53H2S + 106HCO 3 + H P O 2 - + 16NH~- + 14OH-.

(5)

The consumption of sulphate ions demands the production of anions equivalent in charge, which is HCO 3, immensely increasing the alkalinity ("alkalinity pump" of Kempe and Kazmierczak, 1994). Figure 6.4-1 illustrates this concept and the fluxes determining the bio-geochemistry of a stagnating oceanic water body. In the Black Sea bottom waters, for example, the alkalinity is as high as 4.5 meq kg -I, that is almost double that of the open ocean (e.g., Kempe and Kazmierczak, 1990a). It increased extremely rapidly since the bottom waters there started to form only some 9000 years BP. At the same time, H2S is produced together with iron, and settles as FeSe or FeS from the water column. Simultaneously, upon slow upwelling, bacterial oxidation replaces some of the alkalinity with sulphate. Nevertheless, part of the bottom water alkalinity is always upwelled, and has increased the alkalinity of the surface waters in the Black Sea to the present 3.3 meq kg -l . During catastrophic overturn events, much of the H2S can be degassed to the atmosphere, in addition to COe and methane gases, as produced by the respiration process. This could have been the case at some of the major extinction boundaries (e.g., Kempe and Kazmierczak, 1994), a hypothesis supported by a strong negative ~ 13C shift in shallow marine carbonates deposited subsequently to extinction events (e.g., Holser and Magaritz, 1987). The sulphate reduction-produced bicarbonate ions are isotopically light, since they derive from organic matter (equation 5). Carbonates precipitated from an overturned anaerobic ocean should therefore display a shift towards more negative values. A simple mass balance calculation demonstrates the power of the alkalinity pump" in the present ocean the sulphate concentration is 28 mmol kg -I or 56 meq kg - l . The average depth (50% volume) is close to 3700 m and the depth of the mixed layer is

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Chapter 6: Evolution of Life and Precambrian Bio-Geology

Fig. 6.4-1. Schematic diagram explaining the alkalinity pump (after Kempe and Kazmierczak, 1994, modified).

around 200 m. This denotes that, per dm 2 unit area, a maximum of c. 56 meq dm -3 x (37000 d m - 2000 dm) = 1.96 x l 0 6 meq could be added to the ocean, compared to a stock of 37000 dm x 2 meq dm -3 = 74 x 103 meq or roughly 26 times the present alkalinity. Such rapid input during an oceanic overturn would definitely be catastrophical to marine life. Widespread precipitation of calcium carbonate of dramatically lower 613C ~ would be the consequence. Such events could have taken place repeatedly since the late Archaean, when the oceans acquired a significant sulphate concentration (see discussion section 5.5). Only within well oxygenated and connected ocean basins, as common in the present interglacial period, is there no threat of such a bio-geochemical disaster.

The history of the Precambrian ocean According to the scenario presented above, we suggest that the initial terrestrial ocean may have been highly alkaline. The supersaturation, in the absence of biomineralisation, would have been relatively constant at c. SIcc = + 1.0. Since the silicate weathering reaction is geologically fast, and because it is not apparent why the Earth should have started with a high pCO2 atmosphere after the Moon-forming impact, we must assume that the bulk of the outgassed CO2 reacted consecutively with water and with the fine-grained glassy surface rocks. Even in a frozen state, lacking a high pCO2 atmosphere to counterbalance the faint early sun (sections 5.1, 5.6-5.8 and 5.12), frequent impacts (sections 1.3 and 1.4) could have provided for the recurring liberation of water and gases to the primordial atmosphere. However, methane is far more effective as

6.4. Microbial Origin of Precambrian Carbonates

551

Fig. 6.4-2. Main areas of calcium carbonate precipitation in the early Precambrian highly alkaline (sodic) ocean: 1 and 2, Ca2+-stressed areas located within photic zone and potentially accessible for CaCO3-producing cyanobacterial mats; 3, deeper located aphotic or dysphotic Ca2+-stressed areas with predominantly abiogenic precipitation of CaCO 3 on the sea bottom.

a greenhouse gas and its production by mantle outgassing and by bacteria in an initially sulphate-free ocean could have provided a concentration large enough to protect the Earth from freezing (see discussion in section 5.2). The highly alkaline conditions in the primordial ocean could account for such widespread phenomena as the deposition ofjaspilites and cherts, the precipitation of the banded iron-formations (section 5.4), the deposition of primary dolomites and limestones and the widespread growth of shallow water stromatolites (Kempe and Degens, 1985). Figure 6.4-2 illustrates schematically the main areas in which carbonates could have formed in an alkaline ocean. In areas influenced by Ca-rich river and groundwater inflows (1 and 2 in Fig. 6.4-2) whitings and carbonate microbialites could prevail, whereas around deeper located hydrothermal injections (3) inorganic precipitation of carbonates may have predominated. With the establishment of deep subduction zones and growing continental surface the alkaline ocean gradually gave way to the halite-dominated ocean. The Na and the carbonates, formerly dissolved in the ocean, ended up in the continental crust and the carbon was stored within organic and calcareous sediments. Once the atmosphere became sufficiently enriched in oxygen and sulphate could be formed and kept in the ocean, the sulphate reduction and the alkalinity pump system probably played a major role in the short term regulation of the alkalinity. This could have been an effective regulation mechanism from

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the end of the Archaean throughout the Proterozoic and Phanerozoic (cf. section 5.2, where such conditions are proposed already for the Early Archaean).

Modern Analogues of Precambrian Carbonates Two modern sites producing microbial and inorganic carbonates are discussed below and compared with diagenetically little altered Precambrian carbonates. The world's largest soda lake, Lake Van, situated in eastern Anatolia, Turkey, can be seen as a modern hydrochemical analogue of the early alkaline ocean, and the sea water-filled crater lake Motitoi, on Satonda Island, Indonesia, can serve as a small scale example of a stratified basin, with a seasonally operating alkalinity pump.

Lake Van (Turkey) Lake Van fills an endorheic drainage basin on volcanically active plate boundaries (e.g., Kempe, 1977). Weathering of fresh volcanic silicate rocks has produced solutions with a surplus of bicarbonate over Ca + Mg, which after evaporation become highly alkaline (Garrels and Mackenzie, 1967" equation 4). The chemistry of Lake Van is characterised by pH: 9.7-9.8, alkalinity: 152.5 meq 1-1 , and salinity: 21.7%0 (equally shared by NaC1 and sodium carbonates, with minor contributions of sulphate, K and Mg). Although concentration of Ca is very low (4.6 mg 1-1), saturation indices (SImsee equation 3) calculated for calcite and aragonite are rather high: S l c a l c i t e - - 1.04 and S l a r a g o n i t e = 0.89 (at 20~ Mg/Ca ratio is about 30 (Kempe, 1977; Kempe et al., 1991). Lake Van produces a whole array of inorganic and microbially mediated carbonates (Kempe et al., 1991" Kempe and Kazmierczak, 1994). Carbonates form abundantly in all places where inflowing calcium-rich water mixes with the highly alkaline water. Best examples of such precipitates are whitings at tributaries of streams and rivers (Fig. 6.4-3a) and inorganically precipitated, bush-like "chemical gardens", forming at the shallow lake bottom (Fig. 6.4-3b). Most spectacular are calcareous microbialites growing along the lake shore. They occur as pinnacles or branched towers, up to 40 m high, extending to a water depth of over 100 m, where coastal aquifers deliver Ca-rich groundwater on the lake bottom (Fig. 6.4-3c), and as thin, distinctly laminated microdigitate (micro-columnar) stromatolites, growing on rocky shores at depths of up to 3 m, where Ca-rich onshore wells

Fig. 6.4-3. Calcium carbonates forming in the highly alkaline Lake Van (Turkey), a modern hydrochemical analogue of the early alkaline ocean. (a) Whitings at inlets of calcium-rich karstic and river waters; (b) Submerged, inorganically precipitated bush-like calcareous structure (water depth 5 m) at an inlet of Ca-rich groundwater; (c) A group of submerged tufa pinnacles growing on the lake bottom at a seepage inlet. Their external zone is built of calcifying mats of colonial coccoid cyanobacteria (shown in (d)) and the porous interior consists of inorganically precipitated CaCO3, often also capping the towers (water depth 17 m). The arrow in d points towards the mat surface; (e) Submerged calcareous microstromatolitic crust (shown magnified in (f)) forming on rocky bottom at an onshore spring seep (water depth 0.5 m); cracks in the volcanic rock are filled with inorganically precipitated calcium carbonate (arrow heads).

6.4. Microbial Origin of Precambrian Carbonates

Fig. 6.4-3.

553

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Chapter 6: Evolution of Life and Precambrian Bio-Geology

seep to the lake. Cracks in the rocks covered by stromatolitic crusts are filled with inorganically precipitated calcium carbonate (Figs. 6.4-3e, f). Colonial coccoid cyanobacteria, identified by Gessner (1957) as Entophysalis granulosa Kiitzing, but representing probably also pleurocapsacean and nostocacean strains, are in both cases the microorganisms mediating the precipitation of calcium carbonate in these mats (Fig. 6.4-3d). A thin veneer of inorganically precipitated anhedral or subhedral low Mg-calcite grains, from a few microns to 25 lam in diameter, is forming on the living surface of the mat (Figs. 6.4-3d, arrow, and 6.4-4a). These grains apparently precipitate due to photosynthetic uptake of CO2 by cyanobacteria and diatoms attached to the mat surface. The removal of CO2 amplifies calcium carbonate oversaturation at the mat surface, resulting in spontaneous nucleation of microcrystals (e.g., Krumbein, 1979; Pentecost and Bauld, 1988; Merz, 1992; Merz-Preil3, 2000). The main load of calcium carbonate precipitates, however, in the form of sub-micrometer-sized aragonite grains within the common mucilage sheaths (glycocalyx), embedding the coccoid cyanobacterial cells (Figs. 6.4-4c, d). The grains are relatively rare in sheaths of cell aggregates located near the mat surface (Fig. 6.4-4b), but their number increases dramatically a few hundred micrometers deeper in the mat, where they almost completely replace the mucilage surrounding the still living cells. In the dead part of the mat, the cytoplasm is decomposed and the spaces remain empty or are filled with later precipitated aragonitic microspar (Fig. 6.4-5). The high values of calcium carbonate saturation inside the sheaths are most probably related to higher pH, compared to the mat ambience, and to higher carbonate alkalinity levels generated by metabolic and katabolic processes (e.g., sulphate-reducing bacteria activity--cf, with Vissher et al., 2000; see also section 5.5), and to an increased amount of Ca 2+ and Mg 2+ in the mat interior. Ca and Mg are known to be stored in sheaths and extracellular polymeric substances (EPS) of living cyanobacteria (e.g., Somers and Brown, 1978; Amemiya and Nakayama, 1984; Decho, 1990). During autolysis, hydrolysis, and bacteriolysis of the cyanobacterial sheaths, Ca and Mg ions can be released (Stal, 2000; Paerl et al., 2001), adding significantly to the CaCO3 saturation level in the mat interior. The CaCO3 microgranules produced by the Lake Van cyanobacterial mats are similar to granular calcareous morphs obtained from bacterial cultures and bacterially decomposed cyanobacterial biomass emplaced in Ca-rich media (e.g., Morita, 1980; Buczynski and Chafetz, 1991; Rivadeneyra et al., 1998), but also to morphs precipitated from purely inorganic solutions (e.g., Tai and Chen, 1998). This hints at passive, physico-chemical calcium carbonate nucleation processes created in living or decaying cyanobacterial biomass by other microorganisms (Knorre and Krumbein, 2000; Ktihl et al., in press). Analogously to Lake Van, Precambrian non-laminated (homogeneous) and weakly laminated micrites, thrombolites (fenestrated and non-fenestrated), pelmicrites and pelsparites may represent products of calcified coccoid cyanobacterial mats. Figures 6.4-4e, f illustrate the textural appearance of non-laminated micrite generated by calcified cyanobacterial mats from Lake Van. The micrite shows typical features of fenestrated and clotted textures as described from many Precambrian carbonate sequences (e.g., Kennard and James, 1986; Altermann and Herbig, 1991 ; Turner et al., 1993, 2000). The origin of Precambrian peloidal carbonates can be explained by analogy to cyanobacterial mats from

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Fig. 6.4-4. Modes of calcification of benthic coccoid cyanobacterial mats forming microstromatolites and exteriors of tufa pinnacles in the highly alkaline Lake Van (Turkey). (a) Surface of living coccoid cyanobacterial mat with dense cover of in vivo precipitated anhedral grains of Mg-calcite. (b) Vertical section of a slimy surface (arrow) and weakly calcified common mucilage sheaths (glycocalyx) of coccoid cell aggregates just beneath it. (c) and (d) Sections of older part of living mat showing remnants of coccoid cells and submicrometer-sized amorphous granules of aragonite precipitated within the glycocalyx. (e) and (f) Clotted, in part fenestral, fine-grained (micritic) limestone produced by in vivo and early post mortem precipitation of microgranular aragonite (c and d) within mats of colonial coccoid cyanobacteria. Similar deposits are common in shallow-water Precambrian carbonates.

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Fig. 6.4-5. Diagram illustrating the distribution of calcium carbonate microgranules and microcrystallites precipitated in vivo by aggregates of colonial coccoid cyanobacteria forming benthic mats in Lake Van (Turkey). Main permineralisation agents are the aragonitic microgranules precipitating mostly within the mucilage sheaths, whereas the calcitic microcrystals covering predominantly the mat surface are temporally shed off the mat. Lake Van, usually composed of sub-globular cell aggregates, embedded in thick mucilage envelopes (Figs. 6.4-6a, b). After calcification they can produce textures almost identical to those found in Precambrian carbonates (Fig. 6.4-6c). As shown by Kazmierczak et al. (1996), weekly calcified mats composed of such aggregates can easily disintegrate into myriads of peloids, transported and accumulated elsewhere as pelsparitic limestones. The crucial role played by coccoid cyanobacteria in mass production of Precambrian carbonate sediments has been described from the Neoarchaean Nauga Formation of South Africa (Altermann and Kazmierczak, 2001; Kazmierczak and Altermann, 2002). Discoveries of benthic colonial coccoid cyanobacteria, preserved as remains of mineralised capsules and sheaths, have been documented in peritidal carbonate deposits. Abundant remains of coccoid sheaths and capsules are easily visible, after etching of flat pebbles and fine carbonate-arenitic matrix surrounding them (Fig. 6.4-7). Two types of etching patterns can be distinguished: spherical pits, 25-35 ~tm in diameter, surrounded by 3-5 ~tm thick rims, forming occasionally groups composed of two or more rimmed pits, and groups of smaller subglobular to irregularly polygonal units, 3-10 lam in diameter, forming pits separated by 2-3 ~tm thick walls (Figs. 6.4-7a, c, e). Microprobe analyses showed that the interiors of the pits are composed of almost pure calcium carbonate, whereas the rims and walls consist of calcium carbonate with high admixture of A1-Fe-phyllosilicates and dolomite. Bulk XRD mineralogical analyses of the layers enclosing the remnants of the coccoid cyanobacteria and high magnification images of rims and walls confirm the electron microprobe data (EDAX), indicating the authigenic character of the minerals forming both

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Fig. 6.4-6. (a) Mats of colonial coccoid cyanobacteria from Lake Van (Turkey) are often composed of globular subcolonies of coccoid cells, surrounded by thick mucilagenous sheaths (arrow points to the mat surface). Calcium carbonate permineralisation of such subcolonies results in formation of peloidal texture (b), characteristic for many Precambrian carbonates as in (c), a sample from the Neoarchaean Nauga Formation of the Campbellrand Subgroup at Prieska, South Africa. the carbonate filling the pit interiors (CaCO3) and their rims and walls (CaCO3 + A1-Fe silicates + dolomite). Comparison to similarly permineralised coccoid modern cyanobacteria from Lake Van and other alkaline lakes (Figs. 6.4-7b, d, f) demonstrates that carbonates are the first mineral phase filling the spaces remaining after the plasmolysis of the cyanobacterial cell contents, whereas the formation of silicates within the exopolysaccharide forming the bulk of the sheaths and capsules was a later diagenetic process.

Lake Motitoi (Satonda Island, Indonesia) Lake Motitoi is a crater lake (Fig. 6.4-8a) filling a double caldera of the small Satonda Island, north of the Tambora volcano on Sumbawa, Indonesia (Kempe and Kazmierczak, 1990b, 1993). The lake is 44 m deep on average (60-69 m in central basin) and was originally filled with fresh water and later flooded with seawater. The large input and decomposition of organic matter from the crater walls led to a rapid consumption of oxygen and to

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Fig. 6.4-7. Prccambrian peritidal carbonate deposits (left column) from the Neoarchaean Nauga Formation (Kaapvaal craton, South Africa), enclosing remnants of cyanobacterial mats mineralised by calcium carbonate and authigenic silicates, compared to modern analogues. (a) SEM image of etched section showing patterns resembling remnants of common mucilage sheaths of modern colonial coccoid cyanobacteria, like those shown in (b) from Lake Van (Turkey), photographed under Nomarski illumination. (c) SEM image of Neoarchaean globular structure (etching pattern) resembling mineralised mucilage capsules of coccoid cyanobacterium as in modern examples (d) from Lake Vai Si'i on Niuafo'ou Island (Tonga). (e) SEM image of Neoarchaean structure (etching pattern) resembling common mucilage sheaths (glycocalyx) embedding groups of cyanobacterial coccoid cells, like those shown in (f) from the modern pleurocapsalean mat growing on the rocky shore of Sulejow Dam (Poland).

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Fig. 6.4-8. Modern cyanobacterial carbonates from the quasi-marine Lake Motitoi on Satonda Island (Indonesia). (a) General view of the lake. (b) Calcareous cyanobacterial-algal reefs extending above the lake level during the dry season. (c) SEM image of surface of in vivo calcifying mats of coccoid (pleurocapsalean) cyanobacteria. (d) SEM image of a cross-section through the living mat, showing the common mucilage sheaths embedding cells of pleurocapsalean cyanobacteria. (e) Optical photomicrograph (vertical thin section) showing laminated (stromatolitic) calcareous deposit representing in vivo (dark laminae, Mg-calcite/Mg-silicate) and early post mortem (light laminae, aragonite) stages of mineralisation of the coccoid mat (cf. Fig. 6.4-10e). (f) SEM image (vertical section) showing sharp structural differences between the porous calcitic/silicate and fibrous aragonitic laminae.

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Table 6.4-3. Chemical data for Motitoi Lake on Satonda Island* Parameter

l

2

3

T (~ Salinity (%0) PH pCO2 (ppmv) Ca (mmol kg -1 ) Mg (mmol kg- 1) C-alk. (meq kg -1 ) Slarag. SIcc SIdol.

29.9 30.87 8.42 359 5.42 42.17 3.47 0.70 0.84 2.81

29.8 36.82 7.27 1204.00 6.47 50.11 6.28 -0.01 0.13 1.40

29.5 34.37 8.27 289 10.6 50.11 1.99 0.59 0.73 2.39

*Water samples taken at the end of dry season, October 1986. Column 1: sample from 10 m; column 2: sample from 30 m depth; column 3: sea water sample from Satonda Bay. For detailed information on lake chemistry see Kempe and Kazmierczak (1993). stratification of the water column. A sharp pycnocline, at a depth of about 23 m, divides the water column into an oxic epilimnion of slightly lower salinity and an anoxic hypolimnion, somewhat more saline than seawater. Table 6.4-3 gives the basic chemical parameters of Lake Motitoi and of the sea water surrounding the island. The change of the alkalinity of the lake water was probably the main reason for elimination of most marine macrobiota (Kempe and Kazmierczak, 1993). It significantly increased the carbonate supersaturation level, apparently enhancing the formation of in situ calcified cyanobacterial microbialites in the epilimnion (Fig. 6.4-8b). These well-laminated or cystous microbialites built by mats of Pleurocapsales (cf. Figs. 6.4-8c, d) encrust lava blocks and tufts of non-calcifying siphonocladalean green algae, or intergrow with arcuate thalli of calcareous red algae. The distinct lamination of these microbialites is the result of alternating micritic Mg-calcite and/or Mg-silicate and fibrous aragonite layers, well visible in optical thin section and in SEM images of etched specimens (Fig. 6.4-8f). Microstructural examination reveals that the varying mineralogy and lamination are the result of seasonally changing supersaturation in the lake (Kempe and Kazmierczak, 1993). During the wet season (November-April) the supersaturation is lowered due to the dilution of the surface water by monsoonal rain and the mat can grow non-calcified or only weakly calcified. During the dry season (May-October) supersaturation increases and in vivo precipitation of microgranular high-Mg calcite proceeds within the mucilage sheaths, as described above from Lake Van. Once the surface zone of the mat is mineralised, the dead mat below is quickly degraded under anaerobic conditions, mainly by sulphate-reducing bacteria and fungi. This results in an alkalinity increase (cf. equation (5)) which, by raising the supersaturation level, can easily trigger precipitation of early diagenetic aragonite (Fig. 6.4-9). Remnants of coccoid cells found in the calcitic, aragonitic and silicic layers are identical to the capsular coccoid aggregates forming the living mat surface (Fig. 6.4-8c). This indi-

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(a) living cells

Mg-Calcite/Mg-silicate

J

l['

'

Aragonite

-C> (b)

Mg-C al cite/M g-si licate

(c) living cells

Fig. 6.4-9. Schematic diagram illustrating the three major stages in the formation of calcareous laminated (stromatolitic) structures, composed of couplets of micrite (Mg-calcite/Mg-silicate) and microspar (aragonite) laminae, as observed in calcifying modern coccoid cyanobacterial mats from the quasi-marine Lake Motitoi on Satonda Island (Indonesia). Coccoid cell layers (c) are periodically, in vivo, permineralised with Mg-calcite/Mg-silicate (b); the non-calcified cells, remaining below the calcified layer, are obliterated post mortem by heterotrophic microorganisms that mediate an early diagenetic formation of aragonite by their metabolic activity, in the cryptic microenvironment (a).

cates that the Lake Motitoi microbialites are the product of a monospecific community of coccoid cyanobactcria. Calcification evoked by bacterial degradation of organic matter below the zone of photosynthesis is a significant factor in CaC03 formation in decaying cyanobacterial mats

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(Krumbein and Cohen, 1977; Lyons et al., 1984; Caudwell, 1987; Stal, 2000; Vissher et al., 2000; Paerl et al., 2001; Kfihl et al., in press), whereas the precipitation of high-Mg calcite at the living mat surface is easily explainable by the high Mg/Ca ratio (9: 1) in the lake water. A plausible explanation for the aragonite origin in the cryptic microenvironment is perhaps that higher Mg/Ca ratio may arise due to liberation of large amounts of Mg 2+ from decomposing cyanobacterial sheaths, significantly enriched in Mg, in comparison to ambient water (e.g., Foerster, 1964; Gebelein and Hoffman, 1973; Amemiya and Nakayama, 1984). The well-laminated cyanobacterial microbialites in Satonda Island crater lake are products of fluctuations in the environmental carbonate supersaturation. Depending on duration and regularity of such fluctuations in Precambrian marine basins, a great variety of internal textures and microfabric could have been produced, as described from Precambrian stromatolites (e.g., Hofmann, 1973; Semikhatov et al., 1979; Wright and Altermann, 2000). The great similarity of some Precambrian stromatolites to those formed in the Motitoi Lake is well exemplified by comparison of well-laminated microdigitate specimens shown in Figures 6.4-10a-e. In both cases, globular structures representing remnants of capsular sheaths surrounding groups of coccoid cells are preserved in some laminae. The Precambrian sheaths are pyritised while the Lake Motitoi sheaths are silicified, but participation of coccoid microbiota in the formation of the carbonate is clearly visible in both microbialites. Subtle periodic differences in supersaturation level were responsible tot the alternating micrite/microspar stromatolite laminae in the Neoarchaean marine environment. A different expression of the same process can be used as explanation for the whole array of laminated carbonates from Precambrian sequences, from almost purely micritic limestones to well laminated micrite/microspar couplets. Conclusions

Studies of calcification processes in modern cyanobacterial mats are fundamental for understanding of the genesis of carbonate microbialites and for evaluating the contribution of cyanobacteria to mass production of Precambrian fine-grained carbonate deposits. Modern

Opposite: Fig. 6.4-10. (a) Partly silicified (chertified) carbonate microdigitate stromatolitic structures from the late Neoarchaean Kogelbeen Formation (Kathu borehole, Altermann and Siegfried, 1997) of South Africa, probably generated by coccoid cyanobacteria with remnants preserved in some laminae (arrow in a), as globular pyritic bodies shown in (b). (c) Modern microdigitate calcareous stromatolite from Lake Motitoi (Satonda Island, Indonesia) composed of similarly alternating micritic and sparry laminae, originated by in vivo (calcitic micrite with Mg-silicate admixture) and early post mortem (aragonitic spar) calcification of coccoid cyanobacteria. The capsules are occasionally excellently preserved due to early diagenetic silicification (d). (e) Laminite built of undulating microspar and micrite layers from the Neoarchaean Kaapvaal craton sequence, South Africa, formed presumably due to in situ calcification of benthic cyanobacterial mats, as indicated by modern analogues from Lake Van and Lake Motitoi.

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sediments formed by cyanobacterial mats are intimately comparable to Precambrian homogeneous and laminated micrites, fenestral, clotted and peloidal limestones. The described modern microbial mats thrive in two hydrochemically different environments: lacustrinem Lake Van (Turkey), and quasi-marinemLake Motitoi (Satonda Island and Indonesia). Their common environmental factor enhancing and controlling calcification of cyanobacterial communities is the significantly higher calcite and aragonite supersaturation, compared to modern surface ocean water. Two modes of calcium carbonate precipitation have been recognised in these environments: (i) in vivo and (ii) early post mortem, often combined in the same mat. For the Precambrian oceans, similar processes of calcification in microbial mats and of carbonate production, in similar environments, can be envisaged following these examples.

6.5.

PRECAMBRIAN STROMATOLITES: PROBLEMS IN DEFINITION, CLASSIFICATION, MORPHOLOGY AND STRATIGRAPHY

W. ALTERMANN Precambrian stromatolites witness the early evolution of life and preserve fossil microbial remains as old as c. 3.5 Ga (sections 6.2 and 6.3). They occur in a vast range of shapes and sizes, as microstromatolites, recognisable only in thin section, and in small, patchy lithoherms, and within widespread stromatolitic lithostromes, up to tens or hundreds of metres in thickness and hundreds of kilometres in lateral extent. Although they have been mystified as "unique objects in Earth history" (Semikhatov and Raaben, 2000), they can be described and understood with an actualistic approach (Donaldson et al., 2002; see also section 7.1), and owe their abundance and morphological diversity solely to Precambrian sedimentary conditions and the early stage of the evolution of life (long exclusivity and subsequent predominance of microbial communities). The interplay of both factors, environment and biology, resulted in widespread and thriving stromatolite-forming communities in the Precambrian. Studies of Recent and Phanerozoic stromatolites, however, help to clarify most of the intriguing problems in stromatolite research. Problems in Stromatolite Definition Stromatolites are mostly defined, after Walter (1976), as lithified organo-sedimentary structures, growing through accretion of laminae by the entrapment of sediment and by direct precipitation of carbonate, under the active participation or direct influence of microbial organisms. This broad definition of stromatolites includes the microbialites of Burne and Moore (1987), thrombolites of Kennard and James (1986) and other possible definitions of at least partly lithified microbial mats, embracing the "cryptalgal limestones" of Aitken (1967), "potential stromatolites" of Krumbein (1983) and "microbolites" of Riding (1991 ). Cyanobacterial metabolic activity, including uptake of CO2, production of free oxygen and 7"hePrecambrianEarth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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the precipitation of CaCO3 (Kazmierczak and Altermann, 2002) plays a major role in such stromatolite-forming microbial ecological systems (cf. section 6.4). The above definition intentionally excludes purely chemical precipitates. Some authors argue that chemical precipitates may be indistinguishable from stromatolites formed under biological influence and moreover, may have been common in the Neoarchaean and Proterozoic. Therefore many stromatolite experts include also chemical or possible chemical precipitates in the stromatolite definition (cf. Hofmann, 1973, 2000; Sumner, 1997, 2000). Hofmann (2000) emphasises that the question of stromatolite definition is particularly important to the Archaean, because stromatolites have been used in the past to demonstrate the existence and evolution of early life on Earth and are now a primary target in the search for extraterrestrial life (cf. section 6.6, on Precambrian geology and exobiology). Of course, if purely chemical precipitates are indistinguishable from biogenic stromatolites and are included therefore in the definition, findings of stromatolites would not unequivocally prove life on the early Earth nor on planets like Mars. Many workers thus passionately defend the exclusively bio-sedimentary definition of stromatolites and elaborate on the criteria to distinguish biogenic from abiogenic stromatolitic structures. In general, morphologic complexity, resemblance to extant examples and the presence of microfossils forming associations similar to modern microbial communities in modern stromatolites are regarded as the most credible criteria to identify biogenic stromatolites (cf. discussion below).

Problems in stromatolite classification and taxonomy Classification of stromatolites has always been contentious. Considering the difficulties in defining stromatolites and in assessing the importance of chemical, clastic and biological sedimentation, and biogenesis in stromatolite formation, this is not surprising. The first description of stromatolites was made by Hall (1883), who described the Cryptozo6n (poliferum) columnar-spheroidal structures and classified them as stromatopora. Gtirich (1906) classified Devonian stromatolites as spongiostromata. Kalkowsky (1908) attempted the first classification of stromatolitic forms and microstructures. Kalkowsky (1908) introduced the term "Stromatolith" and ascribed the morphological variety of stromatolites to environmental factors such as the search for nutrients and day-light energy. Almost half a century after Hall (1883), however, Precambrian stromatolites were still not easily recognised and their biological origin was not widely accepted (Young, 1928). For the Earth's earliest stromatolites the question of biogenicity is still valid and not fully resolved (Lowe, 1994b; Buick et al., 1995b; Grotzinger and Rothmans, 1996; Hofmann et al., 1999), although most workers agree on biogenic origin of the majority of these stromatolites (Fig. 6.5-1). A broad-ranging discussion of stromatolite morphology and terminology, including the problematic Linnean nomenclature and taxonomy, is given by several authors (e.g., Walter, 1972; Preiss, 1976; Hofmann, 1976; Sastry, 1980) and many scientists from the former Soviet Union (for recent reviews, see Semikhatov and Raaben, 2000; Hofmann, 2000). Recently Semikhatov and Raaben (2000) attempted to revive the classical biological and biostratigraphic school of stromatolite research. Stromatolite morphological studies arising from the Linnean classification approach are typically highly elaborate, requiring three-

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Fig. 6.5-1. Early Archaean, conically laminated stromatolites of the Panorama Formation, Warrawoona Group, Australia, "Trendall locality". These c. 3.45 Ga old stromatolitcs display high morphologic complexity of pseudocolumnar, conically and partly wavy laminated, laterally linked structures, with erect axial zones and intervening horizontal laminae, clearly evidencing biogenic origin (Hofmann et al., 1999). dimensional studies and dense sectioning of the samples in order to reconstruct threedimensionally detailed branching and lamination patterns. Among the advocates of the binominal Linnean stromatolite taxonomy, the shape of the bioherm, bifurcation patterns of the columns and the form of the lamination are the most diagnostic morphological features. However, the Linnean classification system is, despite a long and hot debate, not descriptive and is difficult to understand and apply in the field; it is also of questionable value because stromatolite morphology bears environmental and sedimentological footprints. Geologists working in Precambrian formations evidently require clear and applicable identification criteria. The geometry of the laminae and their arrangement above each other rule the inner structure and gross morphology of stromatolitic bioherms. The size of the bioherm is defined by the continuity of the accretion processes. Walter (1972) and Walter et al. (1992) distinguished diagnostic morphological criteria of stromatolites, such as: mode of occurrence (including bioherm and biostrome forms), branching (including: bifurcation forms, e.g., parallel or divergent), shape of columns and margin structure (including: bridging between columns, bumps, ribs or wall formation), shape of lamination (including: convex, rhombic or rectangular), ornamentation (including smooth, bumpy, lobate and others), and lateral linkage between "cumulate stromatolites". Walter (1972), discussed seven criteria for classifying (columnar) stromatolites in the field. Criteria describing the bed morphology are: shape of the stromatolitic bed, structure of the top of the bed, and

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vertical and lateral changes within beds and at bioherm margins. Criteria describing the column morphology are: mode of origin of columns, the orientation of columns, the shape of columns in transverse section, and the frequency and type of branching. Similar criteria are used in the not strictly biological nomenclature. Many informal stromatolite names are in common use, in conjunction with purely descriptive terminology, such as "domical", "domal", "divergently branching columnar", "curly laminated", "stratiform" or "bulbous" stromatolites (see summary by Krylov, 1976). Such descriptive terminology is convenient to utilise in the field and easy to understand, but it lacks standardised definitions. The huge variety of stromatolite forms can not be explained by biological factors alone. Differing bioherms are often occupied by similar microbial communities, in which however, different microbial species may dominate the community (Schopf and Sovietov, 1976; Awramik and Semikhatov, 1979). Environmental impact on stromatolite microstructure and morphology has often been noticed for Precambrian and younger stromatolites, since Black's (1933) pioneering observation of the Bahamas stromatolites (e.g., Logan, 1961; Horodyski, 1977a; Playford, 1980). It is obvious (at least to sedimentologists) that in the same way that environmental factors restrict the occurrence of stromatolites, they also influence the mode and form of stromatolitic growth throughout the "life time" of a stromatolitic bioherm or biostrome. Morphological changes within one stromatolite structure, in vertical and horizontal directions, thus also reflect, at least to some extent, changes in environment (different currents, sheltering, wave action, water depth, light penetration, sediment availability, and other factors). Such factors, however, are purely physical, although they may in consequence lead to lateral or vertical alteration in the composition of microbial communities. Because the participating species of an ancient microbial community are always impossible to trace even approximately in a stromatolite, biological evolution can not be proven to control stromatolite morphology in ancient stromatolites (Monty, 1977). It is therefore unjustified to summarise all these factors under purely biological terminology and pretend that the form mirrors biology alone. Logan et al. (1964) have pointed out the deficiencies caused by use of the International Code for Botanical Nomenclature for describing the influence on morphology of sedimentary structures and the interplay of colonisation by microbial communities and environmental (sedimentary facies) conditions. An alternative, geometric stromatolite description was proposed. Logan et al. (1964) have clearly demonstrated that the Linnean taxonomy of stromatolites, based on morphology, is untenable and depends to a large extent on the degree of the exposure or on the individual sample under investigation. Under such circumstances, it is usually impossible to judge how significant the given taxonomic grouping is on genus or species level. Other attempts to create an alternative stromatolite classification were initiated. Maslov (1960) applied a polynominal classification scheme in which each morphological feature of a stromatolite was given a Latin name. This led to an even more complicated and less practical nomenclature. Komar (1989) also attempted to establish a new stromatolite classification, based on microstructural features. Four hierarchic taxa: supertype, type, genus and species were defined, but the scheme found little entry to stromatolite literature because of its poor description and definition.

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Fig. 6.5-2. Definition of processes contributing to stromatolite genesis, and classification scheme of stromatolites, simplified after Hofmann (1973, 2000).

Probably the most appealing attempt at stromatolite classification (because of its simplicity and universal applicability) is that by Hofmann (1973). As a solution to the definition problem (see above), Hofmann (1973, 2000) proposed classification of stromatolites in a pyramidal diagram of triangular base, where each corner of the pyramid represents one of the four main stromatolite-forming processes. The base of the pyramid is thus defined by the corners representing purely chemical precipitation, mechanical--clastic accretion and biological--nonskeletal accretion. This base includes all Archaean stromatolites. Younger stromatolites, displaying biological skeletal accumulation, form the peak of the pyramid (Fig. 6.5-2). However, Hofmann (2000) also noted that in many Archaean stromatolite examples it is difficult to pinpoint exactly the proportion of the various genetic

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processes involved. Furthermore, Hofmann (1973) defined morphological types of stromatolitic lamination such as flat, concave-up, convex-up, globoidal, which form nodular, columnar (branched and unbranched) and oncoidal stromatolite patterns, by stacking of laminae upon each other. To this simple morphological classification a metric scale ranging from gm to hectometres (hm) was applied in vertical (bioherm height) and lateral (bioherm width) directions. Up till today, stromatolite researchers are divided into two groups, one following the biological approach to stromatolites and the other treating stromatolites rather as sedimentary structures. Naturally these groups also differ in their opinions on the stratigraphic usefulness of stromatolites in the Precambrian.

Problems in Stromatolite Stratigraphy The lack of Precambrian body fossils and the need for stratigraphic subdivision of thick sedimentary successions lacking suitable rocks for isotopic dating prompted many attempts to use stromatolites as biostratigraphic markers. Over 800 taxa of Precambrian stromatolites are known. A decade ago, only ten of these taxa were regarded as Archaean (Awramik, 1992b). With the improvement of dating techniques and a greater interest in early life forms, many more stromatolite forms can now be assigned to the Archaean. Archaean stromatolites are generally accepted to be of less complex morphologies compared with Proterozoic examples (Awramik, 1992b; Grotzinger and Knoll, 1999; Hofmann, 2000; Semikhatov and Raaben, 2000). Hofmann (2000) also noted that Archaean stromatolites show a gradual increase in size with time, ranging over two orders of magnitude (centimetres to decimetres) in geon 34 (3.4-3.5 Ga) to six orders of magnitude (micrometre to dekametre) in geon 25 (2.5-2.6 Ga). Additionally, Hofmann (2000) noted that Archaean stromatolites, unlike Proterozoic examples, rather occur in limestone than in dolomite rocks, with sideritic, ankeritic and cherty lithologies also present. Mainly Soviet researchers in the 1960s attempted to establish a stromatolite-based subdivision of the Proterozoic, relying on the assumption that stromatolite morphology reflected stages of microbial evolution (Cloud and Semikhatov, 1969; Raaben, 1969). In the Precambrian basins of Siberia, a local and approximate stratigraphy based on stromatolite morphology could be established, and even in part extended to Australian Proterozoic basins (Walter, 1972; Grey, 1984). The stratigraphic value of stromatolites is nevertheless very low. The reasons for the illusive stratigraphic usefulness of stromatolites are probably grounded in the evolution of sedimentary facies belts and in the cyclicity of sediments (i.e., regressive and transgressive cycles; see chapter 8) rather than in biological evolution of microbial communities. Global environmental factors (see below and section 6.3) also have influenced stromatolite distribution. Two major periods of stromatolite decline and subsequent growth were identified for the Proterozoic, until a major decline of stromatolites, at the end of Proterozoic, was caused by the appearance of grazing and burrowing organisms. A Mesoproterozoic stromatolite decline was recorded at c. 1.8 Ga. Grotzinger (1990, 1994) suggested that at that time the ocean became halite-rich, from a preceding

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soda-dominated ocean, and that the first sulphate evaporites appeared. The substantial decrease in the calcium carbonate content and in the total HCO 3 to Ca~- ratio may have led to the subsequent stromatolite decline. At 800-700 Ma, a major decrease in atmospheric CO2 and an increase of 02, led to conditions unfavourable for photosynthetic activity of algae and thus caused a major decline in stromatolite abundance and diversity. As this period is also characterised by widespread to global glaciations (sections 5.6-5.8) conditions unfavourable for microbial mat formation and stromatolite growth were prevalent. For the reasons discussed above, for problems in stromatolite definition and classification, stromatolite stratigraphy can not work satisfactorily. According to Fedonkin (1996), it is possible to divide the Proterozoic into stromatolitic zones, which correlate with a precision of 100-300 million years. However, at this precision, stromatolites rarely contribute to the geochronology of the formation in question. Even Semikhatov and Raaben (2000) admit that a biostratigraphic subdivision of the Proterozoic, based on stromatolites is not possible. When we compare their figure 1 (describing "representative Precambrian stromatolite morphotypes") to Hofmann's (2000) figure 4 (describing "main attributes of Archaean stromatolites") the reason for the failure of stromatolites to serve as biostratigraphic markers becomes clear: exactly the same morphology and laminae arrangements are described in both figures. Microfossils in Precambrian stromatolites were first described from the Belt Supergroup by Walcott (1914), who compared these to cyanobacteria. Since that time, hundreds of occurrences of microfossils in Precambrian stromatolites have been reported, together with an equally large number of reports of modern contaminants, dubiofossils, misidentiffed mineral grains and possible, but not definitely recognisable fossils (for summary see Schopf and Klein, 1992; Schopf, section 6.2). Stromatolites bearing microfossils were termed "biophoric", in contrast to biogenic stromatolites (i.e., microbially influenced but not fossiliferous; Hofmann, 1973, 2000). In this definition a biophoric stromatolite is, however, not necessarily biogenic, because a chemical precipitate can mechanically or chemically trap microbes that can subsequently fossilise. Thus, even the occurrence of perfectly preserved microfossils in stromatolites rarely allows for the conclusion that the stromatolite is definitely an organic, rather than a chemically precipitated structure. Clearly only few biogenic stromatolites will simultaneously be biophoric. In Archaean stromatolites microfossil preservation is extremely exceptional. Because of the simple morphology and because of very slow evolution of Precambrian microbial life (sections 6.2 and 6.4), even well-preserved microfossils in Precambrian stromatolites do not allow for a stratigraphic subdivision of the strata. For Proterozoic stromatolites, identical microfossil taxa are reported from different stromatolites of various ages (Schopf and Klein, 1992, their Table 22.3). Stromatolites as Environmental Indicators

Stromatolites were already diverse in the Archaean and thrived in the Proterozoic. They occupied a wide range of sedimentary environments, formed at all sites of carbonate precipitation and were preserved as dolomites, limestones, diagenetic cherts, banded iron-

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Fig. 6.5-3. Examples of stratiform Neoarchaean (c. 2.6 Ga) stromatolites, intertidal to supratidal facies. In (a) the stratiform mat exhibits a surface with polygonal desiccation structures. The conical, inflexed laminites are interbedded with carbonate sands and rare aeolian (?) lithic fragments. (b) A flat-wavy laminated mat of low intertidal facies (scale bar is 10 cm). (c) Pale to white patches and bands delineate silicified portions of the laminite. The lower part of this thin section is occupied by wavy to curly laminated microbial mat with recrystallised micrite (scale bar is 2 mm). Binding activity is not clear. On the wavy laminae and between the crests, stratified, dark organic matter is enriched and transitional upwards to a birds-eye, crudely laminated mat. In (d) a subtidal stratiform microbial mat has been rolled up by storm wave action (scale bar is 5 mm).

formation (BIF; section 5.4) deposits or even in siliciclastic sedimentary rocks and between lava flows and tufts (Altermann, 2002). Although shallow pools and peritidal realms, to evaporitic and hydrothermal basins are envisaged as the main sites of stromatolite formation, caution is necessary when judging the palaeoenvironment from the occurrence of stromatolites. In the Precambrian rock record, microbial mats thriving in deep, aphotic conditions were reported (Simonson et al., 1993b; cf. Figs. 6.5-3 and 6.5-4) and microbial mats can thus be expected to have covered the sediment surface in the deep oceanic realms of the Precambrian. Indicators for mi-

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Fig. 6.5-4. Examples of Neoarchaean domal stromatolites, c. 2.5 Ga, South Africa: (a) giant domes (in cross-sectional view) up to 50 m long, and with wavy irregular and crude lamination. The elongation of the domes is parallel to palaeocurrents. (b) A rare example, where sediment binding in the lamination of a domal stromatolite can be demonstrated: the visible ripples are strongly accentuated by the bushy growth pattern of the stromatolitic mat, and the regular and equidistant arrangement of the tufts strongly implies current activity rearranging the sediment bound by the mat. In (c) domal stromatolites with lateral linkeage can be seen. In these domes a vertical change from fine columnar laminae to crude lamination is evident and can be interpreted as reflecting change from a lower to a high energy environment. (d) Small deep subtidal domes (white scale is 10 cm long) with an irregular, cortexiodal surface. The lithoherms are of sub-parallel arrangement and are embedded in shale. Their internal lamination constitutes indistinct columns and sediment-fill between the columns.

crobial colonisation of terrigenous clastic sedimentary environments were described by Schieber (1998), from the Middle Proterozoic Belt Supergroup, Montana. Here domal buildups resembling hemispheroidal stromatolites were found in shales and sandstones, next to mat-like crinkly lamination, cohesive deformed laminae, inverted hemispheroidal

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structures (such as found in polygonal stromatolites), and many other criteria for recognition of microbial mats (see summary of recognition criteria by Schieber, section 7.9). It is difficult to asses when the microbial mats entered non-marine realms, because it is often impossible to distinguish between Archaean marine and non-marine environments. Microbial communities colonised supratidal flats (Altermann and Herbig, 1991) and from there might also have inhabited moist soils on land. Stromatolites survived strong siliciclastic and volcaniclastic sedimentation within lacustrine, fluvial and deltaic settings within the late Archaean Ventersdorp Supergroup and Wolkberg Group of South Africa (Buck, 1980; Bosch et al., 1993; Fig. 6.5-5) and in the Tumbiana Formation of Western Australia (Buick, 1992). Conclusions

Two competing definitions are applicable to Precambrian (and younger) stromatolites: one including, and the other excluding, laminated chemical precipitates. The interpretation of Precambrian stromatolites involves consideration of various environments and processes. It requires a detailed facies analysis of the associated sedimentary rocks, and an analysis of possible biological remains and possible biological influence on the genesis of the structure. The proportion of biogenic versus environmental influence on stromatolite morphology and genesis can not be generally answered at present, and it probably varies from case to case. Biological classification of stromatolites appears inadequate in the light of sedimentary influence on stromatolite morphology, and does not lead to a reliable biostratigraphic subdivision of the Precambrian. However, universal alternative classification schemes are not established yet and none of the proposed classifications has been widely accepted. Precambrian stromatolites are also equivocal as environmental indicators and can be found in extremely different environments. Nevertheless, Precambrian stromatolites

Fig. 6.5-5. Examples of laterally linked, pseudocolumnar and isolated columnar stromatolites. (a) Regular, laterally linked pseudo-columns and wavy laminites. These stromatolites are associated with coarse tuffaceous sand redeposited in a wave agitated, lacustrine environment, Ventersdorp Supergroup, South Africa, 2.7 Ga. In (b) small columns, 1-2 cm high and about 2-3 cm in diameter, colonise ripple mark crests in the upper intertidal Malmani Group dolomites, South Africa, 2.5 Ga. In (c), a thin section of micro-stromatolitic columns and pseudocolumns is shown. The photomicrograph demonstrates the filigrane lamination in non-sediment-binding stromatolites and the variety of branching and column forms. Vertical and horizontal alternation between columnar and pseudocolumnar to laterally linked, wavy lamination can be observed. In (d) sediment-binding and -trapping columns are visible. Dark, saucer-shaped laminae between the columns delineate periods of slow or non-sedimentation and of formation of microbial mat (lateral linkage). Sapropel and fine stromatolitic debris are trapped between the columns and within the laminae. (Scale bars in c and d = 2.0 nun, transmitted, polarised light.) In (e), a large conical column, cut in an oblique view, and with a pronounced axial zone and silicified walls and some laminae is visible. These columns have a strongly elongated base, with long axis parallel to palaeocurrent direction, Nouga Formation, 2.6 Ga, South Africa.

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are "archives of biological, chemical and mechanical processes on the primitive Earth" (Hofmann, 2000) and are therefore important in understanding Earth's evolution. Stromatolite research requires a multidisciplinary approach and actualistic (see section 7.1) studies. At present stromatolite research suffers mainly from insufficient knowledge of biogenic processes in sedimentary environments and particularly from a lack of understanding of microbial metabolic and mineralisation processes under varying chemical and clastic conditions (cf. section 6.4).

Fig. 6.5-5.

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6.6.

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PRECAMBRIAN GEOLOGY AND EXOBIOLOGY

F. WESTALL Understanding of earliest Precambrian geology as the context for the appearance and evolution of life has profound consequences for exobiology. The reason for this is that the only example of life that we know arose and evolved within a particular geological context on Earth. However, to what extent is life, as we know it, conditioned by terrestrial environments and what are the possibilities of finding life under different conditions on other planets? Study of the Precambrian Earth can provide some clues to these questions.

The Requirements of Life There are few basic requirements for life: carbon, liquid water and energy. Despite the simplicity of the ingredients, we have not yet been able to recreate life in vitro (Brack, 2002a). What we are lacking is the precise knowledge of the conditions in which life arose. The situation is not helped by the fact that the first 500 million years of Earth's history have been destroyed by later plate tectonics. Thus, we have to rely on inferences from inherited crystals and geochemical signals in younger rocks, as well as comparative planetology (cf. chapters 1 and 2).

Carbon The carbon upon which our life is based could have originated from a number of sources. Carbonaceous material would have accreted together with the planetesimals that formed the planet (see section 1.2 for a summary): there is a wide range of carbonaceous particles in the interstellar medium ranging from amorphous carbon to polycyclic aromatic hydrocarbons (PAHs), and even more complex molecules such as fullerenes (ball-shaped molecules composed of 60, 70 and more C atoms arranged in rings); these molecules and, additionally, C2H6, N~- and CO + occur in comets, whereas carbonaceous meteorites contain a large variety of organic particles including amino acids that do not occur on Earth (Ehrenfreund and Menten, 2002). Carbon on the Earth could have come from the original volatile-containing planetesimals that collided to form the Earth (Drake and Richter, 2002). It could also have had an extraterrestrial origin. The delivery of extraterrestrial organic matter is particularly interesting since the mildly reducing conditions of the early Earth (Kasting, 1993; see also section 5.2) would not have been conducive to the production of prebiotically-important molecules, as produced in vitro by Miller (1953). After the consolidation of the Earth (and formation of the Moon, probably by collision with a Marssized planet, Canup and Asphaug, 2001; see, however, section 5.9), organic molecules would have continued to have been imported by comets, meteorites/asteroids, micrometeorites and cosmic dust (Maurette et al., 2000). This importation is on-going but the flux was much greater in the first 500-600 My of Earth history, during the heavy bombardment period. Other potential sources of prebiotically-important organic molecules include hydrothermal vent environments (Shock, 1992) and impact (sections 1.3 and 1.4) synthesis (Blank et al., 2002). The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Aitermann, D.R. Nelson, W.U. Mucllcr and O. Catuneanu

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Table 6.6-1. Distance of planets in the solar system from the Sun in km and astronomical units (AU: 1.0 AU is the distance between the Sun and the Earth, or 149,600,000 km) Planet

Distance (km)

AU

Mass (kg)

Mercury Venus Earth Mars Jupiter Saturn Uranus Neptune Pluto

57,910 108,200 149,600 227,940 778,330 1,426,940 2,870,990 4,497,070 5,913,520

0.39 0.72 1.0 1.52 5.2 9.54 19.22 30.06 39.5

3.30 x 4.87 x 5.98 x 6.42 x 1.90 x 5.69 x 8.69 x 1.02 x 1.31 •

1023 1024 1024

1023 1027 1026 1025

1026 1022

Water

Liquid water offers the most suitable phase for the diffusion and exchange of organic molecules and salts, the ingredients of life (Jakosky, 1998; Brack, 2002b). Water exists on a variety of bodies in the solar system, such as in dust grains, comets, asteroids and around some planets, but it is generally in the form of ice. In fact, distance from the Sun, as well as geological processes on a planet, are critical to presence of liquid water (Hart, 1979). In general, if the planet is too far from the Sun, for example Jupiter, any water will be in solid form (although Europa and possibly Callisto, moons of Jupiter, are examples where this is apparently not the case; Showman and Malhotra, 1999). If it is too close, for example as in the case of Mercury, water will be boiled off (cf. Table 6.6-1). During the early history of the solar system, three planets most probably had liquid water at their surfaces. These are the so-called terrestrial planets, Venus, Earth and Mars (Mercury is also a terrestrial planet but its proximity to the Sun precludes the presence of water; we are not sure whether Venus had liquid water but it is highly likely that it had). A certain amount of water, together with other volatiles would have accreted with the planetesimals forming the planets (section 1.2). It is thought that this water would have been rapidly degassed and probably removed from the early atmospheres by impact erosion (Kasting, 1993). Further water and other volatiles would have been accreted to the surface of the planets by importation through comets, meteorites and micrometeorites (Drake and Righter, 2002; Delsemme, 1998; Owen and Bar-Nun, 2000). Liquid water would have condensed on the surface of the planets as soon as they had cooled down to a suitable temperature. A planet needs to have a sufficiently heavy atmosphere to keep volatiles, such as water, on its surface and to prevent them from dissipating into space. This requires that the planet also has a certain critical mass and gravitational pull (Table 6.6-1). Water on Venus. Venus has about the same size as the Earth. Its early atmosphere was probably similar to that of the early Earth, that is mostly CO2, H20 and N2, and it may have initially had as much water as the early Earth (Hunten, 1993). However, the proximity of Venus to the Sun would probably mean that much of the water was in the form of vapour

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in the atmosphere (Hunten, 1993), even if the Sun's output was lower during the early history of the solar system than it is today (Sagan and Mullen, 1972; see also section 5.2). CO2 and water vapour are well-known greenhouse gases. The presence of these gases on Venus led to a "runaway greenhouse effect" (Kasting, 1988), which occurs when a planet absorbs more energy from the sun than it can radiate back to space. The water vapour in the atmosphere was readily susceptible to dissociation resulting in the escape of hydrogen to space. Total loss of H20 was the consequence: Venus is now a very dry planet and present-day surface temperatures are about 470~ Water on Earth. On the early Earth, on the other hand, there was, and still is, a delicate balance between the down-drawing of atmospheric CO2 in the form of H2CO3 alteration of rocks to produce carbonate weathering products, and their burial (removal from the atmosphere) in the mantle by plate tectonic activity, and re-supply due to active volcanism (see also section 5.3). In fact, the problem with the early Earth is the opposite to that of Venus: temperatures during the Hadaean needed more than just a CO2 atmosphere to keep water at the surface liquid. An admixture of other greenhouse gases such as CH4 has been proposed (Sagan and Chyba, 1997; Kasting, 1997; Pavlov et al., 2001c; detailed discussion in section 5.2). The CH4 could have had an abiotic origin from volcanic venting (Sagan and Chyba, 1997) and it has also been proposed that CH4 exhaled by methanogenic bacteria on the early Earth could have been another important source before the rise of oxygen in the atmosphere (Walker, 1977; Kasting, et al., 1983). There appears to be a close link between plate tectonic processes and climate on Earth (Evans et al., 1997; Hoffman et al., 1998b; Lindsay and Brasier, 2002; Lindsay and Brasier, section 5.3) and, consequently, the physical state of water. The amount of carbonate deposited on the continental shelves depends, among other factors, on the amount of landmass available for weathering and the rate of weathering. Burial of this carbonate removes CO2, a greenhouse gas, from the atmosphere. Consequently, atmospheric temperatures decrease leading to the global freezing of the Earth, creating thick ice caps on the surfaces of the oceans. Such an event may have been the cause of the global freeze-over that apparently occurred on Earth in the earliest Proterozoic, at about 2250 Ma, and also a number of times between 800 and 600 Ma (Evans et al., 1997; Hoffman et al., 1998b; cf. also sections 5.6-5.8). Coverage of the land by snow and ice results in lower weathering rates, as does the decrease in precipitation. The bright snow produces an albedo effect that reflects the Sun's energy away, further increasing the cold conditions. However, this deadlock breaks because continued tectonic activity and volcanic emissions inject more CO2 back into the atmosphere, the build-up of which warms up the planet so that liquid water can again exist at the surface. The situation was somewhat different in the Hadaean to early Archaean eras in the sense that there were no large extents of continental landmass to be weathered (Lowe et al., 1992). The broad, stable, continental platforms around the growing continental areas, typical of the later Archaean and Proterozoic, on which the carbonates were deposited did not exist either (Grotzinger, 1994). Thus, during the earlier part of the Earth's history, the CO2 in the atmosphere was probably buffered by chemical alteration (carbonitisation) of the large amount of fresh lava surfaces available (Westall, 2002a).

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W a t e r on Mars. There is abundant evidence for the presence of water at the surface of Mars during its early history (Noachian period, roughly 4500-3700 Ma; e.g., Carr, 1996; Baker, 2001; Clifford and Parker, 2001), although there is much debate as to whether it was liquid or frozen (Squyers and Kasting, 1994; Clifford and Parker, 2001). There is also some discussion as to whether the geomorphological features on the surface of the planet, similar to those formed on Earth by water, ice and glaciers, were actually formed by liquid CO2 rather than liquid H20 (Hoffman, 2000). However, the prevailing consensus is that water best explains all the observed geomorphological features. Based on the calculated water content of the early planet, and on the presumed early heat flux, Clifford and Parker (2001) postulate that there must have been a large body of standing water in the low lying areas of the Northern Plains. This body of water (and other water bodies, such as in impact or volcanic craters) may have been capped by ice. Towards the end of the Noachian period, the climate of Mars changed, becoming colder and drier. There is increasing evidence for the presence of ice and, finally, the disappearance of evidence of liquid water at the surface (Jakosky and Phillips, 2001). There is, however, evidence for intermittent, moderate to very minor, aqueous activity throughout the history of the planet up to the present-day (Cabrol and Grin, 1999; Malin and Edgett, 2000; Baker, 2001). The reasons for the change in climate around 3.7 Ga are complex. Jakosky and Phillips (2001) summarise the sequence of events thus:

(1) The small size of the planet meant rapid cooling that led to the shutdown of the planet's dynamo and therefore loss of its protective magnetic field. (2) Atmospheric loss due to ionic erosion from the solar wind, facilitated by the shutting down of the magnetic field. (3) Removal of the residual CO2 atmosphere as precipitated carbonates disseminated throughout the crust. (4) With little atmosphere to stabilise the presence of volatiles at the surface of the cooling planet, ice would sublimate. The H2 would have been lost to space owing to the photodissociation of water vapour in the atmosphere (perhaps as much as a third of the water budget was lost this way). (5) Finally, the remaining water would have been trapped in the increasingly thickening subsurface cryosphere. The presence of water on Mars necessitates a substantial atmosphere that was probably CO2-rich. The present atmosphere of Mars is only 6 mbars, of which 96% is CO2. What happened to all the CO2 in the earlier, heavier atmosphere? If much of the atmosphere was removed as carbonate precipitates, as is postulated, where are they? Large deposits of carbonate have not, to date, been discovered by the orbiters. The early Precambrian Earth could offer a clue. As noted above, large carbonate deposits are not found in the early Archaean supracrustal rocks. On the other hand, there is plenty of evidence for alteration of the lavas by carbonatisation. Possible explanations for the lack of large deposits of carbonate on Mars include its dissemination as alteration products mixed throughout the volcanic crust, or that there are really relatively little carbonates anywhere on Mars owing to the total dominance of the other loss processes.

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Water on Jupiter's Moons Europa and Callisto. The Gallilean satellite, Europa is an interesting exception to the rule, that planets distant from the Sun have no liquid water, because it appears that liquid water exists beneath its icy crust (Squyers et al., 1983; and possibly beneath the surface of Callisto: Showman and Malhotra, 1999). Schenk (2002) calculates a 12 km thickness for the ice crust although other estimates range from 5 to 30 km. The hypothesis that liquid water exists beneath the crust is based on a number of data including measurements of density, IR spectra, albedo, magnetism, the moment of inertia (derived from gravity measurements), as well as the geomorphological features at the surface (Lane et al., 1981; Squyers et al., 1983; Ojakangas and Stevenson, 1989; Kivelson et al., 2000; Pappalardo et al., 1998; Melosh et al., 2002). It is thought that heat provided by the decay of radiogenic elements in Europa's rocky crust and energy produced by tidal motions provoked by Jupiter would be sufficient to produce a sub-ice ocean. It is the presence of water on Europa that makes it of exobiological interest. As with the terrestrial planets, Europa would have had the same endogenous and exogenous sources of organics. Heat flux during the early history of the planet must have been greater than it is today due to the decay of radioactive nucleides, probably resulting in submarine volcanic and/or hydrothermal activity. Redox reactions between various mineral species associated with hydrothermal activity and sea water (possibly sulphate-rich) would have provided energy for prebiotic mineral formation and for metabolism, once life had appeared (if it ever did?) (Zolotov and Shock, 2002). Modelling has shown that nutrients could reach the ocean from the surface of the ice crust (Barr et al., 2002). However, calculations suggest that there would have been very little energy available at the rock-water interface on Europa for metabolic reactions, consequently limiting the potential biomass (Jakosky and Shock, 1998). Europa is very different from any of the terrestrial planets and perhaps the only terrestrial analogue in terms of its physical characteristics is Lake Vostok on Antarctica. The lake has apparently been isolated from the atmosphere for over a million years by an icy crust that is at present about 4 km thick. Beneath the ice there is a layer of liquid water, kept in this state by geothermal heat gradients, as well as the pressure exerted by the weight of the ice. Preliminary reports of bacteria living in the lake come from cells contained in the ice layers just above the lake (Karl et al., 1999; Priscu et al., 1999). These organisms must obtain their energy from sources other than sunlight, such as from chemical reactions involving the crushed up rock pieces brought into the lake by the grinding action of the glaciers over the underlying rocky surface. Confirmation that life can survive in Lake Vostok might strengthen the argument for the possible presence of life on Europa. Lake Vostok was formed after life had bloomed on Earth, whereas Europa would offer an example of the origin of life ex novo. If life did arise on that planet, analysis of the geological and geochemical conditions would provide us with some idea of how life could have started in the oceans on Earth. Energy

There were a number of potential sources of energy for the reactions leading to the appearance of life and for fuelling the metabolism of early life. They include sunlight and

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the energy from chemical redox reactions. The Sun was apparently weaker than it is today (section 5.2), but the heat flux on the terrestrial planets was much higher, leading to increased volcanic and hydrothermal activity (e.g., section 3.6). To this can be added the energy imparted during impact events (sections 1.3 and 1.4), and the subsequent volcanic/hydrothermal activity arising in such situations. Modelling of the energy available for organic molecule synthesis and metabolic processes in hydrothermal systems has shown that such loci would have been highly suitable either for the origin of life and/or as a habitat for early (and extant) life (Baross and Hoffman, 1985; Shock, 1992, 1997; Fisk and Giovannoni, 1999). This hypothesis is supported by phylogenetic clues that the earliest common ancestor to life may have been a thermophilic organism (Forterre et al., 1995), but probably not a hyperthermophilic one as has been suggested (Stetter, 1998).

The Geological Context for Early Life on Earth We have seen above that a number of planets in the solar system satisfy or have satisfied in the past the basic requirements for hosting potential life. At this stage it is necessary to examine the early terrestrial rock record to determine the manner and extent to which early geological processes (may) have influenced life, and the implications of this for life on other planets. Severe limitations are imposed by the fact that the early terrestrial rock record has been more or less effaced. Plate tectonics is generally given as the reason for the lack of crust older than about 4000 Ma (see also discussion in section 3.6), but destruction through massive meteorite bombardment (termed "gardening" in the planetary community) will also have been a significant factor. It is generally believed that there was a gradual decrease in the flux of impactors affecting the planets during the first 500-650 My of the history of the solar system, ending between 3850 and 3900 Ma (cf. section 1.2). Based on revised dating events on the Moon, Ryder (2002) suggests that there was a separate peak in these impacts between 3900 and 3820 Ma. Estimations of the severity of the impact events ranges from between 0 and 6 planet-sterilising impacts (and life-destroying, if life had already developed anywhere) in the Hadaean and earliest Archaean period up until about 3850 Ma (Maher and Stevenson, 1988; Sleep et al., 1989), and no impacts large enough to completely volatilise the surface of the Earth (Ryder, 2002). Mars, on the other hand, may possibly be able to provide critical data for the time period lacking on Earth, since the surface of Mars is generally old, even that beneath the Northern Plains. However, this planet also underwent the destructive bombardment that has cratered the Moon and, therefore, intact crust older than 3850 Ma may not be common. One testimony of this ancient crust is, however, the ALH84001 meteorite, which has a crystallisation age of 4.5 Ga.

Potential habitats for early life We can consider the environmental conditions of the early Earth to be extreme with respect to those of the present Earth (hot, no or only trace amounts of 02, possibly saltier oceans, higher UV flux; sections 5.2-5.5), and it is in these conditions that life arose. The early

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Fig. 6.6-1. Potential habitats for life on early Earth. Three main environments were present: deep ocean, including hydrothermal and volcanic submarine rocks; open ocean planktonic environment; and shallow water basins and lagoons, including volcanic and hydrothermal/hot spring environments.

Earth provided a relatively wide range of potential habitats for life, ranging from deep water to subaerial (Nisbet, 1995; Nisbet and Sleep, 2001; Westall, 2002a), irrespective of the tectonic development of the planet. The potential habitats (see also chapters 2 and 4) may be summarised as follows (Fig. 6.6-1): 1. The deep ocean environment would have included hydrothermal, and volcanic rock surfaces (including fractures), the sediment-covered ocean floor, and possibly even the subsurface sediment. The sediments would have been derived in situ from the alteration products of the oceanic basalts (carbonates and clays) and chemical precipitations due to hydrothermal exhalations, as well as allochthonous, sedimented volcanic ash-fall and maybe some turbiditic input from tectonically unstable sedimented slopes. (It is unlikely that the latter would have played an important role because of the lack of large, exposed continental areas during the HadaearffEarly Archaean periods. Detritus from subaerial erosion contributes vastly to the sedimentary input in modern oceans.) 2. The open ocean provided a planktonic environment. 3. Around the flanks of exposed hot spot volcanoes or emergent portions of basaltic crust in the early oceans, and around the small exposures of fractionated (continental) crust there would have been: (a) shallow water basins and/or lagoons, which hosted lava surfaces, sediment surfaces, hot springs, as well as planktonic environments; (b) an intertidal environment with the same types of substrate (i.e., sediment/rock surfaces, hot

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springs); and (c) a subaerial environment with hot springs, rock surfaces, regolith surfaces and maybe pools of standing water would also have been available for organisms capable of surviving in those habitats. Given the limited rock record from the Early Archaean, it is difficult to estimate the rate of colonisation of the various environmental niches. With the exception of one observation of turbidites in the Isua (see also section 2.3) greenstone belt (Rosing, 1999), no deep water environments are preserved in rocks older than 3300 Ma. (These turbidites are, however, also interpreted as highly deformed mylonites, cf. with Myers, section 2.2.) Already the oldest microfossil remains from the Barberton and Pilbara greenstone belts document that life was widespread in shallow water areas and had invaded the intertidal zone, possibly even becoming subaerial (Walsh, 1992; Hofmann et al., 1999; Westall et al., 2001a, b; Westall, 2002a, section 6.2). There is also a strong association between early microbial mats and hydrothermal activity (Nijman et al., 1999; Rasmussen, 2000; Westall et al., 2001 a, c; Westall, 2002a; see also section 2.7). The lack of environmental indications of free oxygen on the early Earth indicates that early terrestrial life must have been anaerobic (but compare with section 5.2, where arguments for an oxygen-containing Archaean atmosphere are discussed). The life represented by these oldest preserved fossils is relatively well developed (sections 6.2 and 6.3), showing identical morphologies and behaviour to modern prokaryotes (bacteria and archaea, to distinguish the simple cells from eukarya). We have no record of the origin of life, nor of the first cellular life. If there is crust old enough on Mars, it may be able to provide this "missing link" (providing that "Martian life" and terrestrial life are sufficiently similar). However, given the difficulties in interpreting the relatively well-developed microbial fossils from the Barberton greenstone belt and the Pilbara craton (Schopf and Walter, 1983; Westall, 1999; Altermann, 2001; cf. also section 6.2), would we be able to interpret fossilised early cellular life (Westall et al., 2001 b)? An important characteristic of prokaryotes in general is that they are surface specific. Habitat, from a microbial point of view, means a very small area ranging from tens to hundreds of square micrometers (although microbial mats can extend for kilometres in suitable aqueous environments, such as the bottoms of lagoons). Within this range, these microorganisms can control their own micro-environment. Together with their incredible capacity for adaptation and their survival strategies, this means that prokaryotes can exist in a wide range of environments ranging from cold and hot deserts, nutrient-limited ocean waters, high pressures, high radiation environments, to highly alkaline, acidic or saline environments (list not inclusive; Nealson, 1997; Rothschild and Mancinelli, 2001), irrespective of the large-scale geological context.

Further Evolution of Life and the Influence of Geological Processes Whereas geological processes during much of the early history of the Earth seemed to have no particular influence on the existence of simple prokaryotic life, this does not appear to have been the case with respect to further steps in evolution, which seem to have coincided with environmental changes brought about by large scale geological phenomena (Ward and

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Fig. 6.6-2. Late Archaean to Proterozoic habitats for life on Earth.

Brownlee, 2000; Lindsay and Brasier, 2002; section 5.3). Early life on Earth was anaerobic and, judging by the extensive microbial mats that developed in shallow water to subaerial environments since the Early Archaean (Fig. 6.6-2), could probably have already developed photosynthesis in that era (Nisbet and Sleep, 2001; Westall et al., 200 lb). Photosynthesis is a metabolic pathway, which converts light energy into chemical energy, resulting in the incorporation of CO2 into the cell. The development of photosynthesis increased the efficiency of cellular metabolism compared to chemotrophy (use of chemical compounds as an energy source), which would have been the earliest form of metabolism. This development called for modifications in the composition of the cell membrane to incorporate lightsensitive pigments. Nisbet (1995) hypothesises that carotinoid pigments in early microbes, which would have protected them from UV radiation, may have been adapted to harvest light or, alternatively, that thermosensors in bacteria around hydrothermal vents could have been thus adapted. The simplest form of photosynthesis is called anoxygenic photosynthesis and involves one single photochemical reaction. An even more efficient metabolic pathway is that of oxygenic photosynthesis which involves two photoreactions. In the latter process, the H20 molecule is split and 02 is released as a by-product. A typical bacterium using oxygenic photosynthesis is a cyanobacterium (but not all cyanobacteria are oxygenic photosynthesisers). On the modern Earth, by far the most important component of microbial mats in aerated, shallow water conditions is formed by cyanobacteria. These organisms are implicated in the construction of the stromatolites (section 6.5) so characteristic of the Late Archaean-Early Proterozoic era carbonates (Walter, 1994). These typically large microbial constructions are not found in the Early Archaean formations, although small domal

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and tabular stromatolites have been described from 3300-3470 Ma horizons in the Pilbara and from Barberton (Walter et al., 1980; Groves et al., 1981; Buick et al., 1981; Walter, 1983; Byerly et al., 1986; Hofmann et al., 1999; Westall et al., 2001a, c). Despite the fact that it is difficult to determine species from simple morphologies, none of the microfossils from these Early Archaean horizons resembles cyanobacteria, which generally have very characteristic morphologies (Schopf, 1993; Madigan et al., 2000; cf. section 6.2). The oldest indubitable cyanobacterial microfossils occur in c. 2600 Ma carbonate rocks from the Campbellrand Subgroup in South Africa (Klein et al., 1987; Altermann and Schopf, 1995; Altermann and Kazmierczak, 2001). They more or less coincide in age with the first chemical biomarker signature for cyanobacteria in 2700 Ma rocks from the Hamersley Group (section 5.4) in Australia (Summons et al., 1999; Brocks et al., 1999). These occurrences also coincide with the period in which there were large continental landmasses, broad carbonate platforms deposited on the continental shelves, and shallow inland seas (cf. epeiric seas, section 7.7) (Lowe et al., 1992; Grotzinger, 1994). The shallow continental shelves and shallow seas therefore provided an ideal habitat for the microbial mat communities that formed stromatolites. Thus, since probably at least 2700 Ma, oxygenic photosynthesising bacteria were pumping oxygen into the atmosphere. The oxygen was initially taken up in the oxidation of reduced species on the surface of the Earth (mostly iron and manganese), producing the banded iron-formations characteristic of the Late Archaean/Palaeoproterozoic era (Rye et al., 1995; cf. also sections 5.2 and 5.4, for divergent viewpoints). Another, coeval process also influenced the rise of oxygen in the atmosphere: the burial of carbonate and organic carbon in the mantle by normal tectonic processes (section 5.3). This removed CO2 from the atmosphere and, as we saw above, led to drastic consequences for the climate, producing a snowball Earth (for critical discussion of this hypothesis, see sections 5.6-5.8). The appearance of more sophisticated organisms, the eukaryotes, which are obligate aerobes, required the presence of significant amounts of atmospheric oxygen. On the basis of evidence for the first appearance of oxidised palaeosols, Rye and Holland (1995) time the occurrence of important quantities of oxygen in the atmosphere at about 2100 Ma. The oldest known (to date) eukaryotic fossils, Grypania, date back to 2.1 Ga (Han and Runnegar, 1992). However, biomarkers with eukaryotic characteristics have been obtained from the 2.7 Ga Hamersley Group sediments, thus putting the biochemical signal for the appearance of eukaryotes far earlier than the first fossil evidence (and also far earlier than the generally accepted significant rise of 02 in the atmosphere; cf. Ohmoto, section 5.2). This may be a consequence of two factors: (1) it is probably very difficult to distinguish between the fossils of a primitive, single-celled eukaryote and a prokaryote (the Grypania fossils already exhibit a considerable degree of sophistication, but could also be a large cyanobacteria), and (2) there is the possibility that the biochemical signal from the Hamersley Basin sediments has been contaminated by younger eukaryotes. Westall and Folk (2003) have noted that contamination of ancient sediments by younger microorganisms is a particular problem in carbon-rich deposits, and those of the Hamersley Basin are particularly rich in carbon (moreover, Brocks et al., 1999, acknowledge this possibility themselves). Despite these questions on the precise timing of events, the fact is that

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evolution on Earth appears to have been strongly driven by 02 in the atmosphere. Tectonic processes are implicated in the creation of a suitable habitat for the development of microbial mats of oxygenic photosynthesisers, and both biogenic and tectonic phenomena contributed to the rise in oxygen in the atmosphere.

Implications for Life in the Solar System The above considerations mean that, on any planet with liquid water, a source of organics and energy and some essential elements (C, H, O, P, S, N especially and also Na, K, Mg, Ca and C1 plus other trace elements), life could potentially have appeared. As we have seen, these conditions in all likelihood existed on early Venus, Mars and Europa; there is no reason to suppose that life did not arise on these planets (Jakosky, 1998). Venus may have briefly had shallow oceans (Jakosky, 1999) and exposed volcanic land masses, just as the Earth did, thus providing a similar range of potential environmental habitats as occurred on the Hadaean to Early Archaean Earth. Noachian Mars, on the other hand, was not an ocean-covered planet, even if a third of the planet could have been covered by a kilometre deep (ice-covered?) ocean for a short period of time in the very early Noachian (Clifford and Parker, 2001). Since large-scale geological processes are irrelevant to the existence of simple prokaryotic life, it is likely that early life on these planets was represented by simple prokaryote-like cells. Given the coincidence between biological and geological evolution on Earth, the question as to whether large-scale geological regulatory processes are really essential to evolution is pertinent. (Note that Altermann (2002), e.g., argues that prokaryotic life had an important influence on the evolution of sedimentary environments (section 6.3) and that, only when life became more advanced, did geological processes gain some control over life's evolution.) However, on Venus, the runaway greenhouse effect would have extinguished any previous life, if it had ever appeared. The high surface temperatures on this planet (over 450~ would also make it unlikely that any evidence of fossil life could survive. Furthermore, even if such fossil life had existed, it would be buried beneath about a kilometre of volcanic resurfacing. In any case, the search for traces of life on Venus would be impossible with the present technology due to the extreme thermal environment. Mars is a more tantalising proposition. Mars is tectonically "quiet" from a plate motion point of view (Sleep, 1994; Zuber, 2001). Although it may have had an ocean very early in its history, continent-forming plate tectonic activity (that requires liquid water), such as was important during the first two billion years of Earth's history (see, however, discussion in section 3.6), did not apparently occur. On a relatively temperate Earth, plate tectonic activity led to the formation of suitable, large expanses of carbonate-rich, shallow water environments in which oxygenic photosynthesising microbial mats could develop. Early Mars offered many shallow water environments in the form of water-filled impact and volcanic craters, as well as around the edges of the early ocean (Fig. 6.6-3; e.g., Cabrol and Grin, 1999). Were these conditions sufficient for the evolution of oxygenic photosynthesis? To date, no large areas of carbonate deposits have been detected. Perhaps this indicates that the necessary environmental conditions did not exist on Mars. Modelling

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Fig. 6.6-3. Potential habitats for life on early Mars (Noachian).

suggests that there was not enough time nor available energy for photosynthesis to have occurred on Mars before complete deterioration of the climate (Jakosky and Shock, 1998). On the other hand, if oxygenic photosynthesisers did develop independently on that planet, close examination of the geological context of such hypothesised microorganisms would provide valuable, comparative information in helping understanding of biological evolution on Earth. Such a possibility is envisaged by Hartman and McKay (1995) to explain the presence of an oxidant at the surface of Mars. They hypothesise that oxygen produced by photosynthesisers that existed early in the history of the planet caused the oxidation of the Martian surface, rather than H202 (or other oxidants) produced photochemically in the atmosphere, which is the favoured solution (superoxides; e.g., Yen et al., 2000; Levin, 2001). Continuing the above line of reasoning, it is unlikely that more advanced organisms, such as terrestrial eukaryotes, could have developed on Mars, given the necessity of an oxygenic atmosphere. Here again, any evidence to the contrary of more sophisticated organisms would necessitate rethinking of current hypotheses regarding the evolution of life on Earth. The ability of life on Earth to survive in inhospitable environments, such as deep under the surface (Stevens and McKinley, 1995; Parkes et al., 2000; Freund et al., 2002), or in frozen habitats (Gilichinsky, 2002) demonstrates that life could potentially still exist on Mars. A possible scenario is as follows: life developed on the surface of early Noachian Mars and, as the climate deteriorated, was driven into endolithic refuges (Friedmann and

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Koriem, 1989) or the frozen subsurface (the cryosphere). It would have eventually died out in the surficial endolithic environments, but it might have been able to survive as frozen spores in the cryosphere, or there may have been (be) a substantial subsurface biosphere. There is abundant evidence of periodic releases of melted ice onto the surface of Mars throughout its history (Baker et al., 1992). Cyclic volcanic activity and heating through impact events could explain this episodic melting of the cryosphere. Microbial spores could be entrained in the melt water and, if still viable, might possibly be able to metabolise slowly for the short period of time that the water could exist at the surface of the planet (volcanic exhalations of CO2 would also provide the necessary temporary atmosphere to keep the volatiles at the surface and, even ponded in crater lakes (Cabrol and Grin, 1999) for a short period of time). The resuscitated cells could have the potential of being fossilised by mineral precipitates from the water body, and/or by hydrothermally precipitated minerals if hydrothermal activity was associated with the outflow of melted groundwater. With its ice-covered liquid water ocean, Europa offers another possibility of extraterrestrial life (Jakosky, 1998), albeit, probably limited by energy constraints. Given the lack of geological processing of the planet in terms of production of suitable, light-bathed, oxygenated, shallow water environments that are essential for further biological evolution on Earth, it is unlikely that life could have evolved past the simple prokaryotic stage on Europa. However, if the opposite is proved to be the case, the lessons learned from Europa will greatly aid our general understanding of evolution and the possibilities of finding life, from primitive to more evolved, on other bodies, not just in our planetary system. Further general reading that treats the topic of extraterrestrial life is to be found in Jakosky (1998), Brack (1998), Lunine (1999), Clark (2000) and Horneck and Baumstark-Khan (2001).

6.7.

COMMENTARY

W. ALTERMANN The fossil record of the Precambrian, but particularly that of the Archaean is difficult to ascertain and to understand. Although the Archaean encompasses almost 45% of Earth history (the Proterozoic encompassing another c. 45%) Archaean unmetamorphosed sedimentary rocks, that could potentially bear ancient traces of life, are rare. Because life in the Archaean entailed solely prokaryotic organisms (bacteria and cyanobacteria) with simple morphologies and of simple organisation, and because the preservation of such structures is usually very incomplete, identification of Archaean microfossils is often ambiguous. The time and the environmental setting for the appearance of early life on Earth are, however, important for understanding of past and present ecosystems, and also for exploring the possibility of extraterrestrial life. Numerous criteria have been set up to ensure the authenticity of Archaean microfossils (Schopf, section 6.2). Among them, the most important are: the samples must be of firmly established Archaean provenance and age, the fossils must be indigenous to the rock and syngenetic with its deposition, and they must allow morphological recognition of llw Precambrian harth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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assured biological origin. Different lines of positive evidence like fossil morphology, the presence of carbonaceous matter, isotopic signatures and mineralogy need to be involved in such testimony. Petrographic thin sections that can provide clear evidence of the relationship between fossil-like objects and their encompassing rock matrix, are the best suited study objects, and can be supported by electron microscopy (SEM), Raman spectroscopy, Atomic Force Microscopy (AFM) and isotopic investigations, performed preferably on individual fossils (section 6.2). Morphologically representative measurements and statistical comparison to extant species are crucial. Simple coccoidal forms, even if arranged in clusters resembling colonies, and of sizes corresponding to extant species, must be treated with scepticism. This is especially valid if such structures cannot be observed in the samples by other means than SEM or AFM. True coccoidal microbial fossils can be expected to be spheroidal, and to have an easily discernible cell wall; they should have a finite range of cell diameters between < 0.5 lam and c. 60 lam; they may exhibit "bean-shaped" cells, reflecting their derivation from the division of a common parent cell; and they should exhibit traces of the typical mucilaginous sheaths surrounding such benthic colonies. The morphology of filamentous fossils offers more unequivocal characteristics, including cylindrical shapes with robust cell walls, possible septation, a finite range of widths (< 0.5 lam to c. 100 lam) and a finite range of lengths, of a few tens to a few hundred lam. They should exhibit varying degrees of sinuosity that can be correlated with width and length of filaments, have distinct terminal cells, and display other typical criteria (section 6.2). Only a few of the reported fossil findings from the Archaean fulfil all of these requirements for unequivocal fossil identification, among them the fossil remains from the 3.43 to 3.49 Ga formations of Western Australia. Eight microfossiliferous occurrences now known from 3.2 to 3.5 Ga Archaean deposits can be accepted as authentic ancient fossils and provide firm evidence that 3500 Ma ago, microbial life was flourishing and presumably widespread (Schopf, section 6.2). However, life in a more primitive form must have existed in the oceans before that time, and the key to proving it lies in finding older rocks or new methods, allowing for identification of cryptic organic remnants. Recognition of Proterozoic fossils presents only minor problems because remains are abundant, and often much better preserved than in the Archaean (section 6.2). They can even be so similar in morphology to modern microbial taxa that both their biogenicity and biological affinities can be established readily. The better preservation of Proterozoic rocks plays an important role therein. On the other hand, the appearance of eukaryotic cells in the Proterozoic, and most significantly, of sexual reproduction, led to a rapid diversification of life that occupied all hydrospheric environments, and left its fossilised remains in great abundance. Eukaryotes appeared in the geologic record in the 2.1 Ga Neguanee BIF (Han and Runnegar, 1992). Acritarchs appeared first in 1.75 Ga rocks and became the most widespread fossils in Meso- and Neoproterozoic rocks, allowing for a good biostratigraphic subdivision of the upper Proterozoic. Cyst-like structures, interpreted as reproductive bodies and evidence of meiotic cell division, appeared first in the fossil record at around 1.1 Ga, when a rapid diversification of eukaryotic phytoplankton took place. Metazoan fossils are known in the geologic record from the Neoproterozoic, first perceptible from trace fossils in a 1.1 Ga sandstone formation from India. The oldest trace fossils interpreted as

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589

burrows produced by bilateralian animals occur in the 548-545 Ma shallow marine siliciclastic rocks of the Nama Group, Namibia, and are age-equivalent to the Ediacara fauna. The degree to which the fate of life and of the atmosphere and biosphere are interwoven is underscored by the fact that the evolution of life had an impact on the changing composition of the atmosphere and oceans, and the reverse holds equally true (Altermann, section 6.3). The increase in photosynthesis resulted in accumulation of free oxygen which is toxic to most prokaryots, and which thus produced an advantage for eukaryotic cells. The late Precambrian biological crisis might have been self-produced, being triggered by photosynthesis and by contemporary fixing of carbon via burial of organic matter in sediments. The rise in 02 and fall in CO2 partial pressure could have lead to an "inverse greenhouse" and eventually to global glaciation and unfavourable conditions for microbial growth (Hoffman et al., 1998; cf. sections 5.6-5.8). Under such conditions, microbial mats would have been forced into environmental niches such as hot springs, subterraneous aquifers or subglacial water bodies. This would again have hindered photosynthesis and eventually gave way to higher CO2 partial pressure. Simultaneously, microbes surviving in the ecological niches conquered the ice-free realms and more advanced forms of life emerged at the terminal Proterozoic, to finally occupy virtually all sedimentary environments. The first regionally extensive stromatolitic reefs are recorded in the Wit Mfolozi Formation of the Pongola Supergroup, South Africa (c. 3.0 Ga) and demonstrate the ability of microbial communities to form large bioherms and to influence and reconstruct the architecture of sedimentary basins. Before 3.0 Ga only small, patchy stromatolitic reefs in lagoonal and hydrothermal settings are known. In the Neoarchaean and Proterozoic, however, stromatolitic reefs were shaping the morphology of large basins, hundreds of thousands of square kilometres in extent, on the Pilbara and Kaapvaal cratons. Sediment accumulation rates and the amount of organic production in these basins were comparable to those on modem carbonate platforms and in microbial mats (cf. section 7.11). Microbial biostromes and bioherms governed the internal basin architecture, water depth and facies distribution in these depositories (section 6.3). The construction of carbonate buildups by burial of carbon and calcium in carbonate rocks and the stabilisation of sediment by microbial mats are the most apparent examples of the direct influence of life on sediments, sedimentary basins and the environment in general. Several models for Archaean and Proterozoic carbonate precipitation challenge each other (Kazmierczak et al., section 6.4). The early oceans on Earth might have been highly alkaline. Alkaline conditions could account for such widespread phenomena as the deposition of cherts and banded iron formations, of primary dolomites and limestones and the widespread growth of shallow water stromatolites. In areas influenced by Ca-rich river and groundwater inflows, whitings and carbonate microbialites could have prevailed, and around deeper water hydrothermal vents, inorganic precipitation of carbonates may have dominated. Alkaline lakes like Lake Van, the world's largest soda lake, situated in eastern Anatolia, Turkey, can serve as a modem hydrochemical analogue of the early alkaline ocean. With the establishment of deep subduction zones and growing continental surface areas, the alkaline ocean gradually gave way to the halite ocean. Two modes of calcium

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carbonate precipitation have been recognised in modern environments: in vivo and early post mortem, often combined in the same mat. Similar processes of calcification in Precambrian microbial mats and of carbonate production, in similar environments, can be envisaged (section 6.4). From the Neoarchaean on, coccoid cyanobacteria must have calcified in a similar way to modern cyanobacteria. Dolomitisation at that time was probably also microbially induced (Kazmierczak and Altermann, 2002). Precambrian stromatolites witness the early evolution of life as far back in Earth history as 3.5 Ga. The sizes of stromatolitic bioherms vary by several orders of magnitude, from millimetres to tens or hundreds of metres in thickness and hundreds of kilometres of lateral extent. Although such giant stromatolitic reefs do not exist today and the variety of shapes of stromatolites declined by orders of magnitude since the Proterozoic, accompanied by the retreat of stromatolites into ecological niches with few or no grazers and competitive organisms, Precambrian stromatolites must be studied with an actualistic approach (Donaldson et al., 2002). Various definitions of stromatolites compete with each other, some including and some excluding chemical precipitates (Altermann, section 6.5). The question of definition is not a purely semantic problem. Hofmann (2000) emphasises that it is particularly relevant to the Archaean, because stromatolites are used to demonstrate the existence and evolution of early life on Earth and are now an objective in exobiology. Little agreement has been reached on the classification of stromatolites since their first description by Hall (1883). Because the immense variety of stromatolite forms can not be explained by biological factors alone, a purely biological (Linnean) stromatolite classification appears to be poorly justified. Environmental factors like sediment supply, water depths, currents and waves influence the mode and form of stromatolitic growth throughout the "life time" of a stromatolite. Linnean taxonomy of stromatolites, based on morphology, is untenable and depends to a large extent on the degree of the exposure or the specific sample under investigation. Geometrical stromatolite description schemes were therefore proposed by various authors. The most important and useful among them are those by Logan et al. (1964) and Hofmann (1973). Hofmann (1973, 2000) proposed classification of stromatolites using a pyramidal diagram of triangular base, where each corner of the pyramid represents one of the four main stromatolite-forming processes. Hofmann (1973) also defined morphological types of stromatolitic lamination such as flat, concave-up, convex-up, globoidal, nodular, columnar (branched and unbranched) and oncoidal. A metric scale is applied to these descriptions to portray the size of the structure (section 6.5). Although notable differences in stromatolite occurrences throughout the Precambrian have been noted, stromatolites are poor stratigraphic markers. A biostratigraphic subdivision of the Proterozoic or Archaean, based on stromatolites is not possible. Microfossil preservation in Archaean stromatolites is extremely exceptional and because of the simple morphology and very slow evolution of Precambrian microbiota, they also do not allow for a stratigraphic subdivision of the strata. Precambrian stromatolites can, however, serve as environmental indicators if treated cautiously, because they can occupy a wide range of Precambrian settings. Normally shallow pools, and peritidal realms to evaporitic and hydrothermal basins are envisaged as the main sites of stromatolite formation; however,

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591

Precambrian microbial mats which thrived in deep water have been reported (Simonson et al., 1993). Thus, Precambrian microbial mats can occur in all marine and lacustrine environments, from deep to shallow (section 6.5). Microbial colonisation of Neoarchaean supratidal flats was described by Altermann and Herbig (1991) and colonisation of terrigenous clastic sedimentary environments in the Proterozoic by Schieber (1998) (see also section 7.9). Early life altered the atmospheric composition, from one dominated by CO2 to an oxygen-rich atmosphere (section 5.2) and thus affected also weathering patterns in terrestrial realms and sediment supply to basins. It has also changed the chemical composition of sedimentary rocks by producing vast carbonate platforms and left identifiable isotopic and morphological traces in almost all sedimentary rocks. Thus, the understanding of earliest Precambrian geology as the context for the appearance and evolution of life has profound consequences for astro- or exobiology (Westall, section 6.6). It can be expected that if life had also developed on other planets and become extinct thereafter due to endogenic or exogenic processes, it should have left some traces of its existence in the sedimentary rock record of these planets. But is an expectation of extinct (or extant) extraterrestrial life not a mere speculation? The daring dream of mankind is to find extraterrestrial intelligent life. The chances for it are, however, infinitesimal. Finding primitive, microbial life seems nevertheless not entirely unrealistic, even in our solar system. Because of the history of our solar system, this life, if ever present, can only be most primitive, in best case prokaryotic. The most basic requirements of life, like the abundance and availability of the elements C, H, O and N, the existence of liquid water, and accessible energy were probably realised on several terrestrial planets in the early history of the solar system. Apart from the Earth, Venus, Mars and even Mercury, despite its proximity to the Sun, most probably had liquid water at some stage of their early evolution. Also, Europa and Callisto, the moons of Jupiter, might host liquid water even at present. This prerequisite of life was lost, however, on all these planets, with the exception of the Earth, at about 3.5-4.0 Ga. Carbon and other elements combined to form polycyclic aromatic hydrocarbons, and other complicated organic molecules, including amino acids, are abundant in impacting comets, meteorites and asteroids. Potential energy sources (light, hydrothermal, volcanic, impact energy) were also readily available on these planets to satisfy the basic requirements of life. If life ever arose there, however, it would not have had much time nor favourable conditions to develop. Nevertheless, it might have left traces with a higher preservation potential than on Earth, because of the lack of plate tectonic processes. Potential habitats for the search for extraterrestrial life are equivalent to those on Earth. Finding remnants of extraterrestrial microbial life would certainly teach us a great deal about the origin of life on the early Precambrian Earth and in the Universe (section 6.6).

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The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) 9 2004 Elsevier B.V. All rights reserved

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SEDIMENTATION

7.1.

THROUGH

TIME

INTRODUCTION

EG. ERIKSSON, A.J. BUMBY AND M. POPA Earth's oldest known supracrustal rocks, from the Isua greenstone belt in West Greenland (sections 2.2 and 2.3), are interpreted as predominantly mafic volcanics, but include subordinate > 3.8 Ga chert and banded iron-formation (section 5.4), and rare intrabasinally derived conglomerates and sandstones (Fedo et al., 2001). Detrital zircons/> 4 Ga have been obtained from the Jack Hills metasedimentary belt, Yilgarn craton, Australia (Nelson, 2001b). Chemical and clastic sedimentation were thus active during creation of Earth's earliest known continental crust, and must have been preceded by oceanic volcaniclastic reworking processes and biochemical deposition (see also section 7.3).

Basic Principles Actualism, the principle that the same processes and invariant natural physical, chemical and biological laws applied in the (Precambrian) past as at present, provides an amplification of modern (i.e., non-gradualistic) uniformitarianism (Donaldson et al., 2002). This definition of actualism also encompasses catastrophic events such as bolide impacts; non-actualism relates essentially to speculations on early Hadaean processes (sections 1.2 and 3.6) and products (Donaldson et al., 2002). These authors demonstrate application of actualism to Earth's sedimentary record, bearing in mind variable rates and intensities of processes controlling weathering (sections 5.10 and 5.11), erosion, transport, deposition (section 7.11) and lithification. Although the relative rates of processes like mid-ocean spreading and subduction, weathering, continental crustal genesis (section 2.8), rotation of Earth (section 5.9), and atmospheric evolution (section 5.2) contrasted with those derived from Phanerozoic successions, the processes themselves were not significantly different (Eriksson et al., 2001a). Precambrian sedimentary structures and lithologies and their inferred genetic processes all have modern counterparts (section 7.2); however, there was significant temporal control on certain depositional settings, such as glaciogenic (sections 5.6 and 5.7) and aeolian erg (section 7.6) palaeoenvironments (Eriksson et al., 1998a). In the absence of land plants in the Precambrian, there is evidence that colonisation of shallow water (Schieber, 1998) and even continental environments (Eriksson et al., 2000) by microbial mats (sections 7.9 and 7.10) may have been significant. Bioturbation was absent in the Precambrian, leading to much better preserved shelf sediments than in PhanerozoicRecent successions (Eriksson et al., 1998b, 2001b).

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Similarities between Precambrian and Phanerozoic basin styles outweigh differences, and like their younger counterparts, evolution of Precambrian basins was controlled primarily by interaction of magmatic-thermal and plate tectonic processes, modified by eustasy and palaeoclimate (Eriksson et al., 2001a, b; Bose et al., 2001) (Table 7.1-1). The range of basin types and their preserved fills in the Precambrian record show no significant differences from Phanerozoic equivalents (Eriksson et al., 2001a). Plate tectonics in recognisable (Modern-Phanerozoic) form has been active since at least the Neoarchaean (section 3.6), and prior to c. 2.0 Ga mantle plumes (possibly global events; Nelson, 1998a; Condie, 1998, 2001a) (sections 3.2 and 3.3) were a significant primary influence on basin formation (Eriksson et al., 2001 a, b). Plumes probably disturbed "normal" subduction and arc systems and likely promoted highly variable plate movements rather than universally rapid plate migrations (cf. Catuneanu, 2001; Eriksson et al., 2001b). These variable rates and changing tempos of continental crustal growth were, in turn, the major controls on the secondary basin-forming factors of eustasy and palaeoclimate (Eriksson et al., 2001 a). Interdependence of Crustal Growth, Freeboard and Eustasy The expression of continental elevation above mean sea level, encapsulated in the freeboard concept (Wise, 1972) is used to infer approximately constant continental and oceanic areas and volumes since c. 2.5 Ga (constant freeboard model; Wise, 1974). The constant freeboard model is thus intrinsically bound to continental crustal growth rates (section 2.8). The freeboard concept rests on average global conditions; chronological and geographic variability thus result (Eriksson, 1999). Modern hypsometric curves are highly variable between continents (Fig. 7.1-1 ). The geological record of the NeoarchaeanPalaeoproterozoic continents supports diachronous continental crustal growth rates near c. 2.5 Ga (Eriksson, 1995), which would have produced analogous variable freeboard conditions between these ancient cratonic terrains. Post-Archaean continental crustal growth rates may have varied between 10% and 40%, resulting in concomitant freeboard variation up to c. 200 m (Schubert, 1988; Windley, 1995), values which are similar to many eustatic and relative sea level changes, particularly those due to mid-ocean ridge activity, glacioisostacy, geoid relief and local tectonism (Table 7.1-2). Eustasy, freeboard and crustal growth rates were thus interdependent variables in the Precambrian, as they still are today (Eriksson et al., 1999a). Precambrian Depositional Systems: Comparison to the Phanerozoic-Modern Wave- and storm-dominated shallow marine systems In the absence of useable fossils (see chapter 6) and bioturbation, distinction between Precambrian shallow marine and continental deposits, and particularly between inner shelf and fluvial overbank facies, can be problematic (Dott et al., 1986; Mueller and Dimroth, 1987; McCormick and Grotzinger, 1993; Donaldson and de Kemp, 1998; Schieber, 1998). These deposits are relatively common in the Precambrian rock record, vary from metrethick upward-coarsening parasequences to homogeneous successions 1.02-103 m thick,

7.1. h~troduction

595

Fig. 7.1-1. (a) Cumulative percentages of Earth's solid surface at different elevations relative to mean sea level are illustrated by the hypsometric curve, based on a histogram of elevations and depths. The continental freeboard is thus equivalent to the maximum elevation above mean sea level. (b) Hypsometric curves constructed for the six present-day continents, based on an interval from 200 m below mean sea level (edge of continental shelf) to a maximum continental elevation of 1000 m above this datum and normalising this interval to 100%. Both figures modified after Schopf (1980).

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Table 7.1-1. Evolution of selected Precambrian basins (after Eriksson et al., 2001a) Basin SE California, U.S.A.

Age (Ga) c. 0.75-0.6

Major influences on basin evolution Rodinia breakup--rift to passive margin; glacial palaeoclimatic influences

Chuar Group, Grand Canyon, SW, North America

c. 0.8-0.074

Interaction of tectonic and palaeoclimatic (greenhouse and icehouse) influences

Uinta Mountain and Big Cottonwood Groups, Northern Utah, U.S.A.

c. 0.8

Rifting related to Rodinia breakup; tropical palaeoclimate at low palaeolatitudes

Midcontinent Rift, Lake Superior region, U.S.A.

c.l.1

Predominantly mantle plume, lesser tectonic shortening from distant plate margins

Vindhyan, Bhandara craton, central India

c. 1.4-0.55

Rift and sag basin with plate margin compression; tectonically-driven cycles; eustasy significant

Belt, western North America

c. 1.45

Synsedimentary tectonism in intracratonic basin; arid palaeoclimate

Sao Francisco craton (three basins)

c. 1.7-0.65

Rift and sag basin; full Wilson cycle with possible plume influence as well as palaeoclimate (glacial) played role in younger two basins

Lake Superior region, U.S.A.

c. 2.4-2.2

Full Wilson cycle; significant palaeoclimatic influence (icehouse, greenhouse, arid)

Hurwitz, Western Churchill Province, Northern Canada

c. 2.45-< 1.9 Tectonism predominated over significant magmatism, eustasy and palaeoclimate. Related to two supercontinental events

Karelian, Fennoscandian Shield

c. 2.45-1.9

Significant roles for tectonism, eustasy and palaeoclimate (icehouse, greenhouse, arid)

Huronian, Superior Province, Canada

c. 2.4-2.2

Partial Wilson cycle--rift and passive margin; tectonism and palaeoclimatic (glacial) cyclicity

Transvaal, Kaapvaal craton, South Africa

c. 2.67-2.1

Magmatism, eustasy and palaeoclimate predominant over tectonism; intracratonic basin

Raquette Lake, Slave Province, Canada

c. 2.6

Tectonic extension in active backarc basin floored by continental crust; magmatism important

Belingwe greenstone Belt, Zimbabwe craton

c. 2.7-2.65

Foreland basin; evidence for synsedimentary horizontal tectonism

Witwatersrand, Kaapvaal craton, c. 3.0-2.7 South Africa

Retroarc foreland system on a young and less rigid lithosphere; flexural tectonics important

Mallina, Pilbara craton, Australia

c. 3.0-2.94

Tectonic shortening within an intracratonic basin

Isua greenstone belt, West Greenland

> 3.7

No direct evidence for the role of plate tectonics; magmatism predominant

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Table 7.1-2. Sea level changes: causative mechanisms, extents, rates and durations (modified after Reading and Levell, 1996) Mechanism Mid-ocean ridges Orogeny Sedimentation onto seafloor Hot-spot seafloor movements Intraplate stress Flooding/desiccation, small basin Local tectonism Glacioisostasy (rebound) Glacioeustasy (including hydroisostatic effect) Tsunamis and landslides Geoid relief

Max. size (m) 350 70 60 100 100 15

Av. rate (mm ky -1 ) 7.5 1.0 1.1 Very slow 10-100 Instantaneous

Time period (My) 70 70 70 100

1000 250 150

10,000 10,000 10,000

< 10 < 0.1 < 0.1

Instantaneous 5,000

Hours < 0.1

27 250

10

< 0.013

and tend to have very uniform associations of sedimentary structures (Cant and Hein, 1986; Tirsgaard, 1996). Although Precambrian inner and outer shelf deposits strongly resemble their younger equivalents, most studies have concentrated on the nearshore sandstones, and there is some evidence for more uniform storm systems in the Precambrian (Chakraborty and Bose, 1992; Tirsgaard and SCnderholm, 1997). Thick and monotonous Precambrian deposits may also reflect wide, low-angle shelves, whose development was related to high denudation rates due to atmospheric composition (sections 5.2, 5.3, 5.5, 5.10 and 5.11) and vegetation-free landscapes (Els, 1998) (section 7.8). The lack of bioturbation in the Precambrian provides much better preservation of fine outer shelf facies than in the Phanerozoic-Modern record. Shoreface deposits from the Precambrian strongly resemble their younger counterparts (Soegaard and Eriksson, 1985; Bose et al., 1988; Walker and Plint, 1992). Within Archaean greenstone belt settings, there is commonly a rapid transition from alluvial facies into high energy shoreface deposits, with varying degrees of tidal action (Mueller and Donaldson, 1992a; Corcoran et al., 1998; Mueller et al., 2002b) (section 7.3). Detailed analyses of shoreface dynamics and depositional architecture are uncommon, with several good Indian studies (Bose et al., 1988; Chakraborty and Bose, 1990). Inferred Precambrian shoreface deposits tend to be significantly thicker and to contain a more limited set of sedimentary structures than Phanerozoic successions; these are difficult to explain, and more uniform circulation systems and a delicate balance between subsidence and sedimentation have been proposed (Eriksson et al., 1998b and references therein). For foreshore deposits, poor preservation and limited studies apply to all ages, with limited Precambrian examples (e.g., Vos and Eriksson, 1977; Eriksson, 1979; Bose et al., 1988; Bose and Chakraborty, 1994) suggesting overall similarity to younger deposits. Precambrian barrier island-lagoon-washover systems appear to be rare (e.g., Eriksson, 1979 for

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an Archaean example), which possibly reflects identification and preservational problems (Eriksson et al., 1998b). Tide-dominated shallow marine systems In the absence of usable fossils, Precambrian tidal deposits and fluvial sediments may share many characteristics and even "diagnostic" sedimentary structures (Alam et al., 1985; Tirsgaard, 1993), and such facies may also be complexly interbedded in preserved coastline deposits (Eriksson et al., 1995). Thick, mature sandstone successions of tidal origin, with limited suites of sedimentary structures, are common in the Precambrian record; those described in the literature are mainly sandwave deposits from the shoreface to inner shelf regimes, whose internal architecture is analogous to modern examples (Eriksson et al., 1998b and references therein). Tidal sand ridges are rarely described for the Precambrian (e.g., Johnson, 1977; Mueller et al., 2002b). Apart from a few studies which indicate similarity with Phanerozoic equivalent deposits (e.g., Eriksson, 1979; Eriksson et al., 1981; Williams, 1998b), tidally influenced Precambrian shelves are thought to have lacked barrier islands, tidal inlets and tidal deltas (e.g., Harris and Eriksson, 1990; J.M. Jackson et al., 1990; Ghosh, 1991; Lindsay and Gaylord, 1992). However, Williams (1998b) used tidal rhythmites preserved in ebb-tidal deltas to analyse Neoproterozoic Earth-Moon dynamics (section 5.9). Muddy back-barrier subtidal deposits, commonly found in the Modern record (Elliot, 1986) are not known from the Precambrian, where vegetation and bioturbation would also not have been relevant. Precambrian lagoonal deposits exhibit subtidal-intertidal sandstones and tidal channel deposits, with apparently rare flood-tidal and washover facies (e.g., Eriksson, 1979; Deynoux et al., 1993). There is a strong resemblance between Precambrian and younger tidal flat sediments. Meandering tidal channels were notably absent in the Precambrian settings, where poorly confined sand sheets tended to develop, which were commonly associated with and which are easily confused with fluvial and braid-delta sheet sandstones (Tirsgaard, 1993; Eriksson et al., 1995; Els, 1998). Indications that Precambrian tidal, wave and storm shelf dynamics may have been more uniform (Eriksson et al., 1998b and references therein) must be tempered with an appreciation that these shallow marine deposits are commonly uniform and homogeneous over large areas and thicknesses, and lack fossils. Interpretations are further complicated by epeiric sea palaeoenvironments, which at certain periods, particularly in the NeoarchaeanPalaeoproterozoic due to enhanced crustal growth rates (section 2.8), transgressed onto large portions of developing, low-freeboard cratons (section 7.7). A combination of braided fluvial, braid-delta and tidal flat depositional systems was common around the margins of these epeiric seas, and these facies are also to be found preserved in many greenstone deposits (Corcoran et al., 1998) (section 7.3). Eriksson and Simpson (section 7.5) examine the recognition and significance of Precambrian tidalites. Deltaic systems Although Precambrian delta deposits are known from the > 3.0 Ga Barberton greenstones (Heubeck and Lowe, 1994) to the Neoproterozoic, the palaeoenvironmental resolution of

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deltaic subenvironments recognised from the Phanerozoic-Modern records in Precambrian successions is hindered by the lack of faunal, floral and trace fossils, coal seams and bioturbation (Eriksson et al., 1998b). Recognition of Precambrian delta successions is based on the same non-biogenic diagnostic features found in younger deposits. Depositional processes and controls inferred from deltas of all ages are similar, suggesting that the necessary delicate balance between eustasy/accommodation space and sediment input (chapter 8) in their evolution has been relevant throughout Earth's history. The absence of vegetation led to poor channel bank stability on Precambrian delta plains, with concomitant braided systems developing (cf. Schumm, 1968; Miall, 1992); tidal reworking locally resulted in point-bar deposits forming in some Precambrian distributaries (Eriksson, 1979). Precambrian deltaic successions provide evidence for the same variable relationships between river regime, wave energy and tidal range as found in Phanerozoic-Modern examples (Siedlecka et al., 1989; Bhattacharya and Walker, 1992). Discrimination of the transition from prodelta to open shelf sedimentation for Precambrian delta deposits is best estimated by the proximal limit of banded iron-formation (section 5.4), as precipitation of iron was a common background chemical "rain-out" sediment prior to c. 2.0 Ga (Barrett and Fralick, 1989). Bioturbation fulfils a similar role in Phanerozoic-Modern equivalent settings (Eriksson et al., 1998b). Thickness of deltaic deposits is a significant differential between Precambrian and younger settings: only rarely do younger successions exceed c. 150 m (Bhattacharya and Walker, 1991, 1992), whereas the Archaean Barberton greenstone belt has preserved deposits over 400 m thick (Eriksson, 1979), and in the Neoproterozoic Basnaering delta complex of northern Norway, a thickness of 3.5 km is observed (Siedlecka et al., 1989). Additionally, Precambrian delta deposits tend to be more immature texturally, contain more conglomerates, and often lie close to major faults; these characteristics support an active tectonic setting for many Precambrian deltas (Eriksson et al., 1998b). The lack of vegetation, interacting terrane amalgamation and accretion during cratonic growth, and rapid denudation regimes on emerging continents likely played a role in these tectonically active and high sedimentation rate type deltaic deposits (Eriksson et al., 1998b). The large, high-discharge braidplain systems resulting from the combination of early Precambrian atmospheric composition (section 5.2), concomitant high weathering rates (sections 5.10 and 5.11 ), and enhanced erosion rates due to a lack of vegetation and well-developed soil profiles led to enormous braid-delta systems being common along coastlines (e.g., Els, 1998), particularly around the margins of epeiric seas (Eriksson et al., 2002a). They were often subject to significant tidal reworking (section 7.7). Alluvial systems As stated above, broad channel systems with abundant bedload and high discharge rates are synonymous with Precambrian continental palaeoenvironments. A general lack of channel bank stability and essentially braided patterns were thus predominant (Schumm, 1968; Cotter, 1978; Long, 1978); however, discriminating fluvial style requires multi-faceted investigation (Jackson, 1978). Precambrian braidplain systems were almost certainly much larger than younger counterparts. The enhanced runoff rates of Precambrian fluvial sys-

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tems, due to an absence of root binding and scarce soil development allied to weathering regime, would have made these systems more sensitive to palaeoclimatic changes; hence, ephemeral rivers likely formed in a broader climatic range than today (Tirsgaard and Oxnevad, 1998). Semi-perennial fluvial systems active under relatively humid palaeoclimates may have been a style unique to pre-vegetational times (Tirsgaard and Oxnevad, 1998). Suffocation of alluvial systems within Archaean greenstone belt settings by rapid and large additions of pyroclastic debris was common, and hyperconcentrated flood flow and sheetflood deposits were often a consequence (Mueller and Corcoran, 1998, 2001; section 7.3). Long (section 7.8) provides a brief review of pre-vegetational fluvial systems. Alluvial fan deposits are not that common in the Precambrian rock record, as these proximal near-source systems are less likely to be preserved than distal deposits within subsiding basins (e.g., Els, 1998; see, however, Mueller and Corcoran, 1998). Williams (2001) describes Neoproterozoic fan deposits from northern Scotland, from which a fan radius of c. 50 km and a catchment of c. 1.8 • 104 km 2 have been estimated. As fans commonly pass downstream into fluvial braidplains, discrimination between the two systems is often problematic, even more so in Precambrian successions (e.g., Els, 1998). In Modern systems, there is a well-defined gap in slopes between rivers (maximum gradient of 0.007 mm -1) and alluvial fans (slopes > 0.026 mm - l ) (Blair and McPherson, 1994). Palaeohydrological parameters estimated from the c. 1.8 Ga Wilgerivier Formation, Kaapvaal craton lie almost precisely in this gap, and may reflect an association of small fault-bounded basins, aggressive weathering and intermittent torrential rainstorms (Van der Neut and Eriksson, 1999). Lacustrine systems Identified lake deposits within the Phanerozoic rock record are sparse, and they are difficult to discriminate from analogous shallow marine deposits (Picard and High, 1972; Tucker, 1991; Martel and Gibling, 1991; Pratt, 2001; see also discussion in Eriksson et al., 1998b). The cyclicity common in lake sediments is mirrored by many marine mesosequences (Friedman et al., 1992) and stromatolites (section 6.5) occur in both (Hallam, 1981). Palaeontology and geochemistry, particularly the presence of certain evaporite minerals considered diagnostic of lakes, are most often used for identification (Reeves, 1968; Hallam, 1981). These are of limited use in Precambrian deposits where invertebrate and plant remains are lacking, and where evaporites are either destroyed or pseudomorphed. However, Martini (1990) was able to use geochemistry and such pseudomorphs to interpret alkaline playa deposits at c. 2.2 Ga in the Transvaal basin, Kaapvaal craton. Rhythmites in colder climate lakes (Sturm and Matter, 1978) may also result from suspension sedimentation and aeolian deposition (Rogers and Astin, 1991). Wind-formed waves and impounding of water masses in lakes due to wind stress may simulate microtidal marine coastlines (Galloway and Hobday, 1983), but marine swells and significant and sustained lunar tides will be absent (Friedman et al., 1992). Eriksson et al. (1996) used boron as a palaeosalinity estimate in discriminating between Palaeoproterozoic epeiric and lacustrine units at the basinal scale.

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601

Precambrian lake successions from the Archaean through to the Neoproterozoic are commonly identified using associated alluvial and aeolian deposits, and they do not exceed c. 600 m in thickness, with several tens of metres being much more common (Unrug, 1984; Kalliokoski, 1986; Donnelly and Jackson, 1988; Karpeta, 1989; Aspler et al., 1994; Fairchild and Hambrey, 1995). Wood (1980) discusses possible lake deposits within Archaean greenstone belts, and Karpeta (1989) identified an alkaline lake deposit within the 2.7 Ga Ventersdorp Supergroup, Kaapvaal craton, where partly magadiitic cherts, stromatolites, good evidence for evaporite minerals, evidence for desiccation and weak tidal activity support his model. Geochemical data combined with physical sedimentary structures and evaporite remains have resulted in saline lake deposits being identified readily in Mesoproterozoic (Collinson, 1983; Donnelly and Jackson, 1988) and Neoproterozoic (Kalliokoski, 1986; Porada and Behr, 1988; Fairchild and Hambrey, 1995) successions. Discrimination of hydrologically open and more permanent lakes is much more difficult in the Precambrian record (e.g., Unrug, 1984; Winston, 1986; Schieber, 1998; Mueller and Corcoran, 1998). Perhaps the most enigmatic non-saline Precambrian lake deposit is that interpreted by Aspler et al. (1994) from the > 2.09 Ga Whiterock Member, Kinga Formation (Hurwitz Group, Nunavat, Canada). Geometry, allied to analysis of ripples and parallel stratification suggest an extent of 100,000 km 2, water depths averaging 2 cm-2 m, and there is no evidence for tidal or desiccation influences through a preserved thickness up to 400 m (Aspler et al., 1994). The preponderance of partly arid or saline lakes within the Precambrian record may merely reflect their easier identification compared to more permanent lacustrine basins. Alternatively, the increase in evaporite and redbed deposits after c. 2.3 Ga in the Precambrian rock record is at least consistent with the apparent abundance of saline lake successions (Eriksson et al., 1998b, and references therein). Desert systems

In a recent review, Eriksson and Simpson (1998) note that large scale aeolianites appear to be absent from the c. > 2.2 Ga geological record, and that they become common and widespread after approximately 1.8 Ga. However, ventifacts are known from the c. 3.0-2.8 Ga Witwatersrand Supergroup, Kaapvaal craton. In this volume, Simpson et al. (section 7.6), discuss the c. 2.6 Ga sand sheet deposits of the Minas Supergroup in Brazil, and document the evolution of large ergs from c. 1.8 Ga onwards. Eriksson and Simpson (1998) suggest that the apparent lack of aeolianites older than c. 2.2 Ga reflects fluvial reworking of non-vegetated floodplains or destruction of coastal sand sheets through transgressions, as well as possible non-recognition. Eriksson et al. (1998b) speculate that early Precambrian palaeoclimates may have influenced wind systems and aeolian transport. Rautenbach (2001) has used sophisticated climatic software packages to model the effects of the enhanced rotation of the early Earth (section 5.9) on wind regimes. The only reliable rotation rate data goes back to c. 0.9 Ga (Williams, 1998b), and applying the concomitant 18.2 h diurnal cycle at 900 Ma, Rautenbach (2001) finds that there would have been a significant equatorwards latitudinal shift of planetary scale circulation cells, combined with a reduced wind speed throughout the palaeoatmosphere. It is uncertain to what degree these parameters may be extrapolated to the Archaean-early Palaeoproterozoic time period, in the

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absence of accurate palaeorotational data prior to c. 0.9 Ga. Eriksson and Simpson (1998) emphasise the important role of early supercontinentality in the evolution of the first ergs at approximately 1.8 Ga. Glacial systems Earth appears to have been non-glaciated for large parts of the Precambrian period, with global refrigeration events in the Palaeoproterozoic and the Neoproterozoic (Young, 1991) (sections 5.6-5.8). Archaean glaciogenic deposits have been recognised from the Kaapvaal craton (c. 3.0 Ga Pongola Supergroup, likely deposited in the greater Witwatersrand basin) and from the basement to the Stillwater Complex in Montana (Page, 1981; von Brunn and Gold, 1993). Widespread glacial deposits such as those from the Palaeo- and Neoproterozoic are very useful for understanding Precambrian global sequence stratigraphy and for inferred supercontinent reconstructions (e.g., Aspler and Chiarenzelli, 1998). Glacigenic and periglacial processes and products from rocks of all ages appear to have been very similar, more so than for many other palaeonvironments (Eriksson et al., 1998b; Table 1 and references therein). However, the absence of land vegetation in the Precambrian would have favoured preservation of glaciomarine deposits. Loess deposits around ice sheets, as were common in the Quaternary, would most likely have been uncommon in the global absence of land plants, but examples are identified from both Palaeo- and Neoproterozoic (Edwards, 1979; Fralick and Miall, 1989). The proposed causes of Precambrian glaciation tend to outnumber the events themselves (Young, 1991), but most would agree that variation in atmospheric CO2 contents allied to c. 400 My long supercontinent cycles and superimposed on the secular increase in solar luminosity and decreasing CO2 were, collectively, of primary importance in determining glacial and non-glacial states (Eriksson et al., 1998b, and references therein). In section 5.6, Young examines critically the snowball Earth hypothesis and details the two great Precambrian glaciations, and in section 5.8 Frimmel discusses the second, Neoproterozoic event. Williams (section 5.7) examines the paradox of low latitude marine glaciation, open seas and strong seasonality implicit in Precambrian glaciation. Reviews of Precambrian glacial rocks are given by Hambrey and Harland (1981, 1985), Eyles (1993) and Eyles and Young (1994).

7.2.

SEDIMENTARY STRUCTURES: AN ESSENTIAL KEY FOR INTERPRETING THE PRECAMBRIAN ROCK RECORD

J.A. DONALDSON, L.B. ASPLER AND J.R. CHIARENZELLI Sedimentary structures provide invaluable clues regarding the processes of transport and deposition in ancient rocks, and the physiochemical conditions during and shortly after sedimentation. They are particularly important for the interpretation of depositional settings of Precambrian successions that lack fossils unique to specific environments. Although biogenic structures such as biofilms (sections 7.9 and 7.10) and stromatolites (section 6.5) The Precambrian Earth: Temposand Events Edited by P.G. Eriksson. W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

7.2. Sedimentary Structures

603

may serve as environmental indicators (they are most prolific in shallow water and supratidal environments), and although traces of soft-bodied burrowing organisms do occur in some Neoproterozoic sequences, the lack of diagnostic body fossils requires a greater reliance on sedimentary structures in studies of Precambrian strata. In partial compensation however, the lack of burrowing organisms has commonly rendered perfect the preservation of structures which, in comparable Phanerozoic environments, have been destroyed due to extensive bioturbation. Not only do sedimentary structures provide excellent evidence for past conditions, they are paramount in establishing geometric relationships of deformed strata through their utility as tops indicators, and also strain markers in deformed sequences (section 7.4). Reliance on the principle of actualism permits reliable application of our understanding of modern examples to records of the past, extending back to the earliest sedimentary rock record (Donaldson et al., 2002). Well-illustrated treatments of sedimentary structures are available in classical texts such as Pettijohn and Potter (1964), Bouma (1969), Reineck and Singh (1980), Tucker (1982), Scholle and Spearing (1983), Scholle et al. (1983), Allen (1984, 1985), Selley (1985), Fritz and Moore (1988) and Collinson and Thompson (1989). Rather than reproducing examples of many relatively well-understood sedimentary structures, we herein present some that are less common, emphasising links between those in Precambrian successions and counterparts from Quaternary and modern settings. (See Figs. 7.2-1-8.)

Fig. 7.2-1. Glacial and periglacial features. (a) Frost-shattered boulder atop bedrock glaciated during the Wisconsinan (final episode of Pleistocene glaciation). An initial hairline fracture has been opened as a result of repeated freeze-thaw cycles during the past 8000 years (up to 100 cycles per year during winter in this area). Cobalt, Ontario, Canada.

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Fig. 7.2-1 (continued). (b) Glaciated polished and striated (during Pleistocene glaciation) outcrop of Palaeoproterozoic Gowganda Formation (Huronian Supergroup) showing bedding-parallel section through a frost-shattered boulder with a fracture that was opened and widened by frost action, and completely filled with silt/sand/gravel matrix. Uniformity of crack width shown by the opened fracture suggests progressive widening through repeated freeze-thaw cycles, comparable to the space created between the matching halves of the boulder shown in (a). Many clasts in the outcrop show similar uniformity of crack widening, reflecting in situ freeze-thaw creation of such gaps in clasts that rested on Archaean glacial pavement, before Palaeoproterozoic infilling due to flooding by outwash streams at the front of a stagnant ice sheet. Cobalt, Ontario, Canada. (c) Pebble-armoured clast, inferred to represent a frozen ball of outwash gravel. Because pebbles protrude from the clast, derivation by erosion of a previously indurated conglomerate is precluded. Palaeoproterozoic Gowganda Formation, Cobalt, Ontario, Canada.

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605

Fig. 7.2-2. Aligned stromatolites. (a) Modern domal stromatolites, elongate perpendicular to shoreline, Shark Bay, Western Australia. Elongation is due to the strong ebb-flow action of tidal currents in an intertidal environment. (b) Elongate domal stromatolites, Mesoproterozoic Dismal Lakes Group, Nunavut, Canada. Although these are presumed to be aligned perpendicular to palaeoshoreline, measurements of elongation through a 30 m section indicate that stromatolites in some units are aligned perpendicular to the prevailing trend, implying that some ovoid-in-plan-view stromatolites owe their orientation to shore-parallel currents.

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Fig. 7.2-2 (continued). (c) Oblique-inclination columnar stromatolites, Mesoproterozoic Dismal Lakes Group, Nunavut, Canada. Similar groups of inclined stromatolites in the Siyeh Formation, Montana, U.S.A., were inferred to represent initially vertical columns toppled by a hurricane. The lack of fracturing at the horizon at which tilting was initiated, the uniformity of tilt, and the asymmetric nature of the internal laminations collectively indicate that the inclination is a primary growth characteristic. Solar control is considered unlikely because of common reversals in direction of inclination in the Dismal Lakes occurrences. More likely, growth was inclined toward a prevailing longshore current (or toward the strongest of the ebb/flow tidal currents). Such currents would provide nutrients for biofilms responsible for the laminations; long-term current reversals explain the switches in direction of inclination. Note also that conical laminations are developed above domal laminations within some of the columns, which poses a problem for the proposition that conical-columnar stromatolites comprise a distinct "form genera".

A full appreciation of the wide range of primary and secondary structures is essential for field geologists to interpret past conditions (section 7.4), especially in the computer age. Such appreciation is best acquired through the study of modern analogues. The importance of developing and maintaining an understanding of basic field relationships was presented eloquently several decades ago by Francis Pettijohn in his essay in defence of field geology (Pettijohn, 1984). We hope that this brief photo essay will promote the continuation of detailed field studies of sedimentary successions. Our understanding of geological history, as provided by occasionally arcane clues in the rock record, is far from complete (e.g., section 7.5).

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607

Fig. 7.2-3. Microdomal stromatolites. (a) Microdomes developed on flexible leathery biofilm sheet capping unindurated carbonate sand (readily penetrated by shovel). Coastal Sabkha, Abu Dhabi, south side of Persian Gulf. (b) Bedding surface showing similar microdomes on a selectively silicifled biofilm in dolostone of the Mesoproterozoic Dismal Lakes Group, Nunavut, Canada. Light-toned (silicified) microdomes poke through a thin cover of dark toned (unsilicified) dolostone. Lighttoned amalgamated patches mark areas in which interspace dolostone has been silicified. The light-toned sinuous linear feature is a vertical, silica-filled fracture.

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Fig. 7.2-4. Beachrock intraformational conglomerate. (a) Modem beachrock, Heron Island, Great Barrier Reef, Queensland, Australia. Surficial aragonite cementation of carbonate sand within the intertidal zone has created a carapace of solid rock, up to 3 m thick, overlying unindurated carbonate sand. Cyclones occasionally rip up angular blocks, separating them along early formed orthogonal joints. These blocks become rapidly re-cemented to form distinctive intraformational conglomerates. (b) Beachrock zones exposed above present waterline of Hudson Bay in dolarenite of McLeary Formation, Belcher Group (c. 2.1-1.8 Ga), Belcher Islands, Nunavut, Canada. By analogy with (a), these successive zones are inferred to mark recurring strandlines.

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Fig. 7.2-5. Biofilm structures. (a) Modern biofilm structures formed in supratidal zone, Belcher Islands, Nunavut, Canada. The originally continuous sheet has broken into fragments in a roughly orthogonal pattern as a result of desiccation, and most margins of the still-flexible fragments have been curled downward. (b) Mudcurls reinforced with surface biofilms, which in part have been responsible for the extreme curling upon desiccation. Dried ephemeral pond in gravel pit near Ottawa, Canada. (c) Sinuous ridges, arranged in a locally orthogonal pattern, on upper surface of a sandstone bed, Mesoproterozoic Hornby Bay Group, Nunavut, Canada. These trail-like pseudofossils are attributed to the infilling of cracks in a desiccated biofilm sheet, analogous to that shown in (a) and (b). See also section 7.10.

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Fig. 7.2-6. (a) Modem tepee structures, shore of Deep Lake, Eyre Peninsula, South Australia. Sheets of halite have expanded on the edge of the lake due to evaporation during shore retreat. This growth was accommodated through upward arching along subparallel linear trends perpendicular to the shoreline (this orientation was probably controlled by lakewards groundwater flow). (b) Comparable tepee structure in dolostone, Mesoproterozoic Dismal Lakes Group, Nunavut, Canada. Gypsum and halite pseudomorphs are abundant in this facies.

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Fig. 7.2-7. (a) Soft-sediment folds in Pleistocene varved laminites. The shock-sensitive Leda clay, deposited in Champlain Sea during recession of the Wisconsinan ice sheet, is particularly susceptible to such deformation, especially as a result of liquefaction in response to earthquakes. West shore, Lake Timiskaming, Ontario, Canada. (b) Soft-sediment folds in argillites, Palaeoproterozoic Gowganda Formation (Huronian Supergroup), Haileybury, Ontario, Canada. Penecontemporaneous en-echelon sandstone dykelets prove the early soft-sediment deformational origin of these folds, which probably also formed in response to seismic disturbances.

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Fig. 7.2-8. Synsedimentary faults. (a) Soft-sediment normal faults in glaciofluvial outwash deposits, near Cobalt, Ontario, Canada. Deformation likely resulted from melt-out of buried stagnant ice that was detached from the Wisconsinan ice sheet during deglaciation. Note diffuse zone of sand along fault in centre of photo. (b) Soft-sediment faults in fluvial sandstones that are intercalated with volcaniclastic deposits of the Palaeoproterozoic Christopher Island Formation, Nunavut, Canada. This deformation likely records seismicity related to volcanism. (c) Synsedimentary fault in Archaean AIgoma-type banded iron-formation, Mesabi Range, Minnesota, U.S.A. Sinuosity of the fault, which is truncated near the top of the field of view, suggests post-faulting differential soft-sediment compaction. Terminations of most siliceous iron oxide-bearing beds (dark layers) are smoothly rounded, suggesting that they were in a gel-like state when the faulting occurred (photograph courtesy of Gordon Gross).

7.3. Archaean Sedimentary Sequences

7.3.

613

ARCHAEAN SEDIMENTARY SEQUENCES

EL. CORCORAN AND W.U. MUELLER

Introduction Archaean terranes are complex amalgamations of volcanic, sedimentary and plutonic rocks (e.g., section 2.4). The sedimentary component is significant because it reflects source and composition of the hinterland, preserves ancient weathering profiles (section 5.10), indicates depositional conditions, and provides evidence of large-scale geodynamic processes that collectively elucidate early Earth's history (section 3.6). Previous comprehensive studies concern Archaean lithofacies (Eriksson et al., 1994), general Precambrian basin attributes (Eriksson et al., 2001 b), clastic sedimentation patterns (Ojakangas, 1985), greenstone sedimentation (Lowe, 1994) and synrift and craton cover sequences (Eriksson and Fedo, 1994). This review emphasises clastic depositional systems that are generally related to large-scale tectonic regimes: (1) craton-cover sequences (synrift or stable platform), (2) volcano-sedimentary sequences (synorogenic) and (3) molasse sequences (late orogenic). These sequence types record the principal stages during which basins formed, although there are numerous hybrid basin settings or subsettings (Ingersoll, 1988). Classifying Archaean remnant basins with respect to their precise tectonic setting is often problematic, but the selected depositional sequences contain features that enable basin distinction.

Craton-Cover Sequences (Synrifi and Stable Platform) Craton-cover sequences, which develop on stable platform and in synrift settings (Eriksson and Fedo, 1994), represent discrete intervals of time, wherein synrift sequences occur with the onset of extension and with subsequent rifting associated with volcanism, and wherein stable platform deposits mark a stage of erosion of sialic basement commonly constrained to passive margins. These stages may be evolutionary, but complex structural features and hiati associated with Archaean terranes make this determination problematic. Stable platform deposits consist predominantly of quartz arenite, sandstone and banded iron-formation, but carbonate, conglomerate, siltstone, and mudstone may be major components. Synrift deposits contain sedimentary lithofacies similar to platform counterparts, but are interstratified with volcanic lithofacies (section 4.2) that may increase overall thicknesses to as much as 6 km. This rifting, considered an evolved stage of extension, is not only supported by interstratified sedimentary and volcanic rocks, but also by an abundance of dykes consistent with crustal attenuation, and well-preserved coarsening-upward sedimentary sequences reflecting tectonic activity. Craton-cover sequences generally overlie granitoid basement, although some rest unconformably on volcanic rocks. The Precambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Depositional settings Craton-cover sequences are characterised by a common stratigraphy, which generally includes from base to top, quartz arenite + conglomerate • stromatolite-bearing carbonate + sandstone -+- siltstone • iron-formation, although the stratigraphy may be reversed in cases of coarsening-upward sequences (Figs. 7.3-1a-d). Sedimentary structures and changes in facies associations are consistent with either fluvial (Srinivasan and Ojakangas, 1986; Donaldson and de Kemp, 1.998) (section 7.8) or shallow-water settings (Eriksson et al., 1981; Thurston and Chivers, 1990; Fedo and Eriksson, 1996; Pickett, 2002). Srinivasan and Ojakangas (1986) interpreted the quartz arenites of the c. 3.2-3 Ga Bababudan Group, India, as braided fluvial plain deposits based on interstratification with subaerial volcanic flows, low variance calculations from palaeocurrent indicators, and abundant trough cross-beds. In contrast, various combinations of hummocky cross stratification, herringbone cross-beds, reactivation surfaces, mudstone laminae, symmetrical ripples, and complex cross-strata reported for quartz arenites, sandstones and siltstones from the c. 2.9-2.8 Ga Beniah and Bell Lake Formations (Pickett, 2002), the craton-cover sequence of the c. 3.0 Ga Buhwa belt (Fedo and Eriksson, 1996) and the c. 2.9 Ga lower part of the Witwatersrand Supergroup (Eriksson et al., 1981) are more consistent with shallow-water deposition affected by waves and tides (Figs. 7.3-2a, b) (section 7.5). Collectively, features of this shallow-water assemblage are consistent with proximal to shoreline (conglomerate; Pickett, 2002), tidal shelf and shoreface, (i.e., quartz arenite; Thurston and Chivers, 1990; Eriksson and Fedo, 1994; Fedo and Eriksson 1996; Pickett, 2002), and shallow-water shoreface to proximal offshore (sandstone, siltstone; Pickett, 2002) settings. Stromatolitic carbonate (sections 6.4 and 6.5) reflects a shallow subtidal to intertidal setting (Wilks and Nisbet, 1988; Beukes and Lowe, 1989) and banded iron-formation (section 5.4) represents orthochemical (Trendall, 2002) and/or late-stage diagenetic alteration in relatively deep water. Where located at the top of craton-cover successions, banded iron-formation is commensurate with water level rise and drowning of the shelf platform. Distinct beds containing both wavy and planar sandstone, and siltstone and mudstone (iron oxide-rich) laminae may, in some cases, represent tidal influence (Fig. 7.3-2c; Pickett, 2002), as they are remarkably similar to other Precambrian (Williams, 1998b; section 5.9) and younger (Nio and Yang, 1991 ) tidalites (section 7.5). Overall fining-upward sequences record a transition from inner shelf to below-wavebase settings (Fedo and Eriksson, 1996), whereas coarsening-upward sequences are associated with eustatic sea level changes and large-scale tectonism. Both basin uplift and repeated fault uplift of the basement may cause coarsening-upward sequences. An example of an Archaean synrift craton-cover sequence characterised by successive coarseningupward sequences is the c. 2.9-2.8 Ga Beniah Formation (Pickett, 2002). Pickett (2002) illustrated how an estuary-embayment complex developed where the coast was fed by a fluvial system (Fig. 7.3-3). Faulting along the marine shelf caused pulses of tectonic uplift, resulting in periods of shallowing and deposition of coarser detritus (e.g., conglomerate and quartz arenite) over finer-grained counterparts. The abundance of mafic intrusions cutting the sedimentary sequence and the depositional contact between sedimentary deposits and

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Fig. 7.3-1. Representative stratigraphic sections of craton-cover sequences from (a) the c. 2.9-3 Ga Keeyask Lake succession, Superior Province (Thurston and Chivers, 1990; Donaldson and de Kemp, 1998), (b) the c. 3 Ga Buhwa greenstone belt, Zimbabwe (Fedo and Eriksson, 1996), (c) the c. 2.8-2.9 Ga Bell Lake Formation, Slave craton (Pickett, 2002), and (d) the c. 2.8-2.9 Ga Beniah Formation, Slave craton (Pickett, 2002). Modified from Pickett (2002).

overlying pillowed flows supports contemporaneous volcanism and sedimentation during crustal attenuation.

Volcano-Sedimentary Sequences (Synorogenic) The selected volcano-sedimentary sequences are associated with arc-related settings, in which lithofacies vary considerably according to basin type, such as back-arc, interarc, fore-arc or trench. Laterally equivalent subaerial and subaqueous sedimentary deposits are commonly interstratified with volcanic flows and volcaniclastic material (Barrett and Fralick, 1989; DiMarco and Lowe, 1989; Eriksson et al., 1994; Mueller and Corcoran, 2001). Synorogenic sequences may overlie volcanic (Eriksson, 1980) and granitoid (Mueller and Corcoran, 2001) basement unconformably. The complex interstratification of calc-alkaline volcanic and sedimentary lithofacies, gradational changes from shal-

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Fig. 7.3-2. Characteristics of the Beniah Formation quartz arenite and sandstone, and the Bell Lake Formation banded iron-formation lithofacies (craton-cover sequences). Large arrow indicates top. (a) Tabular sets of quartz arenite containing complex composite cross-strata. Seven sets are numbered and indicated with dashed lines. Scale, pen 15 cm long (small arrow). (b) Planar cross-bedded quartz arenite (P) with high angle planar foresets (F) and a mudstone drape at the top of the underlying bed (Md). Scale, pen 15 cm long. (c) Banded iron-formation with alternating bands of mudstone/siltstone-sized (M/S) and sandstone-sized (Ss) grains, resembling tidal rhythmites. Scale, pen 15 cm long.

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Fig. 7.3-3. Palaeogeographic reconstruction of the Beniah Formation illustrating the location of the sandstone-siltstone (SaSL), quartz-pebble conglomerate (QPSL), and quartz-arenite (QAL) lithofacies, and the iron formation (IFSL) and planar- to wavy-bedded sublithofacies along a shallow water coastal setting. For more details concerning the description of lithofacies, see Pickett and Mueller (2000) or Pickett (2002). Diagram modified from Pickett (2002). low water and subaerial coarse clastic deposits to deeper water turbidite deposits, multiple flow directions as determined from palaeocurrent indicators, and abundant synvolcanic plutons and dyke swarms in basement rocks support an arc-related tectonic setting. Depositional settings Volcano-sedimentary sequences are characterised by a plethora of lithofacies representing different environments of deposition, but a succession marking the gradation from shallow to deeper water sedimentation is commonly recorded (Eriksson, 1978; Barrett and Fralick, 1989; Smithies et al., 1999; Mueller and Corcoran, 2001). This sequence of sedimentation is represented by an up-section change: conglomerate + sandstone + siltstone/sandstone (turbiditic) -+- mudstone + iron formation + chert. In most cases, not all lithofacies are preserved, and abrupt lateral and vertical changes are common. Sedimentary structures and facies associations in the conglomerate lithofacies are consistent with mass flow, hyperconcentrated flood flow, talus scree and traction-current deposition on the proximal parts of alluvial fans, fan deltas and braided streams (Eriksson, 1978;

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Mueller and Corcoran, 2001), or represent deposition on subaqueous volcaniclastic aprons (Eriksson, 1982; DiMarco and Lowe, 1989). Subaerial coarse clastic deposits prograded onto sandy braidplains (Eriksson, 1978), coastal transition zones (Mueller and Corcoran, 2001), and shallow marine shoreface (Eriksson, 1980) settings, as represented by the sandstone lithofacies. The overlying siltstone/sandstone lithofacies, mainly characterised by complete and incomplete Bouma divisions, indicate the onset of subaqueous deposition on submarine ramp and fan settings (Eriksson, 1980; Barrett and Fralick, 1989; Krapez and Eisenlohr, 1998). The limited extent and thickness of the parallel-laminated mudstone and its close association with conglomerate, as reported by Mueller and Corcoran (2001), reflect a low-energy lagoonal setting. Thicker and more laterally extensive shale interbedded with sandstones, described by Barley (1987) is more consistent with deepwater basinal deposition. Thick and laterally extensive fine-grained lithofacies, including iron formation and chert are closely associated with the turbiditic siltstone/sandstone, indicating a low-energy submarine fan channel or basin plain setting (Eriksson, 1980; Barrett and Fralick, 1989). Synorogenic sequences generally represent subaerial or shallow-water to subaqueous depositional environments along the fringes of volcanic edifices associated with continental margins (Barrett and Fralick, 1989; Eriksson et al., 1994; Mueller and Corcoran, 2001). The overall fining-upward sequences reflect deposition on a transgressive shelf, characterised by wave- and tide-influence, although the coarse claStic detritus interstratified with abundant volcanic flows, breccia and hyaloclastite reported for the c. 3.4 Ga Duffer Formation records only the shoaling-upward stage of edifice construction (DiMarco and Lowe, 1989). The lateral transition from shallow to deep water deposits is often complex, as a result of basin subsidence, sea level rise, and contemporaneous faulting and volcanic edifice construction. Tectonic activity is indicated by faults, volcanic intrusions, unconformable basement-cover relationships, rapid lateral and vertical lithofacies changes, fining- and coarsening-, or coarsening- then fining-upward sequences, and abundant mass- and sheetflow deposits (Fig. 7.3-4). Interbedded volcanic flows and intrusions, formed during extensional tectonism, would have affected the depositional system by increasing topography, providing natural barriers, stabilising slopes, and redirecting alluvial dispersal patterns. An example of an Archaean synorogenic sequence characterised by an unconformable relationship with granitoid basement, a lateral subaerial to subaqueous transition, abundant interstratified volcanic flows and intrusions, fining- and coarsening-upward, and coarsening- then fining-upward sequences, and abundant mass- and sheetflood deposits, is the 2.68-2.69 Ga Raquette Formation (Fig. 7.3-5; Mueller and Corcoran, 2001). The depositional model encompasses a complex interaction of volcanic and sedimentary processes along the subaerial/subaqueous interface of a continental arc. Explosive felsic and effusive mafic volcanism was concomitant with erosion of granitoid basement and deposition of proximal conglomerates and breccias and more distal finer-grained deposits. The lateral transition of the Raquette Formation deposits into the Burwash Formation turbidites, in addition to the lateral interdigitation of the Cameron River volcanic belt supports extensional tectonism.

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Fig. 7.3-4. Characteristics of the Raquette Formation (c. 2.68-2.69 Ga; Slave craton) stratified conglomerate and pebbly sandstone, which were deposited by hyperconcentrated flood flow. Large arrow indicates top and scale, pen, is 15 cm long. (a) Stratified conglomerate and pebbly sandstone with inverse graded beds (Gi), massive to poorly stratified beds (Ms), and low angle scours (Sc). (b) Pebbly sandstone with angular sandstone rip-up clasts (R), and quartz (Q) and plutonic pebbles.

Molasse Sequences (Late-Orogenic) Molasse sequences develop during the terminal stages of orogenic events, during which detritus is shed from high relief basement rocks. The resultant basins, typically rich in coarse, clastic fluvial and alluvial (section 7.8) deposits, are bound by unconformities, marking significant hiati (Mueller and Corcoran, 1998). Basin-margin faults are common and are consistent with pull-apart or strike-slip movement during deposition (Krapez and Barley, 1987; Eriksson et al., 1994; Mueller and Corcoran, 1998). Lithofacies architecture, in addition to clast sizes in conglomerate, reflect tectonic influence on sedimentation (Krapez and Barley, 1987; Corcoran et al., 1998, 1999), and lateral offsets between source rocks and sedimentary deposits support horizontal displacement (Eriksson et al., 1994). Molasse sequences contain lithofacies that change abruptly laterally and vertically and are often arranged in one or more fining-upward or coarsening- then fining-upward sequences

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Fig. 7.3-5. Palaeogeographic reconstruction of the Raquette Formation illustrating the complex interaction of sedimentary and volcanic lithofacies in a subaerial to subaqueous coastal setting affected by extensional tectonism. For more details concerning the description of lithofacies, see Mueller and Corcoran (2001 ).

(Fig. 7.3-6a; Hyde, 1980; Mueller and Corcoran, 1998; Corcoran et al., 1999). Some basins are characterised by interstratified mafic and felsic volcanic flows, and pyroclastic and volcaniclastic deposits (Fig. 7.3-6b; Teal, 1979; Hyde, 1980; Mueller et al., 1994b; Mueller and Corcoran, 1998) (see also chapter 4). Molasse sequences unconformably overlie granitoid, mafic and felsic volcanic, turbiditic, as well as craton-cover successions. In several cases, quartz-feldspar and feldspar porphyry stocks are located along faulted basin margins and the presence of similar porphyry clasts in conglomerates argues for contemporaneous fault movement, intrusion of stocks and sediment deposition (Teal, 1979; Hyde, 1980; Mueller et al., 1991; Mueller and Corcoran, 1998; Corcoran et al., 1999). Depositional settings The general overall fining-upward sequences typical of late-orogenic basins are represented by conglomerate + sandstone + siltstone/mudstone, although small-scale coarsening-upward sequences are found locally. Sequence thickness varies from 0.2 km up to 3 kin, and where volcanic lithofacies form an integral part of the stratigraphy, successions may be up to 5 km thick. The facies sequence and sedimentary structures of the clastic lithofacies are consistent with two types of depositional settings: (1) a high-relief, fault- and unconformity-bound basin with alluvial fans, fan-deltas, braided streams and small ponds or lakes (Hyde, 1980; Krapez and Barley, 1987; Mueller et al., 1991, 1994b; Mueller and Corcoran, 1998; Corcoran et al., 1999), and (2) a high-relief, fault- and unconformity-bound basin with alluvial

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Fig. 7.3-6. Stratigraphic sections from the late-orogenic c. 2.6 Ga Beaulieu Rapids Formation, Slave craton (a) and the c. 2.7 Ga Stormy basin, Wabigoon subprovince (b). The Beaulieu Rapids Formation contains two fining-upward sequences, consistent with tectonic influence on sedimentation. The overall fining-upward sequence of the Stormy basin sedimentary deposits is characterised by interstratified volcanic flows. Diagrams modified from Corcoran et al. (1999) (a) and Mueller and Corcoran (1998) (b). fans, fan-deltas, and a shallow water shoreface with access to the open ocean (Corcoran et al., 1998; Mueller et al., 2002b). General restriction of the conglomerate lithofacies along remnant basin margins, abundant angular clast conglomerate interpreted as talus or rock avalanche deposits (Krapez and Barley, 1987; Corcoran et al., 1998), and the presence of boulders up to 5 m in size indicate deposition in high-relief settings. Sandstones and siltstones characterised by trough cross-beds, planar beds, and minor mudstone, were deposited on the distal reaches of alluvial fans and fan-deltas (Mueller et al., 1991) or on sandy braidplains (Fig. 7.3-7a; Hyde, 1980; Krapez and Barley, 1987; Corcoran et al., 1999). In contrast, abundant tabular beds, composite cross-strata, reactivation surfaces, ripples and abundant mudstone laminae and drapes are more consistent with shallow water deposition affected by waves and/or tides (Figs. 7.3-7b, c; Corcoran et al., 1998; Mueller et al., 2002b) (section 7.5). The mudstone-dominated unit of the siltstone/mudstone lithofacies represents a low-energy offshore (Mueller et al., 2002b), ephemeral pond or lacustrine

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Fig. 7.3-7. Characteristics of the Beaulieu Rapids Formation, Jackson Lake Formation (both c. 2.6 Ga; Slave craton) and Keskarrah Formation (c. 2.6 Ga; Slave craton) lithofacies. Large arrow indicates top. (a) High angle truncating sets of trough cross-beds (St) in the Beaulieu Rapids Formation sandstone. Scale, coin 2 cm in diameter. (b) Planar beds (Pb) and mudstone drapes (Md) in the Keskarrah Formation sandstone. Scale, pencil 15 cm long. (c) Sandstone-argillite lithofacies of the Jackson Lake Formation with sigmoidal tidal bundles (Stb and dashed lines) and mud-draped upper (ub) and lower (lb) bounding surfaces. Note also the trough-shaped cross-beds (Cb). Scale, pen 15 cm long.

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(Krapez and Barley, 1987; Mueller et al., 1991; Corcoran et al., 1998), or foodplain setting (Hyde, 1980; Mueller et al., 1994b). Laminated siltstone and mudstone, ripples, and graded beds are consistent with wave-induced currents and storm activity. Soft sediment deformation structures represent rapid sedimentation as a result of synsedimentary faulting, or disruption during syndepositional volcanic activity. The predominance of fining-upward sequences in late-orogenic successions records the deposition of braidplain sands and lacustrine silts and muds on alluvial fan and fan-delta conglomerates, respectively, or shallow water shoreface sands and lower shoreface to proximal offshore silts and muds on fan-delta conglomerates. The latter depositional sequence is not common in modern settings and may signify distinct Archaean Earth-Moon dynamics (section 5.9). High-relief, unconformity- and fault-bound basins with coalescing fan-deltas prograding directly onto a shallow marine shelf affected by tides (section 7.7) could have been more plausible in Archaean times if the smaller mean Earth-Moon distance created higher tidal regimes (Fig. 7.3-8a; Mueller et al., 2002b). The more common depositional setting for molasse sequences involving alluvialfluvial-lacustrine deposits is well represented by the c. 2.6 Ga Beaulieu Rapids Formation (Fig. 7.3-8b). Two fining-upward sequences are recorded in the narrow fault-bound basin and mark deposition during two distinct cycles related to tectonic influence. At the base of each sequence, basin margin conglomerates represent the tectonic response to basement uplift. Up-section changes into siltstone-sandstone and sandstone lithofacies represent fluvial and local lacustrine deposition. Numerous basins display the interaction between volcanism and sedimentation, whereby volcanic lithofacies acted as high-relief features damming and redirecting fluvial dispersal systems (Fig. 7.3-8c; Hyde, 1980; Mueller et al., 1994b; Mueller and Corcoran, 1998). The sudden input of abundant volcanic debris congested fluvial dispersal patterns and facilitated the runoff of unconfined hyperconcentrated floodflow and debris flow deposits, and aided in forming new lakes. The lack of vegetation during the Archaean would have favoured an initial predominance of unconfined flows rather than channelised flows on high relief slopes.

Summary and Conclusions The three basin-forming events, synrift-passive margin, synorogenic and late-orogenic, encompass the principal geodynamic settings for Archaean basin formation. Each sequence has a distinct facies architecture that can be identified readily in the rock record, despite common fragmental basin preservation. The constant interaction between volcanism (chapter 4) and sedimentation indicates that extensional processes had a significant role in Archaean basin evolution. Archaean terranes on all continents appear to have a similar basin configuration and timing of events so that plate tectonic processes must have been operative (see also sections 2.4-2.7 and 3.6). The preservation of abundant sheetflood deposits in both syn- and late orogenic sequences and the transition from alluvial to shallow water deposits along a high energy coastline may be features specific to the Archaean, as a re-

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Fig. 7.3-8. Examples of late-orogenic basins. (a) The high relief coastal setting of the Jackson Lake Formation with basin margin alluvial deposits and shallow water tidal features. Modified from Mueller et al. (2002). (b) The first depositional cycle of the high-relief alluvial/fluvial Beaulieu Rapids Formation, represented by the large-scale fining-upward sequence of conglomerate and siltstone-sandstone. Modified from Corcoran et al. (1999).

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Fig. 7.3-8 (continued). (c) The high-relief alluvial/fluvial Kirkland Basin (c. 2.7 Ga, Abitibi greenstone belt) with contemporaneous volcanism. Modified from Mueller et al. (1994). sult of the absence of vegetation and inferred higher tidal regimes (sections 5.9 and 7.5), respectively.

7.4.

DISCUSSION OF SELECTED TECHNIQUES AND PROBLEMS IN THE FIELD MAPPING AND INTERPRETATION OF ARCHAEAN CLASTIC METASEDIMENTARY ROCKS OF THE SUPERIOR PROVINCE, CANADA

J.R. DEVANEY Introduction

It is the small details, the raw data collected from outcrops, which constitute much of the foundation upon which broad tectono-stratigraphic, regional geological interpretations are made. Unfortunately, this foundation is largely ignored in most papers, including review articles on Archaean metasediments which summarise various regional stratigraphic case studies and take a necessarily broad view (e.g., Ojakangas, 1985; Eriksson et al., 1994; Mueller and Corcoran, 1998). Any subtleties and uncertainties involving the identification and interpretation of various sedimentary features (e.g., section 7.2), from small scale structures and sequences to large scale units or formations, can result in differences of opinion, controversy, and new interpretations. Because improvements in our ability to assess these foundational details lead to better and more constrained interpretations, and as a partial remedy to the frequent neglect of many of the smaller details noted above, this short essay discusses some field mapping techniques and interpretative concepts. These have proved to be useful or important repeatedly during detailed outcrop mapping and subsequent regional interpretations of parts of The Plecambrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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four Archaean greenstone belts in the Superior Province (Williams et al., 1992) of northern Ontario, Canada: (1) the Beardmore-Geraldton belt of the eastern Wabigoon subprovince (Devaney, 1987; Devaney and Williams, 1989), (2) the Sioux Lookout belt of the western Wabigoon subprovince (Pettijohn, 1934; Walker and Pettijohn, 1971; Devaney, 1999a, 2000a, b), (3) the Melchett Lake belt of the eastern English River subprovince (Devaney, 1999b), and (4) the Birch-Uchi belt of the western Uchi subprovince (Devaney, 1999c, 2001 a, b). Despite this chosen focus on c. 2.7 Ga greenstone belts (e.g., section 2.4), most of what is discussed below could apply to metasediments of any age, and may be of interest to many geologists interested in the difficulties involved in trying to "see through the metamorphic haze" back to the original depositional environments of various sedimentary units.

Descriptive Aspects of Lithofacies in Outcrop Initial questions at the outcrop, the "tuff-wacke" problem, and structural considerations When performing field work on the supracrustal units in a greenstone belt, upon a mapping geologist's first arrival at an outcrop, often three questions must be answered: (1) is it a volcanic or sedimentary rock?; (2) are the layering and any Structures present of "primary" (syndepositional) or "secondary" (hard-rock tectonic deformation) origin?; (3) in which direction are the strata younging? Answers to these most basic questions are sometimes difficult to obtain. Question (1) is troublesome where "volcaniclastic" facies, presumably transitional between proximal pyroclastic and distal sedimentary facies, are found: e.g., ambiguous "tuffwackes". Lateral facies changes in the natural world are commonly gradational, versus the abrupt boundaries in our arbitrary classification schemes. Mapping of a proximal-todistal, volcanic-to-sedimentary facies gradient in along-strike exposures, or comparing the facies and petrographic characteristics within a range or continuum of beds irrespective of their location, will likely be required to place any equivocal "tuff-wackes" in context. Based on various case studies and models (e.g., Cas and Wright, 1987; Devaney, 1999d, 2001a), in an orderly transect from proximal pyroclastic to distal sedimentary facies, it would be expected that: (a) beds become thinner, finer-grained and better sorted (with nontuffaceous slate-mudstone as "background" sedimentation from suspension); (b) mass-flow deposits are replaced by current deposits; (c) evidence of deposition of originally hot pyroclasts diminishes greatly; and (d) the abundance of sedimentary structures and small scale sedimentary sequences increases. These generalisations use relative, not absolute criteria; interesting and problematic exceptions should be expected. Answers to questions (2) and (3) (above) will depend on the geologist's experience and skill in recognising small scale sedimentary and tectonic structures. For example, initially puzzling sandstone laminae which pinch and swell laterally might be suspected to represent a hard-rock boudinage fabric, but careful observation with a hand lens may reveal tiny cross-laminae (ripple foresets defined by grain size changes), and perhaps also regularly laterally spaced ripple shapes and coarser sand grains in the larger ripples, features which would confirm the sedimentary origin of the pinch and swell laminae. Reports such as that

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by Decker (1990), illustrating soft sediment folds (including isoclinal and sheath forms) and truncation surfaces which many geologists might misidentify as hard-rock deformation features, are educational regarding potential exceptions to the "rules". In the case of folded strata, if the folds are small and/or tightly spaced, the answer to question (3) will probably be that the strata young in two or more directions. This is commonly seen in isoclinally to tightly folded, thinly bedded exposures of turbiditic strata: e.g., younging reversals indicated by thin graded beds showing opposite directions of younging, representing separate fold limbs, plus younging directions near any small fold hinges. The younging directions at any hinges displaying axial plane cleavage (i.e., the structural facing direction) can be important to structural analysis.

Way-up (top) indicators The better one's knowledge of sedimentary (and volcanic) structures and small scale sequences, the more top indications can be identified in an outcrop. There are many types of way-up indications other than graded bedding and cross-stratification, so well illustrated texts (e.g., section 7.2), including photo atlases, should be consulted. A great many of the non-sedimentologists working in Archaean greenstone belts assume that all graded bedding is of the fining-upward type. Although beds displaying the fining-upward type, "normal grading", are far more common than coarsening-upward "inversely graded" beds, and normal grading is nearly ubiquitous in thinly bedded turbiditic successions, inversely graded beds can be important locally. Identifying inversely graded beds obviously requires some knowledge about the types of sedimentary processes which form them. A survey of the various depositional settings (alluvial fan, braided fiver, beach, submarine fan channel, etc.) and specific processes relevant to inverse grading is beyond the scope of this short section; the reader is referred to relevant texts, particularly Reineck and Singh (1980, pp. 118-120), Lowe (1982), Clifton (1984), Nemec and Steel (1984), Koster and Steel (1984 and references therein), Bluck (1986) and Cas and Wright (1987). Also, geologists unfamiliar with the intricacies of clastic facies often do not recognise subtly defined bedding surfaces, and may erroneously lump two beds together. For example, a coarse conglomerate bed overlain by a fine conglomerate bed, with the base of the latter being a vaguely defined bedding surface, could be misidentified as a single finingupward bed. It is not uncommon for multiple types of top indications to be present in one bed or one outcrop, even very small ones; e.g., a 2 m thickness of well exposed turbiditic beds, with small scour and load structures giving the same way-up as normally graded interbeds. Such multiple types of top indications increase the local reliability of the resultant structural interpretations. Two-dimensional outcrop surfaces with views of steep foresets offer good approximations of the palaeoflow direction (i.e., estimates of the "apparent palaeocurrent" orientation). Using a real example from an Archaean subaerial stratovolcano succession, in which current tipples are consistently oriented away from a volcanic vent source (Devaney, 2000a): in north-striking, vertically dipping beds, steep foresets and ripple form asymmetry

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show tops to the west and approximate palaeocurrents to the north. Despite the imprecision regarding palaeocurrent azimuths and their tectonic rotation history, such "apparent palaeocurrent" data can be valuable when combined with various other sedimentological features to support a facies model interpretation. Some orderly, small scale sequences of strata can also be used as top indicators. Sequences of beds up to a few metres thick in which the strata thin and become finer-grained upward are present locally in many Archaean fluvial (braided river) deposits. A relatively thin mudstone (slate) bed may cap a fining-upward sequence, or mudstone may be present as intraclasts (rip-up clasts) at or near the base of an overlying, erosively-based bed or sequence. Note that the "identification" of such a sequence is also an interpretation, based on a close comparison with well known undisturbed cross-sections (vertical profiles, real or modelled) of similar appearance (e.g., Miall, 1978), and based on the assumption that because such an orderly sequence is present, the original layering has not been seriously disturbed by tectonic shearing and transposition.

Clast size measurements in deformed conglomerates and related sedimentological patterns Measurements of "maximum clast size", measured as the average size of the 10 largest clasts at a site, can reveal interesting and well preserved sedimentological patterns, even in significantly strained rocks. At a chosen outcrop, the long and short dimensions (length, apparent width) of the 10 largest clasts are measured in the horizontal plane (in the erosionally peneplaned terrain of the Canadian Shield, most outcrops have near flat to low sloping surfaces). Where the schistosity is dipping, the short dimension is not the true width, so the apparent width must be converted to the true thickness (i.e., apparent width multiplied by the sine of the schistosity dip angle). Results from (meta-)conglomerate consisting of oblate clasts must be treated separately from any data derived from exposures with prolate clasts, as they record different structural processes (flattening versus stretching) and, in most cases, represent different structural domains. Having a clast size sampling area of some standard, maximum, or minimum size will reduce a potential source of variance in the results. For polymict conglomerates (those with heterolithic clast compositions), it might be best to record the 10 largest clasts for each of the main clast compositions present. Although laborious to collect, such a subdivided database may help to show different trends among the various clast populations, including subtle patterns. As can be seen in modern environments and undeformed sedimentary formations, and summarised in a typical facies model, sedimentological parameters such as maximum clast size, conglomerate-sandstone or sandstone-mudstone ratio, bed thickness, and the presence and type of sedimentary structures are, with few exceptions, all mutually interrelated. For example, within a braided river conglomerate-sandstone facies assemblage, as the maximum clast size decreases, the local percentage of sandstone typically increases, and conglomerate bed thicknesses decrease (Miall, 1978; Devaney, 1987). Commonly, the effects of deformational stresses on clasts and beds do not have to be removed in order to recognise the types of primary sedimentological patterns and relationships listed above; e.g., thick, strained conglomerate beds will tend to contain large strained clasts, and thin-

7.4. Discussion of Selected Techniques and Ptvblems

629

ner interbeds of similarly strained conglomerate and pebbly sandstone will contain, on average, smaller clasts, reflecting the greater flow power that produced the original undeformed coarse thick beds versus the original finer and thinner interbeds. These and other orderly sedimentological patterns can be recognised in strained rocks at a variety of scales, from the highly local scale (within a single outcrop) to that of a formation-scale unit (e.g., 1 km thick, tens of kilometres along strike; Devaney, 1987; Devaney and Williams, 1989). These methods might also work for coarse pyroclastic facies, as long as the many major differences between volcanic and sedimentary processes are considered (see Cas and Wright, 1987).

Palaeoenvironmental Interpretations The methods of interpretation of the depositional environment of a given Archaean lithofacies assemblage are in most respects identical to those for younger suites of tectonised sedimentary rocks, aside from the obvious secular differences (e.g., sections 7.1, 7.5, 7.6, 7.8 and 7.10) such as the presence of fossils or red beds in younger formations. Although braided river, deep marine, and subaqueous volcanic settings are the ones most commonly interpreted for Archaean clastic deposits (Ojakangas, 1985) (see also section 7.3), familiarity with the full variety of sedimentary and volcanic facies models, well documented in numerous texts, is the ideal to strive for. Of particular use to Archaean stratigraphers would be consideration of a number of observations: (1) braided river conglomerates are not always easy to distinguish from submarine fan channel facies (Hein, 1984); (2) herringbone cross-beds (bimodal-bipolar palaeocurrents) are no longer considered diagnostic of intertidal settings, but tidal bundles with paired mud couplets are diagnostic of tidal processes (Terwindt, 1981) (see also section 7.5); (3) the morphology of wave ripples versus current ripples (de Raaf et al., 1977; Reineck and Singh, 1980) is useful to know in order to identify shallow marine or lacustrine deposits; (4) not all graded beds represent deep-water turbidites; storm deposits include graded beds deposited above wave base (Aigner, 1985); (5) ancient fan-deltas (McPherson et al., 1987) and Gilbertian deltas may have had a steep delta front/slope, allowing deposition of sediment gravity flows in shallow water, with rapid deposition potentially burying and preserving such gravity flow deposits (typically viewed as deep-water facies) above wave base (Devaney, 1991); and (6) in some greenstone belts, the primary volcanic-sedimentary stratigraphy has been overprinted by a "semi-conformable", partly cross-cutting, hydrothermal alteration stratigraphy (Galley, 1993; Devaney, 1999b). Older, simpler facies models, such as those developed in the 1970s, may be more useful for Archaean workers than more recent models which have a greater degree of subdivision and complexity, particularly for generalists unfamiliar with specialised jargon and approaches. For example, older facies models of braided rivers (Miall, 1978) will serve the purpose of most Archaean mappers well, versus more recent fluvial architecture models (Miall, 1985) and other subsequent, more sophisticated approaches which require good to excellent quality exposures (e.g., laterally continuous cliff exposures of flat-lying, unde-

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Chapter 7: Sedimentation Through Time

formed sandstone), very different from the type of database most Archaean workers deal with.

Larger-Scale Aspects: Stratigraphy and Tectonics It has been suggested that Archaean geologists adopt a "sequence stratigraphic" (detailed discussion and Precambrian examples are given in chapter 8) approach (Krapez, 1996), but the practice of sequence stratigraphy (e.g., Galloway, 1989; Van Wagoner et al., 1990; Embry, 1995) relies heavily on the recognition and correlation of unconformities and other key stratal surfaces (e.g., widespread subaerial erosion surface, shoreface ravinement surface, transgressive lag beds, maximum flooding surface). These sorts of data are available in subsurface databases for young rocks (petroleum exploration work) and in well-exposed areas but are mostly unavailable in the small, discontinuous exposures of tectonised rock which typify the Superior Province; i.e., most critical sequence stratigraphic surfaces appear to be hidden rather than exposed. Indeed, in order to understand a formation-scale "package" of strata in the Superior Province, it is more likely that one will have to perform a sophisticated structural analysis of the multiple episodes of tectonic deformation; hence the locally well-known dictum, "you can't do stratigraphy without doing structure". Unlike the case for normal stratigraphic work, it cannot be assumed that "outcrop sections" represent internally undisturbed stratigraphic cross-sections; experienced geologists are not surprised if the layering has been sheared and transposed at a variety of scales (e.g., tectonised sedimentary units can be envisioned as an array of "panels", 1 cm-1 km thick, bounded by shear zones). In the Superior Province, major lithostratigraphic contacts tend to have become deformation zones, typically recessively weathered and predominantly covered by Quaternary glacial overburden, soil and lakes. Given this general absence of exposed contacts, units of similar lithology but different ages may be erroneously lumped together; e.g., one wacke formation may be grouped with a similar-looking, adjacent but much older wacke formation because a mapping geologist might not take the time to look for subtle differences in sand grain composition between the two formations. The presence of a basal conglomerate bed or unit, which will lie directly on a sharp contact (a local erosion surface, or a regional unconformity) and typically contains clasts of a lithology similar or identical to the lithology of what must have been an older rock unit beneath an erosion surface, provides strong evidence of relative ages. However, exposures of basal conglomerates can be frustratingly rare; in both the Beardmore-Geraldton belt (Devaney, 1987; Devaney and Williams, 1989) and the Sioux Lookout belt (Turner and Walker, 1973; Devaney, 2000a), coarse fluvial conglomeratic units extend for tens of kilometres along strike but contain only one small outcrop of a basal conglomerate in each belt. Study of the interplay of tectonics and sedimentation (e.g., Busby and Ingersoll, 1995), usually leading to a classification and interpretation of (palaeo-)basin type for either an individual stratigraphic unit or an entire greenstone belt, has been the subject of much recent research (Eriksson et al., 1994). It is important to note that many (most or all?) Archaean

7.5. P r e c a m b r i a n Tidalites

631

greenstone belts evolved through more than one major basinal stage (e.g., section 2.4), with each stage (or "mega-sequence") having a distinct stratigraphic style reflecting the regional tectonic controls or influences. For example, in the Sioux Lookout belt, an arc succession (stages 2-4) was compressed and thrusted into a foreland basin (stage 5), followed by wrenching and the formation of a smaller strike-slip basin (stage 6), illustrating increasingly smaller-scale structural partitioning of Archaean synorogenic basins (Devaney, 1999a, 2000a). Notably, the youngest sedimentary unit in this belt (stage 6, strike-slip basin-fill) was thought by previous workers to be the oldest sedimentary formation in the belt, illustrating the value of sedimentology in the re-mapping of certain greenstone belts. This and other studies also provide a broader lesson: rather than relying on bold new ideas and models to advance our understanding of Precambrian history, many of us could make better use of old ideas (e.g., know more about basic sedimentary structures) as part of more thorough and competent field mapping, leading in turn to significant advances in our regional interpretations.

7.5.

PRECAMBRIAN TIDALITES: RECOGNITION AND SIGNIFICANCE

K.A. ERIKSSON AND E.L. SIMPSON Introduction

In the absence of fossils, physical sedimentary structures of tidal origin provide the best evidence of marine conditions in Precambrian basins. Recognition of marine conditions is of particular relevance to the ongoing debate concerning the global versus local origin of Neoproterozoic and Palaeoproterozoic carbon isotope excursions (sections 5.3 and 5.8). For example, Melezhik et al. (1999, 2000) have attributed positive carbon isotope excursions in Palaeoproterozoic carbonates to local 13C enrichment in restricted lacustrine basins, whereas other authors (e.g., Hoffman et al., 1998b) argue that such excursions are global in origin and reflect major changes in ocean chemistry in the Neoproterozoic and Palaeoproterozoic. In addition, sedimentary structures produced by tides may be helpful in providing information on Earth-Moon dynamics (section 5.9). Based on our studies of the Archaean Moodies Group (c. 3.3 Ga), Witwatersrand Supergroup (c. 3.0 Ga) and Palaeoproterozoic Waterberg Group (c. 1.8 Ga) in South Africa, the Upper Mount Guide Quartzite (c. 1.8 Ga) in Australia, and the Ortega and Uncompahgre Groups (c. 1.7 Ga) in southwestern U.S.A., we summarise the evidence for tidal influences on their sedimentation, and evaluate the arguments in favour of Precambrian tides. The equivocal nature of some sedimentary structures used as evidence for tides (Table 7.5-1), warrants a re-evaluation of the evidence because of the need to establish a consistent set of criteria for recognising tides in Precambrian units. The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

632

Chapter 7: Sedimentation Through 7~me

Table 7.5-1. "Tidal" sedimentary structures present in stratigraphic units discussed in text Bimodal-bipolar Herringbone T i d a l Compound Rhythmic palaeocurrents cross-beds bedding ripples bedding Uncompahgre Group NO NO YES NO YES (c. 1.7 Ga) Ortega Group YES YES YES NO YES (c. 1.7 Ga) Upper Mount Guide NO NO NO YES NO Group (c. 1.8 Ga) Waterberg Group YES (?) NO YES (?) NO NO (c. 1.8 Ga) Witwatersrand Super- YES YES YES YES YES group (2.7-3.0 Ga) Moodies Group YES YES (?) YES NO YES (c. 3.25 Ga)

Foreset bundles YES NO YES NO YES YES

Recognition Criteria Bimodal-bipolar palaeocurrents A widely used criterion in the above examples is palaeocurrent data (Table 7.5-1). Bimodal-bipolar patterns commonly are accepted as indicating current reversals as for example in the Moodies Group (Eriksson, 1977; Heubeck and Lowe, 1994), the Witwatersrand Supergroup (Eriksson et al., 1981; Karpeta and Els, 1999), the Waterberg Group (Vos and Eriksson, 1977), and the Ortega Group (Soegaard and Eriksson, 1985). Palaeoftow indicators in each of the above examples, with the exception of the Waterberg Group, are cross-beds that provide reliable evidence for reversing currents. Data for the Waterberg Group are based on asymmetrical or combined-flow ripples that probably are the product of waves rather than tides. An absence of bimodal-bipolar palaeocurrent patterns should not exclude a tidal interpretation because Holocene tidal systems often are characterised by the dominance of ebb or flood flow (Dalrymple et al., 1990). Interpreted subtidal sandwave deposits in the upper Mount Guide Group (Eriksson and Simpson, 1990) and the Uncompahgre Group (Harris and Eriksson, 1990) are characterised by unimodal palaeocurrent patterns.

Herringbone cross-bedding Herringbone cross-bedding provides unequivocal evidence of current reversals but, based on observations in modern tidal settings, the likelihood of forming and preserving this structure is low. Ebb and flood currents typically advance and retreat along mutually exclusive pathways or, in those rare instances where the two currents follow similar pathways, one of the currents is characteristically much stronger than the other (Dalrymple et al., 1990). Herringbone cross-bedding may also be misidentified, an example being its probable confusion with overlapping troughs in the Moodies Group (Fig. 7.5-I a; Eriksson, 1977). However, examples of herringbone cross-bedding in the Ortega Group (Soegaard

7.5. Precambrian Tidalites

633

and Eriksson, 1985) and Witwatersrand Supergroup (Karpeta and Els, 1999) are convincing because set boundaries as well as foresets are planar.

Tidal bedding Tidal bedding (flaser, wavy and lenticular) is often cited in support of a tidal origin. Examples are documented from the Moodies Group (Eriksson, 1.977) and particularly from locally in the Waterberg Group (Vos and Eriksson, 1977), but similar structures may develop in response to storm- and fair-weather conditions (e.g., Allen, 1981). However, sandstone components in the Waterberg examples are structured exclusively by wave and combinedflow ripples (Fig. 7.5-lb) and it is likely that the purported tidal bedding was generated in wave-influenced lakes rather than on a tidal flat as previously argued.

Compound ripples Compound ripples such as flat-topped, washed-out, double-crested and ladder-back forms (Fig. 7.5-lc) are common in the Lower Witwatersrand (Eriksson and Fedo, 1994) and Upper Mount Guide (Simpson and Eriksson, 1991) successions. These tipples indicate modification associated with falling water levels and likely reflect retreat of tides. In association with desiccation cracks, these ripples indicate tidal flat emergence. However, compound ripples should not be used as the sole criterion in support of a tidal interpretation because their origin may be complex.

Rhythmic bedding The most compelling sedimentological evidence for ancient tides is based on quantitative data derived from vertically accreted, rhythmically interlaminated sandstone/siltstone and mudstone (tidal rhythmites). Such data are available from the Carboniferous of the U.S.A. (e.g., Miller and Eriksson, 1997; Kvale et al., 1999), the late Neoproterozoic (c. 650 Ma) Elatina Formation and Reynella Siltstone of South Australia (Williams, 1989a, 1994c), and the early Neoproterozoic (c. 900 Ma) Big Cottonwood Formation of Utah (Chan et al., 1994) and have been used to constrain Earth-Moon orbital parameters, including the rate of lunar retreat, back to 900 Ma (e.g., Kvale et al., 1999; Sonett et al., 1996b) (section 5.9). Quantitative data from the pre-900 Ma record are restricted to iron formations from the c. 2.5 Ga Hamersley Group (Trendall, 1973b). The accuracy of the microbanding counts was questioned by Williams (1990) who speculated that the cycles may be annual. Thick-thin pairs of siltstone-shale couplets are common in the Coronation Shale of the West Rand Group, Witwatersrand Supergroup (Hopkins et al., 2000) and in interlaminated siltstones and shales below the Livingstone Reef of the Central Rand Group, Witwatersrand Supergroup (Kuklis et al., 2000; Fig. 7.5-2). These pairs of couplets probably reflect dominant and subordinate semi-diurnal tidal currents, respectively. Less well expressed in "noisy" data sets are thickening and thinning, possible neap-spring-neap cycles (Figs. 7.5-1d and 7.5-2). Rhythmites are also present in the Moodies, Ortega and Uncompahgre Groups (Table 7.5-1 ) but no cyclicity has been identified to date.

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Chapter 7: Sedimentation Through Time

Fig. 7.5-1. (a) Pseudo-herringbone cross-bedding from the Moodies Group, Saddleback Syncline, Barberton Greenstone Belt, South Africa (scale = 14 cm long). (b) Flaser and wavy bedding from the Waterberg Group, South Africa (scale = 5 cm diameter).

7.5. Precambrian Tidalites

635

Fig. 7.5-1 (continued). (c) Ladderback ripples from the Upper Mount Guide Quartzite, Mount Isa, Australia. (d) Interlaminated siltstones and mudstones from the Central Rand Group, Witwatersrand Supergroup, South Africa (rock slab 8 cm long).

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Chapter 7: Sedimentation Through Time

Fig. 7.5-1 (continued). (e) Tidal sand-wave deposit in the Moodies Group, Eureka Syncline, Barberton Greenstone Belt, South Africa showing bundles of foresets separated by mudstone drapes (scale bar -- 5 cm). Note the thickening and thinning of foreset bundles. (f) Cross-bed set showing an increase in thickness of mudstone drapes from left to right corresponding with an increase in thickness of foreset bundle, Moodies Group, Saddleback Syncline, Barberton Greenstone Belt, South Africa (scale in cm).

7.5. Precambrian Tidalites

637

Fig. 7.5-2. Bar chart of siltstone laminae thicknesses in Figure 7.5-1d. Note the common presence of thick-thin pairs of siltstone laminae and the crude thickening and thinning cycles. Central Rand Group, Witwatersrand Supergroup, South Africa.

Foreset bundles A foreset bundle represents the deposit of the dominant portion of a tidal cycle and often is separated from the overlying bundle by a mudstone drape deposited during slack water (Visser, 1980). Less commonly, foresets within bundles display an increase followed by a decrease in dip angle and record acceleration-deceleration of diurnal or semi-diurnal tidal currents. In these instances, bundles of foresets are separated by reactivation surfaces rather than mudstone drapes and are characterised by sigmoidal shapes (cf. Kreisa et al., 1986). Foreset bundles commonly are associated in diurnal, thick-thin pairs related to the semi-diurnal inequality of tidal current velocities. Convincing evidence of tides is provided by systematic thickening and thinning of foreset bundles related to variations in current velocities associated with neap-spring-neap cycles (Visser, 1980). Sigmoidal foreset bundles separated by reactivation surfaces are widely developed in the Upper Mount Guide, Uncompahgre and Witwatersrand successions (Harris and Eriksson, 1990; Simpson and Eriksson, 1991; Karpeta and Els, 1999) but cyclicity has not been quantified. In contrast, neap-spring-neap cycles are well developed in the Moodies Group as described below. The oldest quantitative record of ancient tides. Quantitative evidence for tides has recently been recognised for the first time in the c. 3.25 Ga Moodies Group in the Barberton Greenstone Belt (Eriksson and Simpson, 2000). Tidal signatures in the Moodies Group are preserved as bundles of sandstone foresets separated by mudstone drapes (Fig. 7.5-1 e) in a tidal sand-wave deposit in the lower part of the succession. Detailed measurements

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Chapter 7: Sedimentation Through Time

of foreset-bundle thicknesses at a millimetre scale were made along traverses through the sand-wave deposit and plotted on a bar chart of foreset bundle thickness versus foreset bundle number (Fig. 7.5-3a). Analysis of this plot by analogy with modern tidal processes and records (Nio and Yang, 1980; Tessier et al., 1995) has led to the identification of a hierarchy of diurnal, semi-monthly, and monthly tidal periodicities (Eriksson and Simpson, 2000). Thick-thin pairs of foreset bundles are considered to reflect deposition from semi-diurnal dominant and subordinate flood-tidal currents, respectively. Similar thick-thin diurnal pairs are widely developed in Holocene tidal sediments (de Boer et al., 1999). Cyclic variations in foreset bundle thicknesses (Fig. 7.5-3a) record longer period changes in strength of the dominant semi-diurnal tidal currents consistent with semi-monthly neap-springneap tidal cyclicity. Alternating thicker and thinner neap-spring-neap cycles (Fig. 7.5-3a) are comparable to monthly anomalistic, perigean-apogean tidal signatures reported by Kvale et al. (1999). Fast Fourier transform analysis on the data set reveals strong peaks at 13.11, 9.83 and 2.18. The last 2 peaks are consistent with the interpretation of diurnal and neap-spring cyclicity discussed above, whereas the 13.11 peak is considered to record neap-spring-neap cycles in which both dominant and subordinate semi-diurnal bundles are developed (Eriksson and Simpson, 2000). Fast Fourier transform analysis on the 4-5 month-long data set from which inferred semi diurnal, subordinate-tide foreset bundles had been removed (Fig. 7.5-3b), reveals only one well-developed peak at 9.33 that is interpreted as a strong semi-monthly signature (Eriksson and Simpson, 2000). Close inspection of Figure 7.5-3b reveals that monthly perigean-apogean cycles in the Moodies sand-wave deposit have a maximum number of 20 foreset bundles. These cycles suggest a lunar synodic orbital period of 18-20 days. This is considered to be an estimate of the minimum number of days in the synodic month during the middle Archaean because of the possibility of missing neap-tide foreset bundles, especially within the apogean component of the monthly cycle when tidal current velocities are less than during perigee. Vertical associations of sedimentary structures Although individual structures can be equivocal, repetitive associations of some of the above structures provide stronger support for a tidal interpretation. For example, the upper Mount Guide Group and parts of the Quilalar Formation in the Mount Isa region are composed of stacked, metre-scale parasequences that consist of sigmoidal cross-bedded sandstones, containing acceleration-deceleration cycles, capped by thinly bedded sandstones with a variety of modified ripples and other exposure indicators. Individual parasequences are considered to reflect shoaling from subtidal shelf to intertidal conditions (Eriksson and Simpson, 1990; J.M. Jackson et al., 1990; Simpson and Eriksson, 1991). Vertical associations of sedimentary structures have also been recognised in the Ortega and Uncompahgre Groups (Soegaard and Eriksson, 1985; Harris and Eriksson, 1990) and in the Moodies Group discussed below as a case study. A case study. Upward-fining, fluvial channel deposits in the Moodies Group record evidence for tidal modification. Facies are arranged in 45-140 cm-thick, fining-upward packages in which the proportion of interlaminated sandstone, siltstone and mudstone increases

7.5. Precambrian Tidalites

639

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Foreset Bundle Number Fig. 7.5-3. (a) Bar chart of sandstone bundle thickness versus bundle number. Moodies Group, Eureka Syncline, Barberton Greenstone belt, South Africa. Note the presence of thick-thin pairs of bundles considered to reflect the dominant and subordinate tides of a diurnal system, and the cyclic thinning and thickening of foreset bundles reflecting neap-spring-neap tidal cyclicity. (b) Plot of sandstone bundle thickness versus bundle number using same data set as (a) but with inferred subordinate tide bundles removed. Note the alternation of thicker and thinner neap-spring-neap cycles considered to represent, respectively, perigee and apogee records of an anomalistic tidal system.

640

Chapter 7: Sedimentation Through Time

Idealized Fining-upward Package Interlaminated sandstone, siltstone and mudstone. Thin-thick pairs of sandstonemudstone laminations and systematic thickening and thinning of laminations Some cosets capped by wave and combined-[ flow ripples I

c5 ::-:-:-:.i:!-:-:-:-!:i:!:!:!:i:i:i:!

Cross-bedded sandstone with pebble stringers defining set boundaries. Foresets are tangential, planar or sigmoidal. Reactivation surfaces. Mudstone drapes on foresets more common up-section. Laterally within sets, foresets thicken and thin; thicker mudstone drapes are associated with thinner foresets and vice versa. Some foresets contain cross-laminations directed up the foresets.

N E O | O

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Erosionally based framework-supported conglomerate composed of quartz pebbles and ripped-up clasts of sandstone, siltstone and mudstone

Fig. 7.5-4. Idealised vertical sequence of lithologies and sedimentary structures in upward-fining, tidally modified fluvial deposits from the Moodies Group, Saddleback Syncline, Barberton Greenstone Belt, South Africa.

upwards (Fig. 7.5-4). Basal conglomerates, up to 30 cm thick, are erosional and consist mainly of quartz pebbles and ripped-up clasts of laminated sandstone, siltstone and mudstone. Clast size decreases upwards within conglomerate beds. Overlying cross-bedded sandstone ranges in grain size from very coarse to fine sand. Locally, pebble stringers define set boundaries. Cosets vary from 20 to 210 cm thick. In several instances, laminated sandstone, siltstone and mudstone, and wave and combined-flow ripple bedforms are preserved below coset boundaries. Within sets, foresets are tangential, planar or sigmoidal in shape and, towards the top of upward-fining packages, commonly are draped with mudstone. In general, thin foresets have continuous mudstone drapes whereas thicker

7.5. Precambrian Tidalites

641

foresets have no drapes, discontinuous drapes or are separated by mudstone chips. In bedding plane views, these chips display polygonal desiccation cracks. Reactivation surfaces are present throughout the section. Laterally within sets a systematic thickening and thinning of foresets occurs with a corresponding increase in development of mudstone drapes associated with thinner foresets (Fig. 7.5-1f). Some foresets contain internal ripple crosslaminations directed up the foresets. These ripple cross-laminations show a complex pattern of mudstone drapes. Interlaminated sandstone, siltstone and mudstone intervals cap the upward-fining packages and attain a maximum thickness of 25 cm but commonly are absent at the tops of fining-upward packages as a result of erosion. Vertically within these intervals, thick-thin pairs and systematic thickening and thinning of laminations are developed. Desiccation cracks are ubiquitous. Where laminations are absent, mudstones are black and desiccation-cracked. The vertical sequence of strata within upward-fining packages records the increased influence of tidal currents with time at the expense of fluvial processes. Evidence for the change from fluvial to tidal processes includes an upward decrease in the proportion of conglomerate, the increase in abundance of mudstone drapes on foresets, the presence of cyclic foresets, and the occurrence of interlaminated sandstone, siltstone and mudstone at the top of upward-fining packages. Vertical transition from fluvial to dominantly tidal facies is considered to be related to sea level fluctuations rather than tectonics. Conglomerates reflect channel processes whereas cosets of trough and tabular cross-bedded sandstone and the laminated sandstone, siltstone and mudstone were generated by flows modified by various tidal beats. Cosets of trough and tabular cross-bedded sandstone with or without mudstone drapes reflect lateral accretion of sediment, whereas interlaminated sandstone, siltstone and mudstone records vertical accretion. In both facies associations, mudstone developed during slack water phases whereas sand and/or silt transport took place during the ebb or flood stages. Within both laterally and vertically accreting facies, alternating thin-thick laminations reflect diurnal twice-daily tides. Thinner groupings of foresets and thinner intervals of vertically stacked sandstone-siltstone-mudstone laminations formed during neap tides, whereas thicker groupings of foresets and laminations developed during spring tides. Desiccated mudstone drapes on foresets indicate that bedforms rarely were exposed during some portion of the tidal cycle. Evidence for exposure is best preserved at the top of upward-fining packages. Discussion

Based on the above review, we conclude that quantitative rather than qualitative data, and vertical associations of sedimentary structures provide the best evidence for the existence of tides in the Precambrian. Detailed studies of rhythmically interbedded sandstonesiltstone-mudstone from Archaean and Palaeoproterozoic intervals are warranted with a view to extracting fortnightly and possibly monthly and even annual signals from the record. Studies of foreset bundling patterns in the Moodies Group has revealed a minimum of 18-20 days per month at 3.25 Ga (Eriksson and Simpson, 2000) but the record is incomplete probably because current velocities during neap tides dropped below the

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Chapter 7: Sedimentation Through Time

threshold for sand movement. If complete records of the number of days in a lunar month can be extracted from foreset bundles and/or rhythmites in the Moodies Group and Witwatersrand Supergroup, it will be possible to reconstruct Earth-Moon orbital parameters including number of hours in the day, number of days in the year, Earth-Moon distances, and lunar retreat rates for the Archaean and Palaeoproterozoic (see also section 5.9). In light of the ongoing debate concerning the global versus local origin of Palaeoproterozoic and Neoproterozoic carbon isotope excursions (section 5.3), it is important to establish a marine origin for sedimentary rocks that display or are in close stratigraphic association with those rocks that display the anomalous isotopic characteristics. Palaeoproterozoic rock units of particular interest in this regard are present below and especially above glacial units in the Huronian Supergroup (see also section 5.6) including the Pecors and Gowganda Formations. These units, as well as rhythmically interbedded sandstones, siltstones and mudstones above other Palaeoproterozoic as well as Neoproterozoic glacials (sections 5.7 and 5.8) represent glacial varves but detailed studies may reveal a tidal overprint and thereby establish their marine origin.

7.6.

SEDIMENTARY DYNAMICS OF PRECAMBRIAN AEOLIANITES

E.L. SIMPSON, EE ALKMIM, RK. BOSE, A.J. BUMBY, K.A. ERIKSSON, EG. ERIKSSON, M.A. MARTINS-NETO, L.T. MIDDLETON AND R.H. RAINBIRD Introduction

Identification of diagnostic wind-ripple stratification permits confident recognition of aeolian processes, and separation of aeolian from subaqueous deposits (Hunter, 1977, 1981; Kocurek and Dott, 1981). Additional diagnostic aeolian features include pin-stripe laminations (Fryberger et al., 1988), adhesion structures (Kocurek and Fielder, 1982) and coarse sand to granule ripples (Clemmensen and Abrahamsen, 1983; Fryberger et al., 1992; Clemmensen and Dam, 1993; Bose and Chakraborty, 1994). These diverse criteria have been employed successfully to identify aeolian processes in the Precambrian (e.g., Eriksson and Simpson, 1998; Eriksson et al., 1998b; Bose et al., 1999). During the Precambrian, the absence of terrestrial vegetation should have led to abundant loose surficial materials that would have been modified easily by streams and winds, resulting in widespread braidplain complexes (Cotter, 1978) and erg margin/sand sheet/dune systems (Eriksson and Simpson, 1998). As a consequence, aeolianites should have been more prevalent and should have developed in more diverse climatic and depositional settings. Eriksson and Simpson (1998) and Eriksson et al. (1998b) noted the paucity of Precambrian aeolianites in depositional settings where aeolian processes should have been commonplace, and hypothesised that numerous possible factors including reworking by braided-rivers, erosion during transgression, and non-recognition could have been responsible. Several new examples of Precambrian-age aeolianites have been reported since those papers and are summarised here, along with others. We also interpret controlling The Precambrian Earth: Tempos and Events P2tited by P.G. Eriksson, W. Altcrmann, D.R. Nelson. W.U. Mueller and O. Catuneanu

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factors that led to the generation and preservation of these aeolianites such as tectonics, climate change and groundwater table fluctuations.

Circa 2.6 Ga Aeolian Deposits of the Basal Minas Supergroup, Quadrilrtero Ferrifero, Minas Gerais, Brazil Geological setting The oldest known aeolian sandsheet deposits (c. 2.6 Ga) occur in the Quadrihitero Ferrffero (Iron Quadrangle), in the southeastern Brazilian highlands, at the base of the Minas Supergroup. The QuadriW.tero Ferrffero is a major iron and gold mining district located at the southernmost part of the S~o Francisco craton (Fig.7.6-1). The Precambrian section of the Iron Quadrangle comprises five lithostratigraphic units: an Archaean 3.2-2.9 Ga basement complex (Machado and Carneiro, 1992); the Late Archaean Rio das Velhas Supergroup, a typical greenstone belt sequence and platform cover sequence; surrounding granitoids; the Palaeoproterozoic Minas Supergroup; and the Itacolomi Group (Fig. 7.6-1; Dorr, 1969; Renger et al., 1995; Machado et al., 1996). The Palaeoproterozoic Minas Supergroup is an 8,000 m thick rift to passive-margin to foreland-basin succession (Dorr, 1969; Machado et al., 1996; Alkmim and Marshak, 1998). The oldest units, the Tamandu~. and Caraqa Groups display a vertical transition from alluvial to marine deposits and represent the rift phase of the Minas Basin. The iron-ore bearing Itabira Group and the Piracicaba Group record the thermal subsidence phase, consisting of deltaic to deep-marine deposits. The Sabar~. Group, the youngest unit of the Minas Supergroup (zircon date on tuff layer of 2.125 Ga; Machado et al., 1996), was deposited in a foreland basin (Dorr, 1969; Renger et al., 1995; Reis et al., 2002). The Itacolomi Group, deposited in orogeny-collapse basins (Alkmim and Marshak, 1998), lies unconformably on older units. The age of the aeolian sandstone-bearing Tamandmi Group, the oldest unit of the Minas Supergroup, is not well constrained. A maximum age is provided by the ages of the youngest detrital zircons (2703 -+-48 Ma; Machado et al., 1996). The minimum age of the Tamandu~i Group is 2.42 Ga (Babinski et al., 1995), given by the age of an Itabira Group marble.

Aeolian deposits The Tamandu~i Group in the Caraqa ridge range (c. 3300 m thick) (Fig. 7.6-1) consists of metasandstones, locally interbedded with thin metapelites. Sandstones are quartz arenites and subarkoses. Martins-Neto and Costa (1985) and Rosseto et al. (1987) recognised three lithofacies associations, from the base to the top: fluvial, alluvial plain and aeolian. The fluvial facies association, a few hundred metres thick, consists mainly of sheet-like to lenticular bodies of cross-stratified sandstone interbedded with thin layers and lenses of pelites, indicating deposition in braided channels with minor floodplain deposits. Distally, the fluvial system lost energy and fine-grained deposits, pelites, predominate, characterising the alluvial plain facies association. A c. 2500 m thick, cross-bedded sandy succession overlies the fluvial and alluvial-plain deposits, representing the aeolian facies

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association. Cross-beds are metre- to decametre-thick and trough-shaped, with tangential or high-angle planar forests (Fig. 7.6-2a). Complex patterns of cross-bedding can be observed locally. Interdune deposits consist of decimetre- to metre-thick, sheet-like sandstone bodies mainly composed of facies Sh and subordinately of facies St and Sp (facies codes after Miall, 1978). Controb The upward transition from fluvial to aeolian conditions in the whole occurrence area of the Tamandufi Group, represented by the progradation of thick and extensive aeolian deposits over the fluvial/alluvial system, characterises an upward drying tendency, possibly marking a climate change during the deposition of the group. Sandsheet deposits formed by aeolian dune migration predominate over interdune deposits. The foreset geometries indicate an origin through the migration of barchan and transverse dunes. Complex patterns of cross-bedding suggest local seif dune development. Waterlaid, tractive processes predominate in the interdune deposits. The preponderance of facies Sh (sheetflood deposits) over facies St and Sp (products of, respectively, three-dimensional and two-dimensional subaqueous dune migration) indicates a dominance of upper-flow regime conditions that typifies ephemeral flood events in desert stream systems (e.g., Sneh, 1983). 2.3 Ga Dhalbhum Formation, India Geological setting The 2.3 Ga old Dhalbhum, Singhbhum, and Jharkhand Formations (Saha et al., 1988) developed in an intracratonic rift basin (Mukhopadhyay, 1994; Singh, 1998; Mazumder et al., 2000). Overlying the marine Chaibasa Formation, the Dhalbhum Formation is c. 300 m thick and is exposed in a c. 150 km long belt (Fig. 7.6-1). The aeolian deposit is variable in thickness but averages 30 m. It overlies and intertongues with coarse-grained and poorly sorted fluvial sandstone and is overlain by pillow basalt, locally komatiitic, and even rhyolite. Dark and light coloured and finely laminated tufts are intercalated with aeolian strata.

Opposite: Fig. 7.6-1. Geological maps of aeolianite localities. Central map shows localities of Precambrian cratons on Earth. Inset maps in surrounding figures show positions of Precambrian-age aeolian deposits in these cratons. Brazil--Geological map of the Iron Quadrangle area (modified after Dorr, 1969) and its location on the S~o Francisco craton, southeastern Brazil. India--Geological maps of aeolian deposits in India. (a) Map of India with study areas demarcated. (b) Lateral distribution of Chaibasa, Dhalbhum and volcanic rocks; see (a)for stratigraphic order, and location. (c) Geological map of central India (Vindhyans) along with the stratigraphic context of the Upper Bhander Sandstone. (c) South Africa---Geological map showing the distribution of the Waterberg Group in the main basin (after Callaghan et al., 1991). Canada--Generalised geology of the Dubawnt Supergroup, eastern Baker Lake basin. Arizona, U.S.A.--Map showing the distribution of the Apache Group.

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Aeolian deposits The Dhalbhum Formation aeolian deposit is fine-grained sandstone dominated by windripple laminae that are stacked locally into c. 20 cm-thick sets separated by second-order iron-stained surfaces (Fig. 7.6-2b). These surfaces are without armour and are possibly non-erosional. Translatent strata occur locally in association with the aeolian ripples and consist of alternating grainflow and grainfall laminae (cf. Hunter, 1977, 1981). At various localities are 6 cm-thick sets of crinkled adhesion laminae, and isolated 40 cm-thick lenticular dune cross-bed sets are present (Fig. 7.6-2c). These structural elements generally occur in c. 90 cm-thick cycles with upward transitions from adhesion laminae through translatent/ripple laminae to cross-beds, and are bounded, below and above, by laterally extensive, roughly planar erosion surfaces.

Controls The upward transition from the marine Chaibasa Formation to the terrestrial Dhalbhum Formation sediments is related to plume-generated crustal uplift (Mazumder et al., 2000), even though the absence of anything coarser than sand size and extreme rarity of massive beds precluding rapid sedimentation indicate a low-relief source near sea level. A draa complex failed to develop because of the proximity of the water table to the depositional surface. Pillow basalts, on the uppermost superbounding surface, terminated the aeolian succession and indicate inundation of the aeolian field; therefore preservation of the aeolian deposits appears to be linked to base level rise (cf. Kocurek, 1996). The 90 cm-thick, drying-upward cycles separated by first-order bounding surfaces are possible reflections of shorter subsidence periodicity, although climatic fluctuations cannot be ruled out. Nonerosional iron-stained surfaces separating lamina sets indicate omissions in sedimentation probably caused by wind diversion, perhaps seasonal in nature.

2.0-1.85 Ga Makgabeng Formation, Waterberg Group, South Africa Geological setting Waterberg Group strata in the northern part of the Waterberg basin (Fig. 7.6-1) consist of three formations, which non-conformably overlie the basement gneiss complex. The lowermost Setlaole Formation records deposition in a braided-fluvial environment (Callaghan et al., 1991; Bumby et al., 2001a) whereas the overlying Makgabeng Formation consists mostly of large-scale cross-bedded sandstone deposited in an aeolian setting (Meinster and Tickell, 1975; Callaghan et al., 1991; Simpson et al., 2002). The uppermost Mogalakwena Formation reflects braided river to proximal alluvial-fan deposits. The contact between the Makgabeng and Setlaole Formations appears conformable with fluvial strata within 2 metres of wind-ripple deposits. Recent studies suggest that the Waterberg strata in the northern part of the Waterberg basin were deposited between 2.0 and 1.85 Ga. (Cheney et al., 1990; Bumby et al., 2001b).

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Fig. 7.6-2. Photographs of aeolian deposits. (a) Aeolian dune deposits of the upper Tamandu~i Group, Brazil. (b) Planar omission surfaces (lighter) terminating sets of translatent strata in the Dhalbhum Formation, India. (c) Aeolian dune cross-strata in the Dhalbhum Formation, India.

Aeolian deposits Erg deposits in the Makgabeng Formation consist of large-scale, trough and planar crossbedded sandstone composed of wind-ripple, grainflow and grainfall strata (Simpson et al., 1999; Bumby, 2000; Simpson et al., 2002). Generally, trough cross-beds characterise the

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Fig. 7.6-2 (continued). (d) Second- and third-order surfaces developed in large scale trough cross-beds in the Makgabeng Formation, South Africa. Note the high dip angles of the foresets. (e) Wind-ripple deposits showing pin stripe laminations, Dubawnt Supergroup, Canada. The scale coin is 2.5 cm in diameter. (f) Interdune deposits with raindrop impressions, Dubawnt Supergroup, Canada. Raindrop impressions have a maximum diameter of 1.8 cm.

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lower and upper part of the formation with planar cross-bedded sets more common in the middle. However, near the top of the formation, massive sandstone beds varying from 0.1 to 5.0 m thickness, are associated with a return to the predominance of trough cross-beds. Second-order and, locally, third-order bounding surfaces, as defined by Brookfield (1977), are developed widely (Fig. 7.6-2d). Planar cross-bedded strata represent deposition from straight-crested, transverse aeolian dunes, whereas the large-scale trough cross-bedded strata are likely to represent sedimentation by either sinuous-crested (akl6) or barchanoid dunes (McKee, 1979). Thin drying-upward interdune deposits are interbedded with dunes in the middle part of the formation (Simpson et al., 1999; Eriksson et al., 2000). Above the interdune deposits, playa deposits with evaporite dissolution features can be traced for over 5 km along strike (Simpson et al., under review). Interbedded with aeolianites near the top of the formation are massive and locally plane-bed laminated sandstones with parting lineations, that are indicative of high-velocity, flashy discharge characteristic of ephemeral streams (Miall, 1996; Bumby, 2000). Controls Erg development may have been related to climate amelioration resulting from a reduction in continental freeboard after the c. 2.2-2.0 Ga "southern" supercontinent formation (Eriksson et al., 1999b). As a result of erosion and isostasy, more humid marine weather systems may then have been able to penetrate further into the continental interior changing the sedimentation styles throughout deposition of the Makgabeng Formation. The dryingupward deposits may have developed in flat-lying interdune areas which are more prone to flooding during periods characterised by heavy, periodic precipitation events (Eriksson et al., 2000). Intercalated wind-ripple strata and playa deposits are thought to reflect the local encroachment of aeolian dunes over the margins of the dried-up playa lakes during times of low precipitation or during sandstorms, during which sand flux may fill a playa lake within hours (cf. Wadge et al., 1994). The massive sandstone facies associated with dune deposits are related to slumping of water-saturated sand down the lee face of sand dunes as a result of periodic torrential rainfall (Wizevich, 1997; Loope et al., 1998, 1999; Simpson et al., 2002). The fact that the massive sandstone facies are absent in the lower part of the Makgabeng Formation, and generally more common towards the top, is evidence for long-term climatic change, suggesting that the desert became wetter through time. This overall increase in precipitation towards the top of the Makgabeng Formation is also recorded in the increasingly common transition from inferred interdune deposits to playa lake deposits. Increased precipitation might also explain the transition from transverse to barchan dunes, as they are associated with reduced sediment supply (McKee, 1979). As rainfall rates increased, sand in the source areas was moved by fluvial action, which bypassed the erg rather than being retained as aeolian bedforms. This is reflected in interbedding of high discharge ephemeral fluvial beds near the top of the formation. The widespread evidence for an increasingly wet palaeoclimate recorded in the upper strata of the Makgabeng Formation also provides evidence for the cessation of aeolian conditions.

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Circa 1.85-1.72 Ga Dubawnt Supergroup, Western Churchill Province, Nunavut, Canada Geological setting The Dubawnt Supergroup is a tripartite Palaeoproterozoic succession of continental siliciclastic and volcanic rocks up to 15 km thick (Fig. 7.6-1; Gall et al., 1992; Rainbird et al., 2001a). The Baker Lake Group and unconformably overlying Wharton Group were deposited in elongate rift and strike-slip basins comprising the Baker Lake basin (see also section 3.5). The younger Barrensland Group is confined mainly to the Thelon basin, interpreted as an intracratonic thermal sag. The Baker Lake Group includes progradational to retrogradational alluvial red bed sequences defined by intervening ephemeral lake and erg deposits. Radiometric dating provides a maximum age of 1.85 Ga (Rainbird et al., in review). The Wharton Group is comprised of widespread basal sandstone, including erg deposits, overlain by 1.76-1.75 Ga porphyritic rhyolitic lava flows and pyroclastic and epiclastic sedimentary rocks. The Barrensland Group includes lower alluvial conglomerates grading upward into fluvial, aeolian and ultimately marine sandstones. Aeolian deposits Aeolianites occur throughout the Dubawnt Supergroup and typically are associated with alluvial fan deposits that formed on the margins of extensional fault-bounded basins (Rainbird et al., 2001 a). In the Baker Lake Group, aeolianites occur in the centres of basins and are closely associated with ephemeral lake deposits that were supplied by transverse and axial river systems. Two types of aeolian deposits are exposed in the lower Baker Lake Group strata from eastern Baker Lake basin: thin sandsheet and thicker erg deposits. Aeolian sandsheet deposits, exposed at Thirty Mile Lake, are a component of ephemeral lake and fluvial facies associations within a series of 100-500 m-thick, alluvial fan-fluviallacustrine cycles, interpreted as third-order sequences (Hadlari and Rainbird, 2000). Sandsheet units consist of small-scale, low-angle, wind-ripple deposits (Fig. 7.6-2e), interbedded with wavy-lenticular bedded sandstone and siltstone and parallel-laminated siltstone and mudstone with desiccation cracks. These sandsheets formed between alluvial flood events associated with the ephemeral lake. Similar strata also are preserved in braided stream deposits on the tops of coarse, erosional-based, fining-upward cycles. Sandsheets developed from reworking of fluvial channel macroforms (terminology of Ashley, 1990) and thin overbank deposits between flood events. At a section along the Kazan River, erg deposits up to several hundred metres thick comprise thick (up to 100 m) packages of stacked and overlapping large-scale cross-beds, that are interbedded with braided river through delta to ephemeral lake deposits (Fig. 7.6-3). The erg section commences with c. 50 m of festoon cross-beds, up to 6 m thick and 150 m wide. Cutting down several metres into the top of the erg complex is a discontinuous, c. 100 m-long, massive sandstone lens that represents a slump produced by flooding and localised liquefaction of the erg complex. The massive sandstone passes upward into plane-bedded to ripple cross-laminated to small-scale trough cross-bedded sandstone. The upper bounding surface is flat, erosive and extends laterally for at least several hundred metres. The surface is overlain by a thin

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mudclast breccia and about 2.5 m of plane to wavy-bedded sandstone. Collectively, the strata above the large-scale festoon cross-bedded sandstone units of the erg complex are interpreted as shallow-water deposits formed during flooding of low-lying interdune areas. Above the interdune deposits is an interval characterised by alternating dune and interdune deposits. This, in turn, passes upward into another thick large-scale cross-bedded sandstone, interpreted as a dune complex. The erg complexes are bounded by a variety of associated facies interpreted as ephemeral lake, lacustrine delta and braided fluvial environments (Fig. 7.6-3; Rainbird et al., 1999; Hadlari, in preparation). Within an erg complex at the southern end of Christopher Island, bounding surfaces of large-scale cross-beds are overlain by thin intervals of ripple-laminated sandstone and siltstone. The ripples are mainly symmetrical but rare current ripples are observed. Elsewhere on Christopher Island similar ripple-bedded interlayers are underlain by thin pebble lags, suggesting wind deflation followed by transgression (flooding) and ponding in low-lying interdune areas (Hadlari, in preparation). Other wellpreserved features of the interdune deposits are adhesion structures and raindrop imprints (Fig. 7.6-2f). Overlying and bounding the erg complex is a > 50 m-thick interval composed of 10-60 cm thick beds of sandstone, overlain by multi-generational, desiccated siltstone and mudstone. These layers are interpreted as distal flood deposits in ephemeral lakes; coarse-grained bases reflect tractional deposition followed by suspension deposits that were subsequently desiccated. Controls

Aeolian sandsheet deposits described from the Thirty Mile Lake area are located at the tops of retrogradational, tectonically controlled, basin-filling cycles interpreted as third-order sequences (Hadlari and Rainbird, 2000). As such, aeolianites were preserved during periods of maximum accommodation and minimum coarse sediment influx from uplifts along inferred basin-bounding faults. Lack of exposure prevents correlation of these sequences eastward; therefore the Thirty Mile Lake sandsheets may represent the distal edges of the thicker erg deposits of the Kazan River and Christopher Island areas (Fig. 7.6-3). Preservation of thicker aeolian and ephemeral lake deposits is due to greater accommodation, partly as a consequence of greater subsidence in the east, as established by fluvial palaeocurrent analysis (Rainbird et al., in press). Alluvial fan deposits also are thinner in eastern Baker Lake basin relative to the Thirty Mile Lake area, suggesting greater accommodation related to reduced alluvial sediment influx. A cyclic alternation of erg complex-dominated (dry) and fluvial-dominated wet intervals is suggested in the sections from eastern Baker Lake basin (cf. Hadlari and Rainbird, 2000; Rainbird et al., 200 l a). Limited exposure precludes a sequence stratigraphic interpretation, although the laterally extensive bounding surface overlain by mud-clast breccia in the Kazan River section could be interpreted as a sequence boundary, representing a significant period of exposure, and succeeded by breccia, erosion and bypass (Kocurek and Havholm, 1993). Altemation of dry and wet aeolian systems may have been tectonically controlled, but was more likely the product of another allocyclic mechanism such as climate. Climate fluctuations are ascribed to numerous mechanisms and controls operat-

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Fig. 7.6-3. Representative stratigraphic sections of typical erg and erg-margin deposits of the Baker Lake Group: (a) Kazan River section; (b) Christopher Island section (see Fig. 7.6-1 for locations).

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ing on a variety of time-scales. Global scale orbital forcing has been invoked for similar scale cyclic alternations of dune complex and interdune deposits from Quaternary aeolian deposits in the Sahara region (Kocurek, 1998; Swezey et al., 1999).

1.26-1.10 Ga Troy Quartzite, U.S.A. Geological setting Late Mesoproterozoic and Neoproterozoic basins occur throughout western North America. These basins comprise isolated intracratonic rifts and aulacogens including the Belt, Uinta, Grand Canyon (Unkar and Chuar Groups), Apache-Troy, and Pahrump basins (Stewart, 1972). Rare palaeoenvironmental interpretations indicate deposition largely in shelf and nearshore marine settings. The Pahrump in southeastern California and the Apache-Troy in central and southern Arizona contain ample evidence of deposition in terrestrial settings (Middleton and Trujillo, 1984; Fedo and Cooper, 2001). The ApacheTroy basin in particular contains the thickest and most complete record of erg margin and erg sedimentation of any Mesoproterozoic deposits in western North America (Fig. 7.6-1). The Mesoproterozoic Apache Group and overlying Troy Quartzite crop out discontinuously throughout central and parts of southern Arizona (Fig. 7.6-1). Maximum thickness for the Apache-Troy is approximately 850 m (Wrucke, 1989) but, due to numerous unconformities, thickness is variable. Studies by Middleton and Trujillo (1984) have documented alluvial fan and braidplain deposits at the base of the Apache Group. Marine and subordinate non-marine units occur throughout the succession. The overlying Troy Quartzite rests with pronounced unconformity on the Apache and contains ample evidence, particularly at its base, of erg margin and erg deposition (Weiss and Middleton, 1986). These comprise the arkose member (informal) of the Troy Quartzite (Shride, 1967). The age of the arkose member is constrained by dating of widespread diabase intrusions and of detrital zircons. Diabase sills and dykes intruding the Apache-Troy package yield a U-Pb zircon age of about 1.1 Ga (Silver, 1978). Detrital zircon geochronology establishes a maximum age of 1.26 Ga (Stewart et al., 2001).

Aeolian deposits Extradunal deposits.

The arkose member, up to 120 m thick, was deposited on an eroded surface of basalt and dolostone. The basal arkose member represents extradunal deposits which are composed of eroded basement debris as well as medium- to coarse-grained sandstone derived from uplifted nearby granites. Ventifacts occur at several localities. Finingupward sequences are common but there is little evidence of cyclicity. This association reflects deposition on a broad braided plain that experienced episodes of sheetfloods and was dissected by broad, shallow braided channels. The broad, shallow channels and sheetlike geometries of the sandstones reflect the unconfined nature of these high energy flows, and migration of in-channel dunes. These extradunal deposits typically are overlain by and interbedded with low-angle cross-stratified sandstones that represent wind-ripple migration during onset of erg-margin sedimentation.

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Sandsheet deposits.

Horizontal and low-angle laminae commonly occur at the base of, and intercalated with, erg deposits. Wind-ripple strata are common. Thin beds of rippled sandstone and siltstone locally contain adhesion ripples and desiccated mudstone. Medium-grained sandstone beds, up to 1 m thick, developed as broad sandsheets peripheral to the main erg. Ephemeral flows from nearby highlands transported coarse sand into the basin, that was modified into coarse wind ripples on the sandsheets. Similar deposits are reported from modern dune fields (Kocurek and Nielson, 1986) as well as from ancient Precambrian deposits (Tirsgaard and Oxnevad, 1998).

Erg deposits. Dune deposits make up more than 70% of all sections of the arkose erg unit. Trough cross-strata occur as 2-4 m solitary sets or as cosets up to 8 m thick (Fig. 7.6-4). Internal stratification is dominated by wind-ripple laminae with subordinate grain-flow deposits. The thick planar-tabular sets and complexly cross-stratified sets resulted from periodic migration of simple, transverse and barchanoid dunes, and draa complexes. The latter exhibit large, erosional discordances reflecting periodic cessation of bedform movement and/or wind erosion by lee-side eddies, produced by winds across the slip faces. These second-order bounding surfaces also separate sets of wind-ripple laminae. Large scale sets are bounded by first-order bounding surfaces, likely the result of bedform climb (Rubin and Hunter, 1982), and are overlain by minor plane-bedded sandstone that accumulated in dry, interdune areas. The scale of the cross-strata and the angle of foreset dip decrease upward. In the upper portions of the arkose member, 2-4 m thick trough cross-strata and up to 3 m thick beds of low-angle strata are common. The trough sets are broad and represent large scale deflation areas filled by migrating wind ripples on sandsheets that formed along the trailing margins of the erg. Controls The arkose member erg formed in an intramontane setting close to nearby highlands. The presence of ventifacts and absence of wet interdune deposits suggest that the climate was arid. If the climate were more humid, considerable reworking of aeolian deposits would be expected due to sheetfloods and erosion by wide, braided channel systems of the nonstabilised, surficial detritus. The thick dune deposits likely accumulated in a confined basin that was experiencing moderate to high rates of subsidence. Such conditions are common in intramontane settings. The setting likewise resulted in complex wind regimes. This is attested to by the complexity of the internal stratification and numerous truncation surfaces. Such conditions would be expected in a basin where wind regimes were varied in intensity and direction.

0.6 Ga Upper Bhander Sandstone, India Geological setting An erg in the 600 Ma Upper Bhander Sandstone (UBS), at the top of the Vindhyan Supergroup, central India, formed in an intracratonic sag basin (Fig. 7.6-1; Bose et al., 1999, 2001 ; Chakraborty and Chakraborty, 2001; Sarkar et al., 2002; Ray et al., 2002). The Bhander Sandstone, the topmost member of the 600 Ma Bhander Formation (Ray et al., 2002),

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Fig. 7.6-4. Outcrop photograph of large scale foresets in the arkose member of the Apache/Troy Groups. Note man for scale.

is underlain by the coeval Sirbu Shale, and is terminated by an unconformity. Bose et al. (1999) report that the basal contact with the Sirbu Shale, deposited in a shelf setting, rises stratigraphically seaward because of progradation punctuated by intermittent marine flooding. Consequently, the succession is divided into a number of parasequences bounded below and above by marine flooding or superbounding surfaces.

Aeolian deposits The overall prograding UBS is terrestrial, except for the seaward fringe where marine supralittoral storm bed packages occur at a number of stratigraphic levels. The maximum flooding surfaces and their lateral equivalents, characterised by erosion, thick iron encrustation and pitting, divide the UBS into a number of parasequences (Bose et al., 1999, 2001). These erosion surfaces dip gently seaward with a convex upward geometry. The superbounding surface, enclosing the parasequences, on the seaward side passes landward into a first-order aeolian surface, separating an underlying draa deposit from the overlying sandsheet. Further landwards, this surface passes into second-order bounding surfaces that

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intervene between draa deposits. Bose et al. (2001) noted numerous low angle and striated slide planes, with a strike of northeast-southwest, in close association with these surfaces. Individual parasequences bounded below and above by superbounding surfaces are upward-drying and are stacked vertically with slight seaward offset. Within parasequences, sandsheet deposits pass upward into dunes and draa deposits. Adhesion lamina sets, with individual thickness of 8-10 cm, dominate the parasequence bases; thickness and frequency decrease eastward. Translatent strata encasing isolated dune cross-sets (draa) dominate the mid-level and multiple dune cross-bed sets dominate the top of parasequences, locally composed of longitudinal dunes (Bose et al., 1999). In a (palaeo-)landward direction, parasequences thicken and dune cross-strata and draa deposits increase in abundance. The longitudinal dune deposits have a maximum set thickness of 1.2 m with cosets up to 3.2 m thick. Individual cross-sets of transverse dunes are only 1.4 m thick. Interbedded with aeolian deposits are thin fluvial and lacustrine beds. Lacustrine deposits are characterised by mudstone-siltstone/fine sandstone interbeds. Fluvial deposits are dominantly sandy, and are less well sorted than either the aeolian or lacustrine sandstones. Controls The association of slide planes with parasequence boundaries identifies basin subsidence as the cause for parasequence terminations. The subsidence-related marine flooding apparently controlled the thickness of individual parasequences and preserved thickness of the draas. Water table rise restricted deflation of sand, and dune and draa growth. Occurrence of longitudinal dunes at the top of some parasequences indicates restriction in sand supply (Tsoar, 1982, 1983; Wasson and Hyde, 1983; Rubin and Hunter, 1987; Rubin and Ikeda, 1990). At the top of a parasequence, water table rise was probably slow initially, but a subsequent rapid rise terminated the parasequences. A slow rise in the water table allowed the next drying-up parasequence to develop, explaining the appearance of terrestrial subaqueous deposits at the base and top of the parasequences. In this progradational succession, the depositional surface had a low gradient. High relief was not generated even at the peak of tectonism. The uniformity of facies assemblages between parasequences indicates that no significant climate change took place. Therefore, relative sea level fluctuation was the principal factor controlling the water table, with the water table remaining close to but not far below the depositional substratum. Chakraborty and Chakraborty (2001) attribute the relative rarity of palaeodunes in this succession to this high water table. Discussion

These above examples of Precambrian aeolian deposits illustrate the diverse ages and processes controlling the preservation of these aeolianites. The oldest known sandsheet deposits are found in the Minas Supergroup, with the only older evidence of aeolian processes preserved as ventifacts in the c. 2.9 Ga Witwatersrand basin (Minter, 1976). Sandsheets are the predominant record of aeolian deposit until the appearance of erg/draa sedimentation in the Palaeoproterozoic to Mesoproterozoic (see Eriksson and Simpson, 1998; Eriksson

7. 7. Early P r e c a m b r i a n Epeiric Seas

657

et al., 1998b). Ground water table control on preservation has been documented from the Dhalbhum Formation and the Upper Bhander Sandstone. Ground water fluctuations controlled both the internal packaging of sedimentary structures and the overall thickness of preserved parasequences. Climate change during the lifespan of an erg is well illustrated in the Makgabeng Formation. The appearance of wet interdune deposits, playa and massive sandstones records an increase in precipitation in the Makgabeng dune field. Thick erg deposits of the Dubawnt Supergroup and Troy Quartzite illustrate the control of tectonics on stratigraphic position and sedimentation style of aeolianites.

7.7.

EARLY PRECAMBRIAN EPEIRIC SEAS

EG. ERIKSSON, A.J. BUMBY AND E MOSTERT As sedimentary processes in shelf and epeiric seas are not fully understood, their discrimination remains problematic (Brenner, 1980). Epeiric seas encompass epeiric embayments and epeiric seaways. Seaways submerged large parts of cratons and were characterised by shelf-like palaeoenvironments, shelf-breaks and strongly directional oceantype palaeocurrents, in contrast to the smaller and shallower embayments (Friedman et al., 1992). Seaways thus closely mirror shallow oceans and open ocean shelves (Bouma et al., 1982) and embayments are epeiric seas s e n s u s t r i c t o . True modern counterparts to both types are absent (Galloway and Hobday, 1983). Shaw (1964) and Irwin (1965) proposed the classic conceptual (Phanerozoic) epeiric sea model, with low oceanwards gradients (c. 1:50 000) dissipating offshore wave energy seaward of the coastline in an open-sea "X-zone". This zone was separated from an equally broad, shallow, landward "Z-zone" by a narrower, high energy wave and tidal "Y-zone" (Fig. 7.7-1). Epeiric sea depths are thought to have varied from c. 30 m (and less) to 100 m, and under such conditions storm waves, wind-driven surface currents and water level changes would have resulted in a well-mixed "Z-zone" (Friedman et al., 1992). Variable salinities characterised the "Z-zone" (Hallam, 1981). Open-ocean tides approaching the wide, shallow epeiric platforms increased tidal range rather than the opposite (Pratt and James, 1986), as did seafloor topography and tidal resonance (Hallam, 1981). The c. 3074-2714 Ma greater Witwatersrand basin, Kaapvaal craton, extended at least 400 km inland and for 600 km along the developing craton margin (Beukes and Cairncross, 1991; Robb and Meyer, 1995; Eriksson et al., 1998b). This retroarc foreland system had lower, transgressive, underfilled and upper, regressive, overfilled basin-fill successions (Catuneanu, 2001). Preserved facies support tidal and wave-storm wave deposition within this sea, with evidence for predominant tidal action along the (inland) coastline and wave action, shore-oblique and geostrophic currents along the (seaward) cratonic marginal regions (Eriksson et al., 1981 ; Beukes and Cairncross, 1991; Stanistreet and McCarthy, 1991; Beukes, 1996). Fluvial braidplain deposits are intimately associated with the epeiric marine sediments (e.g., Els, 1998). The microplate structure of the still developing Kaapvaal craton likely played a role in the evolution of this inferred epeiric embayment. The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

Chapter 7: Sedimentation Through Time

658

I

xzo.E H.re~

I YZO"E I ---..-

Tenso, m

ZZO"E 1

"."

"

Om

200m

i ~Wavebase ......... ue!r,[a, sea,men[s n'om nigh energy Y zone

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Tides =,. . . . ,,, ~,,uu.... sediments

I ~I I I I

Some ~ Tides ~ None Chemical sediments (Dolomite~

~ Evaporites)

Fig. 7.7-1. Classic Shaw (1964)-Irwin (1965)epeiric sea model.

Periodically elevated global sea level related to high continental crustal growth rates (section 2.8), combined with high denudation rates (sections 5.10 and 5.11) and concomitant lowered freeboard during the Palaeoproterozoic promoted epeiric transgression onto growing cratons (Eriksson et al., 2002b). Two major embayments are inferred for the c. 2.4-2.1 Ga Pretoria Group, Transvaal Supergroup (Kaapvaal). For the Timeball Hill Formation sea, preserved over 500 x 300 km, Eriksson and Reczko (1998) suggest a relatively deep-water embayment (Fig. 7.7-2), formed by tectonic subsidence during global glaciation (section 5.6) at c. 2.4-2.2 Ga. The preserved Silverton Formation embayment, of similar dimensions, has a basal arenaceous facies association ascribed to braid-deltaic and turbidity current origin (Eriksson et al., 2002a). Overlying argillaceous facies are thicker and are ascribed to sub-storm wave base pelagic sedimentation under transgressive to highstand conditions, formed in transitional and offshore mud belts. Storm waves locally formed graded siltstone to fine sandstone beds. Muds were derived from fine fluvial sediment bypassing a high energy coastal sand belt, preserved as the Magaliesberg Formation (regressive systems tract) which overlies the Silverton Formation (Eriksson et al., 2002a) (Fig. 7.7-3). Compared to the Irwin-Shaw model (Fig. 7.7-1), the "Silverton sea" had a much enlarged "Y-zone", characterised by ephemeral braid-delta systems debouching into high energy peritidal flats marking the inland margin of the embayment (Fig. 7.7-3). Braid-delta and tidal channel dynamics were analogous and are best distinguished from palaeocurrent data (Eriksson et al., 1995). Ripples abound on the upper surfaces of these sandsheets and reflect 90% wave-formed structures, with 4% due to current action and 6% to wind (n - 194). Applying Tanner's (1971) techniques to estimate wave height and water depths from wave ripple forms, suggests waves of 1.5-23.5 cm (average 7 cm), and depths of 7 cm-1.52 m (average 31 cm). These sub-fair weather wave base water depths (cf. Aspler et al., 1994) negate a subtidal sandsheet interpretation, and suggest that tidal sandsheet

7. 7. Early Precambrian Epeiric Seas

659

Fig. 7.7-3. Shallower water, low gradient model proposed for the Palaeoproterozoic Silverton Formation epeiric sea by Eriksson et al. (2002).

thicknesses (c. 0.5-5 m) approximate tidal range (sensu Klein, 1971). Meso-macrotidal conditions can thus be inferred for the Silverton-Magaliesberg coastline. Maximum preserved channel-fill thicknesses support tidal downcutting ~< 70 cm and fluvial erosion ~< 1m (Eriksson et al., 1995). In the lower Transvaal Supergroup, the c. 2642-2432 Ma Ghaap-Chuniespoort succession contains up to 2500 m of dolomitic carbonate rocks, predominantly of organosedimentary or stromatolitic origin (sections 6.4 and 6.5), and 700 m of succeeding banded

Chapter 7: Sedimentation Through Time

660

iron-formation (BIF) (Altermann and Siegfried, 1997; Eriksson and Altermann, 1998). Thermal subsidence formed an intracratonic sag basin in which this chemical succession was laid down under highstand conditions (Catuneanu and Eriksson, 1999). Water depths are estimated at 40-80 m during carbonate sedimentation (Klein et al., 1987; Eriksson and Altermann, 1998) and at > 100 m during BIF deposition (Klein and Beukes, 1989); an area of c. 600,000 km 2 (Beukes, 1987) suggests a seaway rather than an embayment. Shallow water depths appear to have been pertinent throughout carbonate deposition, and intertidal to supratidal facies have been identified (Altermann and Herbig, 1991). Due to a major transgression at c. 2550 Ma, early carbonates in the southwest of the basin were drowned, and a large carbonate platform developed over much of the Kaapvaal craton lying to the north and northeast (Altermann and Wotherspoon, 1995). Deposits from a younger terrigenous-carbonate epeiric sea are preserved in the Belt basin, western North America. Pratt (2001) describes syndepositional tectonics and palaeoenvironmental conditions during sedimentation of the carbonate-dominated c. 1.45 Ga Helena Formation. Lime mud deposited as low energy tempestites accumulated at depths of about 50 m and this epeiric basin was characterised by tsunamis, a thermocline, a shallow aragonite compensation depth in warm water, and by temporary salinity stratification (Pratt, 2001).

7.8.

PRECAMBRIAN RIVERS

D.G.E LONG Introduction

In order to understand the behaviour of fluvial systems before the advent of rootedvegetation it is critical to realise that many of the processes that influenced pre-Devonian fluvial systems were significantly different from the present. The greatest difference would have been that the lack of sediment binding, baffling, and trapping by plant roots would have promoted a tendency for flashy surface run-off, lower bank stability, and faster rates of channel migration than in present-day vegetated areas (Schumm, 1968; Cotter, 1978; Long, 1978; Fuller, 1985; Els, 1990). Although anaerobic microbial communities may have been important in Proterozoic soils and groundwater (Martini, 1994; Horodyski and Knauth, 1994; Ohmoto, 1996b) (see also sections 7.9 and 7.10) they would have had little effect on sediment binding. This is supported by the near absence of algal- and microbial-bound sand-chips (Pfltiger and Gresse, 1996; Schieber, 1998) in Precambrian fluvial deposits. In modern systems sediment supply and characteristics are directly influenced by climate (Blum and T6rnqvist, 2000). In the Archaean, climate zones may have been influenced by the faster rate of the Earth's rotation and differences in the tilt of the rotational axis (section 5.9). Despite lower solar luminosity, Archaean climates were probably predominantly warm (e.g., Eriksson et al., 1998b). It has been suggested that enhanced levels of greenhouse gasses promoted aggressive weathering of labile material (Donaldson and de Kemp, The Precambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

7.8. PrecambrianRivers

661

1998; Corcoran et al., 1998) (sections 5.10 and 5.11). If these were not dissolved, or removed by the wind, a greater availability of fines should have promoted mass flow and hyperconcentrated flow processes in Archaean fan and river systems (section 7.3). By the latest Archaean and early Palaeoproterozoic, the presence of glacial deposits (sections 5.6 and 5.7) indicates development of a broader range of climatic zones. Warmhumid climates may have prevailed during much of the early Palaeoproterozoic (at least to 2.3 Ga), as there is little preserved evidence for arid and hyper-arid conditions (Eriksson et al., 1998b). Acid rain effects should have declined in this period in tandem with decreases in greenhouse gasses, especially after the onset of fully oxygenic conditions after c. 2.2 Ga (Kasting, 2001) (see, however, section 5.2). The progressive colonisation of terrestrial environments by rooted plants in late Silurian to mid-Devonian times would have affected significantly microclimate, by modifying albedo and moisture retention, leading to a greater role of organic acids in decomposition of labile components. Pre- Vegetation River Systems

The depositional products of most pre-vegetation fluvial systems appear to be fairly similar to those of modem braided and ephemeral systems in dry-land climates, although individual river systems may have developed in a broader range of climatic settings. Based on the sheet-like geometry of many pre-vegetational deposits, it is clear that on unconfined braid-plain systems, flood channels were significantly wider, with width to depth ratios from 200:1 to more than 1000:1 (Fuller, 1985; Els, 1990; Rainbird, 1992). Interpretation of these fluvial systems has relied largely on direct comparison with idealised models based on small Holocene river systems, predominantly from valley-confined systems in humid-temperate climates (Cotter, 1978; Friend, 1978; Long, 1978; Fuller, 1985). This has resulted in the identification of numerous bed-load dominated, sheet-like, braided alluvial deposits similar to modern sandy (Platte- and South Saskatchewan-type) and graveldominated (Scott- and Donjek-type) rivers and fans, using models developed by Miall (1977, 1978, 1996). Descriptions of pre-vegetational high-sinuosity meandering systems are rare (Sweet, 1988). Gravel-dominated systems

Many studies of pre-vegetation gravel-dominated systems have suggested deposition on alluvial fans without collaborative evidence of radiating dispersal patterns, down-stream changes in maximum grain size or three-dimensional geometry. Systems containing or dominated by matrix-supported conglomerates are typically interpreted as Trollheim-type fan deposits with in-channel and possibly lobate debris-flows. As the availability of fines encourages the production of mass-flow and hyperconcentrated flow deposits, it should be expected that enhanced weathering of labile components by acidic rainfall in the Archaean should have led to a greater climatic range of this facies, and may even have allowed production of debris-flows in non-fan fluvial systems (Buck and Minter, 1985). Supporting

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evidence of this comes from observations by Pfltiger and Seilacher (1991), who describe unusual contact framework conglomerates, containing stacks of imbricated, well-rounded boulders and cobbles, with up-stream directed false bedding produced by plastering of clasts on to the back of gravel bars by hyperconcentrated flows during flash-flood events in a valley-confined setting. Hyperconcentrated flows are also thought to have produced channelised and sheet-like units in which matrix-supported conglomerates and overlying massive sandstones were deposited by sheet-floods, which degraded into normal flows to produce an upper layer of trough cross-stratified sandstone (Dillard et al., 1999). Wandering and meandering gravel systems (Miall, 1996) have yet to be identified in prevegetation fluvial systems, although a number of authors have identified shallow (Scott, type) to deep (Donjek-type) mixed sandy-gravelly systems (Ethridge et al., 1984; Long, 1987, 2001a, b; Eriksson and Simpson, 1993) based largely on the preserved thickness of depositional cycles. Sand-dominated systems Pre-vegetation sandy systems typically contain stacked sequences of bedforms, which form small-scale depositional sequences with high lateral continuity ("sheet-braided" of Cotter, 1978). Clear evidence of channel elements, or even channel margins, are very rare. Sequences dominated by planar cross-stratification tend to be interpreted as Plattetype river deposits (Siedlecka and Edwards, 1980; Sweet, 1988; Eriksson and Simpson, 1993; RCe and Hermansen, 1993; Long, 2001 b); those with greater abundance of trough cross-bedding, and some evidence of lateral accretion tend to be interpreted as South Saskatchewan- or Brahmaputra-type rivers (Etheridge et al., 1984; Sweet, 1988; Eriksson and Simpson, 1993; Amireh et al., 1994; Bose and Chakraborty, 1994). Eriksson and Simpson (1993) suggest that monotonous sequences of trough cross-bedded sandstone may reflect relatively constant, perennial discharge in pre-vegetation fluvial systems. An unusual feature of many pre-vegetation sandy systems is the abundance of flat lamination, low-angle cross-stratification and massive beds, suggestive of upper flow-regime conditions and hyperconcentrated flows in ephemeral stream settings (Bhattacharyya and Morad, 1993; Simpson and Eriksson, 1993; Eriksson et. al., 1993; Hjellbakk, 1993, 1997; SCnderholm and Tirsgaard, 1998; Tirsgaard and Oxnevad, 1998; Long, 2001a, b). Flat lamination is not confined to the tops of larger bedforms, as in Platte- and South Saskatchewan-type rivers; Martins-Neto (1994) suggested that the abundance of planar laminated sandstone at the base of sandstone sheets may reflect development of upper-flow regime conditions, at the onset of flooding in ephemeral settings. RCe (1987), RCe and Hermansen (1993) and Hjellbakk (1997) have described extensive sequences dominated by sigmoidaly bedded sandstones, passing into flat laminated and low-angle cross-stratified sandstones. These probably formed at the transition from lower to upper flow-regime conditions, and may be a distinctive feature of sheet-floods in pre-vegetation dry-land fluvial systems. Abundant mudstone clasts in these and other pre-vegetation fluvial deposits suggest that fines may have been deposited in channel thalwegs, and on floodplain surfaces during falling flood stage, but were eroded by subsequent flows.

7. 9. M i c r o b i a l M a t s in the Siliciclastic R o c k R e c o r d

663

Sandy-muddy systems Reliable descriptions of high-sinuosity sandy and sandy-muddy river deposits in the Precambrian are rare (Cotter, 1978; Long, 1978). Sweet (1988)described a possible meandering sequence of pebbly to non-pebbly sandstones with marked lateral accretion surfaces, capped by 3 m of mudstone. These are remarkably similar to surfaces which Rainbird (1992) interpreted as lateral accretion surfaces associated with bar-form migration in an extensive, 150 km wide, sandy braided river system with individual channels up to 20 km across. Mudstone units in this case are interpreted as both over-bank (inter-fluvial) and channel-fill deposits, and are thus not diagnostic of a meandering style. Although flow direction indicated by cross-stratification appears to be at a high angle to the lateral accretion surfaces in both cases, too few cross-bed directions are recorded to confirm systematic changes in flow up-section. Architecture Although architectural element analysis of post-Devonian fluvial strata is now commonplace and has been used to extract useful information (Miall, 1996), little attention has been paid to the anatomy of pre-vegetation fluvial systems. Only a limited number of studies (RCe, 1987; Rainbird, 1992; RCe and Hermansen, 1993; Bhattacharyya and Morad, 1993; Amireh et al., 1994; Hjellbakk, 1997; Tirsgaard and Oxnevad, 1998; SCnderholm and Tirsgaard, 1998; Chakraborty, 1999; Long, 2001a, b) provide substantial architectural detail; some (Chakraborty, 1999; Van der Neut and Eriksson, 1999) have attempted to use this information to reconstruct palaeoflow conditions. Future research should concentrate on detailed architectural analysis of laterally extensive outcrops. Inclinations of all surfaces should be measured to allow a better understanding of stream sinuosity, geometry of bar forms, and hydrologic significance of these systems. Close attention should be paid to modem dry-land fluvial systems (Tooth, 2000) as these may provide useful clues to the behaviour of pre-vegetation fluvial systems.

7.9.

MICROBIAL MATS IN THE SILICICLASTIC ROCK RECORD: A SUMMARY OF DIAGNOSTIC FEATURES

J. SCHIEBER

Introduction Although in many instances quite subtle and often overlooked, microbial communities are nonetheless a ubiquitous component in many modem siliciclastic depositional environments. In modern environments the overall impact of microbial communities on sedimentation processes is somewhat diminished as a consequence of metazoan grazing. In the Precambrian, in contrast, they probably colonised most surfaces where their moisture, light, and nutrient requirements were met (e.g., Hagadorn et al., 1999; Schieber, 1999) (see also section 7.10). The Precambrian Earth: Tempos and Events Edited by P.(}. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Chapter 7: Sedimentation Through Time

Typically, the organic components of these communities are degraded upon burial, and what remains is mainly the impact they had on physical and chemical sediment properties (e.g., erodibility, cohesion, redox conditions and authigenic minerals). These indirect indicators are in a way analogous to trace fossils, in that the absence of a preserved trace-maker does not preclude the preservation of a record of animal-sediment interactions. As far as studies of modern examples of microbial mats in siliciclastic sediments are concerned, most progress has been made with regard to mats in shallow marine and tidal settings, primarily because of ease of access (e.g., Gerdes at al., 2000). While still lagging behind research on microbial mat recognition in carbonate rocks, work on microbial mats in siliciclastic sediments has accelerated substantially in the past few years (e.g., Hagadorn et al., 1999; Schieber, 1999; Pfltiger, 1999; Gehling, 1999). As a result, there is now a much larger array of sedimentary features to draw upon when searching for microbial mats in the siliciclastic rock record. A schematic summary of sedimentary features attributed to microbial mats in mudstones and sandstones (Figs. 7.9-1 and 7.9-2) and this short narrative provide the necessary leads to the relevant in-depth literature. Generally speaking, microbial mats influence the depositional fabrics of sedimentary rocks across a broad spectrum of physical, biological and chemical processes. Their imprints have long been neglected in sedimentological research, in part because knowledge of modern analogues was lacking, and in part because of their cryptic nature. The most telling features that attest to former presence of mats are usually those that indicate uncharacteristic sediment cohesiveness (e.g., for a layer of sand), impermeability (e.g., to gas), tensile strength, erosion resistance, and geochemical behaviour during early diagenesis (see also section 7.10). Microbial Mat Features in Sandstones

Figure 7.9-1 provides an overview of features that might be found in sandstones where microbial mats flourished in the past. The processes resulting in these features are arranged clockwise, along a continuum from active mat growth to final destruction during diagenesis. Proterozoic examples of a few of these features are shown in section 7.10. Mat growth Binding, trapping and baffling are typical processes associated with mat development (Gerdes et al., 2000). Depending on the amount of time available for unhindered mat growth and the overall rate of sediment supply, mats may develop (1) as layers of intermingled microbial filaments and extracellular polymers with little mineral content (up to several cm thick), or (2) as thin biofilms of intermingled filaments and sand grains. The latter tend to stabilise sediment surfaces after episodes of physical reworking. Microbial binding "freezes" surface morphology and can in that way lead to, firstly, surfaces with palimpsest ripples (Fig. 7.9-1a) when new sediment is brought in (Pfltiger, 1999) and, secondly, to surfaces with multi-directional ripple marks (Noffke, 1998). With sufficient

7.9. Microbial Mats in the Siliciclastic Rock Record 665

Fig. 7.9-1. Overview of features that might be found in sandstones where microbial mats flourished in the past. Processes that produce these features are arranged clockwise, along a continuum from active mat growth to final destruction during diagenesis.

oI

E2

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Chapter 7: Sedimentation Through Time

energy, partial erosion of ripple crests may occur, revealing an erosion-resistant top veneer as narrow ridges (Fig. 7.9-1 b; Pfltiger, 1999) (see also Fig. 7.10-1, next section). Aside from surface features, sediment binding and trapping can also produce characteristic lamina features. For example, in Figure 7.9-1c, graded laminae record brief depositional events, whereas laminae with horizontal mineral grains record episodes of mat formation (Noffke et al., 1997). In these latter laminae sand grains were either embedded horizontally, or rotated into a horizontal orientation as the mat decomposed and compacted. Although potentially useful, this feature has not yet been reported from the rock record, and may be difficult to differentiate from other compaction-related features. Biolamination can also be caused by lamina-specific grain selection, such as enrichment with heavy minerals (Fig. 7.9-1d; Gerdes et al., 2000) or micas (Garlick, 1981, 1988). The energy level of the environment is another factor in microbial mat development. Under conditions of high current or wave activity three-dimensional forms such as domes (Fig. 7.9-le) may develop (favoured by rapid synsedimentary lithification), whereas at low energy levels planar forms are prevalent (Hoffmann, 1976; Sami and James, 1993). Domal structures in sandstones have been reported from various Proterozoic and Phanerozoic occurrences (Davis, 1968; Garlick, 1981, 1988; Schieber, 1998). Interaction between the different members of microbial mat communities, filament abundance, water depth and flooding history (Horodyski, 1977b; Horodyski et al., 1977; Gerdes et al., 2000) can lead to a wide range of surface morphologies, including tufts, pinnacles, and pustules (Fig. 7.9-1f), bulges and reticulate ornamentation that has been described as "elephant skin" (Gehling, 1999; Fig. 7.9-lg), and a variety of wrinkle structures (Hagadorn and Bottjer, 1997, 1999; Schieber, 1998, 1999; Fig. 7.9-1h). Although of lesser preservation potential than comparable structures in carbonate producing environments, there is a growing number of reports on these features from Proterozoic sandstones worldwide (e.g., section 7.10). Winds, currents, and gas development, as well as intermittent drying, can lead to intermittent disturbance of mat growth and produce buckling, doming, and rupturing of microbially bound surface layers. Modern examples of such antiform structures in microbial mats have been described as petees (Reineck et al., 1990; Gerdes et al., 1993), and ancient examples have been identified by Gehling (1999). Depending on the intensity of disruption, simple polygonal networks of petee ridges (Fig. 7.9-1i), or complex sinuous ridges with rupturing of microbial surfaces (Fig. 7.9-lj) may be seen.

Metabolic effects Study of modern mats indicates that a metabolic process, such as photosynthesis, can shift carbonate solubility within mats sufficiently to lead to carbonate precipitation between and along the filaments of growing mats (Krumbein, 1974, 1986; Gerdes and Krumbein, 1987; Chafetz and Buczynski, 1992; Chafetz, 1994). Visible effects in the rock record may be the formation of irregular ooids (Gerdes and Krumbein, 1987), disseminated carbonate grains (e.g., Kropp et al., 1997), micritic cement between terrigenous grains, and highly lamina-

7. 9. Microbial Mats in the Siliciclastic Rock Record

667

specific carbonate cementation of otherwise terrigenous laminae (see also section 6.4). The presence of high Mg concentrations in sheaths of filamentous cyanobacteria may also favour the formation of very early diagenetic dolomite (Gebelein and Hoffman, 1973). In sandstones, due to their inherent high permeability, it is very likely that these essentially syngenetic signatures are overprinted by subsequent diagenetic processes. Certain textural features, however, such as terrigenous grains "floating" in a carbonate matrix, would be suggestive of precompactional and possibly syngenetic carbonate formation (Garlick, 1988; Schieber, 1998). In addition, highly lamina-conformable distribution of pyrite may be reflective of the activity of sulphate reducing bacteria beneath the photosynthetic surface layer (Schieber, 1989). Physical mat destruction

Drying out of mat-bound sand layers can either lead to polygonal or incomplete crack networks that are themselves filled with sand (Fig. 7.9-1 k), as well as complexly superimposed sets of spindle-shaped cracks (Bouougri and Porada, 2002; S. Banerjee, 2001, pers. comm.; not illustrated here). The critical observation in that case is that the cracks be in a sand layer (e.g., section 7.10). In a non-mat sand layer, the inherent grain support makes shrinkage impossible, thus a shrunken sand layer must have had an additional component that could shrink during dehydration. In the absence of clays, which could produce similar features during dewatering, a water-rich microbial substrate is the most likely candidate. A special case of this type of sand-based crack is sinuous-circular cracks known as Manchuriophycus (Fig. 7.9-11), probably formed in ripple troughs with thicker mat development (Pfltiger, 1999; Gehling, 2000). Microbial mats may also maintain some of their initial cohesiveness for some time after burial. Thus, during deformation, microbially-bound sand layers may show contrasting behaviour to over- and underlying layers of loose sand. Non-penetrative microfaults (Fig. 7.9-1m) in sand have been interpreted as indicative of microbial mats by Pfltiger (1999) and Gehling (1999). Although microbial mats render a sand surface substantially more resistant to erosion (Neumann et al., 1970), erosion and reworking will commence once currents are sufficiently strong. The binding of the sand surface, however, leads upon erosion to sedimentary features that are distinctively different from those expected from erosion of a loose grain substrate. For example, local erosion of mats can expose underlying sand to wave and current action, leading to rippled patches in an otherwise smooth surface (Fig. 7.9-1n) (see also section 7.10). This feature has been observed on modern tidal flats (Reineck, 1979; Gerdes et al., 1985), as well as in the rock record (MacKenzie, 1972; Reineck, 1979; Schieber, 1998). The cohesiveness of mat-bound sand surfaces also leads to formation of flipped over edges (Fig. 7.9-1o) of partially eroded mat surfaces, as well as redeposition of deformed and rolled up mat fragments (Figs. 7.9-lp and 7.9-1 w). Reports on modern examples include those by Reineck (1979) and Gerdes et al. (2000), and on ancient exampies, those by Schieber (1998, 1999), Garlick (1981, 1988), Simonson and Carney (1999) and Eriksson et al. (2000) (see Fig. 7.10-3b, next section).

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Microbial sand chips (Pfltiger and Gresse, 1996; Bouougri and Porada, 2002) are a variation on this theme (Fig. 7.9-1 q). They are typically smaller (a few cm at most) than the irregular and rolled up mat fragments pictured in Figure 7.9-1 p, of similar size in a given occurrence, plastically deformed, and often current-aligned (Pfltiger and Gresse, 1996) and even imbricated (Bouougri and Porada, 2002). These observations suggest that microbial sand chips are a subclass of eroded mat fragments, abraded and sorted due to a longer transport history. Dried-out microbially-bound sand surfaces typically yield rigid curved chips (several cm across; Fig. 7.9-1 v) that can be transported and form intraclasts in high energy sand deposits. Fossil examples are reported by Garlick (1988) and Schieber (1998) (Fig. 7.10-3a, next section). In the absence of textural differences (grain size, lamination) between sand chips and their sand matrix, diagenetic effects (mat decay mineralisation) related to the organic content of the former (Garlick, 1988; Schieber, 1999) may be the only clue to their recognition (see below).

Mat decay and diagenetic effects Gas development from decaying portions of microbial mats can lead to physical disturbance of the sediment and disruption of surface mats. Observed features are gas domes and convoluted internal lamination (Fig. 7.9-1r), produced by gas buildup beneath mats (Gerdes et al., 2000; Bouougri and Porada, 2002), as well as ruptured gas domes termed "Astropolithon" (Pfltiger, 1999; Fig. 7.9-1 s). In the latter case, the substrate cohesiveness that is implicit in the radial ruptures of the dome (Fig. 7.9-1 s) is again a good indication of the former presence of a mat. Gas development also contributes to the formation of the more severely disturbed and ruptured petee structures (Fig. 7.9-lj). Kinneyia style ripples (Fig. 7.9-1t) show considerable similarity to wrinkled mat surfaces (Figs. 7.9-1g and 7.9-lh). On account of the steep slopes of their troughs and their flat tops, however, they were interpreted by Pfltiger (1999) to reflect gas trapping beneath flat mats. Whereas those described by Pfltiger (1999) really seem to represent gas trapping beneath mats, many Kinneyia described in the literature show more resemblance to the round crested microbial wrinkle marks described by Hagadorn and Bottjer (1999). Thus, attention to detail is clearly needed to interpret properly wrinkled surface features. Due to the permeability of sand, organic matter is readily metabolised by microbes during early burial, making it unlikely that organic matter will survive as a microbial mat indicator. Fortunately, microbial mats also constitute sharply defined geochemical boundaries (Bauld, 1981), and anaerobic decay beneath mats favours formation of "anoxic" minerals such as pyrite, siderite, and ferroan dolomite. Cementation of sand grains by these minerals constitutes "mat-decay mineralisation" (Schieber, 1998). Ghosts of filaments may be preserved in these cements. Observing thin, stratiform horizons of these minerals (Fig. 7.9-1u) in a shallow water sandstone (above wave base) is suggestive of the former presence of microbial mats (Gerdes et al., 1985; Garlick, 1988). Depending on water chemistry (e.g., marine versus freshwater), different minerals will be favoured (e.g., pyrite versus siderite). Burial of rigid (Fig. 7.9-1 v) or soft fragments (Figs. 7.9-1 q and 7.9-1 w) of resedimented mat can, upon

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decay, give rise to comparable cementation that preserves the former outline of transported mat fragments (Garlick, 1988; Pfltiger and Gresse, 1996; Schieber, 1998). Microbial Mat Features in Shales

Figure 7.9-2 provides an overview of microbial mat features that might be found in ancient mudstones. As in Figure 7.9-1, causative processes are arranged clockwise, and their effects illustrated with drawings and photographs. Mat growth Although binding, trapping, and baffling are equally well associated with microbial mats on muddy substrates, preservation of surface relief is of a more subtle nature, due to the intrinsically high degree of compaction. Nonetheless, the initial surface relief leads to wavy-crinkly laminae (Fig. 7.9-2a) that are distinctively different from the parallel laminae that form in mudstones as a result of physical sedimentation processes (Schieber, 1986; Fairchild and Herrington, 1989; O'Brien, 1990; Goth, 1990; Wuttke and Radtke, 1993; Goth and Schiller, 1994). There are also examples where mat colonisation of an irregular surface (e.g., an intraclast conglomerate) had a smoothing effect (Fig. 7.9-2b). In non-mat mudstones, compactional effects over comparable relief tend to be visible for a greater distance upward from the underlying surface irregularities. Surface stabilisation by mat cover can also be deduced from differences in loading behaviour (Schieber, 1986). For example, in mudstone units where silt layers were deposited on mat-bound surfaces as well as on non-mat muds, comparable silt layers produce miniature ball-and-pillow structures on the latter (Fig. 7.9-2c), and only minor load features on the former (Fig. 7.9-2d). Whereas the wavy-crinkly carbonaceous laminae discussed above have been reported mainly from inferred subtidal and shelf deposits (Schieber, 1986; Fairchild and Herrington, 1989; Logan et al., 1999), domal buildups of various amplitude and spacing have been observed in nearshore mudstones (Figs.7.9-2e and 7.9-2f; Schieber, 1998). It is quite likely that the inherent rapid weathering of mudstones has thus far concealed a variety of other occurrences in the rock record from scrutiny. By burying a growing mat under a sudden influx of sediment, event sedimentation (storms, floods) can cause interruption of mat growth. Intermittent event sedimentation in an area of mat growth can lead to "striped shales" with alternating mat and event layers (Fig. 7.9-2g; Schieber, 1986; Logan et al., 1999). Occasional deposition of thin clay drapes in areas of incomplete but expanding mat cover may lead to false cross-lamination (Fig. 7.9-2h) at the edge of expanding mat patches. In this situation mats re-establish themselves (vertical movement of filaments) on top of recently deposited clay drapes and expand laterally (Schieber, 1986). In many instances, the resulting false cross-lamination probably will look quite a bit more irregular than as depicted in Figure 7.9-2h.

670 Chapter 7: Sedimentation Through Time

Fig. 7.9-2. Overview of microbial mat features that might be found in ancient mudstones. Processes that produce these features are arranged clockwise, along a continuum from active mat growth to final destruction during diagenesis.

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The processes leading to petee structures are not dependent on a particular substrate (e.g., Reineck at al., 1990; Gerdes et al., 2000), and analogous structures (but at a smaller scale) occur in modem mud puddles. It is probably only a matter of time before fossil analogues will be recognised in the rock record. Enrichment of mat laminae with mica flakes (Fig. 7.9-2i) is one type of lamina-specific grain selection that has been observed in mud-based microbial mats (Schieber, 1998). Just as for sandy microbial mats, the underlying causes for this type of grain enrichment are not well understood (Gerdes et al., 2000).

Metabolic effects Just as in sandy microbial mats, it is to be expected that syngenetic carbonate precipitation associated with mats growing on a muddy substrate will also occur. Observation of randomly oriented (instead of subhorizontal) mica flakes in conformable carbonate-rich laminae can, for example, be a suggestion of syngenetic carbonate deposition (Schieber, 1998). Cementation later in burial history would most likely be accompanied by partial rotation of mica flakes into the horizontal. Terrigenous grains floating in a carbonate matrix similarly would suggest essentially syngenetic carbonate precipitation. Although pyrite formation also happens quite early, because it results from mat decay under anaerobic conditions, it is considered with diagenetic effects (see below). In cases where bituminous substances can still be extracted from suspected fossil mat deposits, carbon isotopes and biomarkers, in conjunction with determination of sulphur isotopes, can be used to deduce the likely metabolic pathways operating at the time of deposition (Brassell, 1992; Logan et al., 1999). These biomarkers may help to determine whether a mat system was dominated by cyanobacteria (oxygenic photosynthesis), photosynthetic sulphur bacteria (anaerobic photosynthesis), or sulphide-oxidising bacteria (chemoautotrophy; Gallardo, 1977; Williams and Reimers, 1983). Because of the implications for the global cycling of carbon and sulphur (see also sections 3.2, 5.3 and 5.5), the differentiation of photosynthetic mats from non-photosynthetic and sulphide oxidising types, as well as the magnitude of microbial mat involvement in black shale formation, is of considerable interest. Physical mat destruction Sedimentary features produced by erosion of mat-bound mud surfaces are broadly similar to those observed in the erosion of sandy microbial mats (Figs. 7.9-1n, o, p). Flipped over mat edges (Fig. 7.9-2j), overfolded mat layers (Fig. 7.9-2k), and "roll-up" structures of various size have all been observed in ancient examples (Schieber, 1986, 1998, 1999). Mat layers distinguish themselves from other mud layers by their display of "within layer" cohesiveness upon erosion and transport (Fig. 7.9-21), as well as by rheological differences between mat layers (firm-doughy, less compactable) and normal mud (soft-fluid, yogurtlike; Fig. 7.9-2k; Schieber, 1986). Because the tearing of a mat is analogous to the tearing of a fibrous fabric, torn mats tend to display frayed edges (Fig. 7.9-2m). This phenomenon has been termed "blotting

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paper effect" in studies of modern mats (Gerdes et al., 1993), and has also been described from fossil examples (Schieber, 1999). Although desiccation of muddy microbial mats will produce cracks and dried-out mat chips, recognition in the rock record is not a trivial task. While in the case of sandy surfaces, shrinkage features (Fig. 7.9-lk) and coherent transport (Fig. 7.9-1 q) are highly suggestive of a binding material with high water content (such as a mat), muds are already watery and coherent in the absence of mats. Thus, even without a mat they will crack and produce chips when dried. Though modern mats on muddy substrate tend to modify crack morphology and crack edges (Gerdes at al., 1993), to date I am unaware of any systematic documentation of desiccation effects in ancient mat-bound muddy sediments. One example of desiccation in mats on a muddy substrate concerns the drying-out of thin mats (biofilms) covering mudflat surfaces. As these microbial films dry out they crack and curl up, and can then be transported by wind (Trusheim, 1936) and water (Fagerstrom, 1967). Because these fragments resist compaction upon redeposition, they leave irregular impressions on mudstone bedding planes (Fig. 7.9-2n), that on occasion are reported from the rock record (Horodyski, 1982, 1993). Dried-out mat fragments can also float out into open water bodies (Fagerstrom, 1967), and thus transport detrital grains from nearshore regions to deeper portions of a water body. Clusters of coarser grains (Fig. 7.9-2o) that occur in otherwise "pure" mudstones may thus be explained as material that was "rafted in" by mat fragments from nearshore areas, and buried collectively once a fragment had sunk to the bottom (Olsen et al., 1978; Schieber, 1999). In Phanerozoic sediments care has to be taken to eliminate alternative mechanisms, such as rafting-in with plant debris and animal carcasses (buoyed by decomposition gases), as well as by fecal pellets.

Mat decay and diagenetic effects An effect similar to that produced by grain rafting via dried-out mat fragments may also occur when gas formation in submerged mats, either due to photosynthesis or to decay processes, induces portions of the mat to detach from the substrate and to float upward (Fagerstrom, 1967). Attached coarser grains may then be rafted offshore and give rise to clusters of coarse grains within a much finer matrix (Fig. 7.9-2o). Anaerobic decay of organic matter beneath a growing mat is a favourable environment for precipitation of "anoxic" minerals, such as pyrite, siderite, and ferroan dolomite. In marine settings, this sub-mat decay typically leads to production of hydrogen sulphide and to pyrite formation (Berner, 1984). Depending on the availability of iron, manifestations in the rock record can range from carbonaceous laminae dusted with tiny pyrite grains (Schieber, 1989), to strongly pyritic laminae (Fig. 7.9-2p) that closely follow the original organic laminae and mimic the wavy-crinkly mat lamination (Fig. 7.9-2a; Schieber, 1989). Later diagenetic effects may include pyrite overgrowth and cementation of the finegrained original pyrite (Strauss and Schieber, 1990), as well as recrystallisation and enlargement of carbonate minerals in layers with syngenetic carbonate accumulations (Fig. 7.9-2q). Maturation of organic matter upon further burial leads to reduction of organic content (hydrocarbon formation), as well as to gradual destruction of biomarkers and

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kerogens. In contrast to sandstones, however, the low permeability of mudstones thwarts complete organic matter destruction and leads to preservation of anastomosing carbonaceous laminae (Fig. 7.9-2r).

7.10.

MICROBIAL MAT FEATURES IN SANDSTONES ILLUSTRATED

S. SARKAR, S. BANERJEE AND EG. ERIKSSON Biota are seldom preserved, and sedimentary features resulting from microbial mats growing on sandy substrates tend to be lost readily upon lithification. Prolific mat growth, nonetheless, can make sand cohesive and even thixotropic like mud. Non-uniformitarian mat growth thus resulted, uncommonly, in the preservation of a host of structures (section 7.9; Figs. 7.9-1 and 7.9-2) in Proterozoic sandstones that seldom attract attention, despite their significance for helping to interpret depositional systems. Examples from different formations from India and South Africa are divided here into four broad categories: (1) inherited structures (palimpsest) (Fig. 7.10-1), (2) deformation structures, comprising (a) brittle, (b) ductile examples (Fig. 7.10-2), and (3) derived structures (Fig. 7.10-3). Prolific microbial mat growth can be traced back in time to at least 2.4 Ga, and continued throughout the Proterozoic eon. Mats encroached on the terrestrial setting before 1.8 Ga. Ubiquitous mat growth must have restricted sediment reworking. As a consequence, bedform evolution and preservation were often different in Proterozoic depositional systems in contrast to their Phanerozoic or laboratory equivalents.

Fig. 7.10-1. (a) Ripple sets replicated beneath cross-bedded sandstone. This replication, particularly for the set with very small magnitude, would have been impossible without mat cover. 2.3 Ga Chaibasa Sandstone, India (marine). (b) Patchy reworking of a set of wave ripples by a secondary flow. The first generation ripples were most likely largely protected by a microbial mat. Circa 0.6 Ga Jodhpur Sandstone, India (marine). The Precambrian lz~rth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.I.J. Mueller and O. Catuneanu

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Fig. 7.10-2. (a-i) Cracks, resembling lips, on wave ripple crests; probably formed due to tensile shear and most likely became accentuated along highs on a leathery mat cover. 1.6 Ga Chorhat Sandstone, India (marine); (a-ii) Cross-cutting and steep-sided grooves, inferred to have formed as mat-covered granular sand became amenable to synaeresis cracking, analogous to such cracks in mud. 1.6 Ga Chorhat Sandstone, India (marine); (a-iii) Cross-cutting ridges, formed by sand flow under confining pressure, around synaeresis cracks, and preserved due to protection by a mat. Circa 0.6 Ga Jodhpur Sandstone, India (marine); (b-i) Wrinkle marks in sandstone, presumably formed on a mat under gentle shear resulting from fair weather waves. Circa 0.6 Ga Jodhpur Sandstone, India (marine); (b-ii) Numerous load marks at bed sole (left block) and load casts (right block) on bed top. A gelatinous microbial mat led presumably to the sand becoming thixotropic on top of the bed carrying the casts. An overlying thin sandstone bed (top left of right hand block; see coin) bears, on its top, a palimpsest ripple obviously inherited from the underlying bed. 1.6 Ga Chorhat Sandstone, India (marine); (b-iii) Small sand mounds (thick arrows) formed by upward sand flowage beneath a mat. Minute craters (thin arrows) formed in adjacent areas where the mat was absent or torn away. Circa 0.6 Ga Jodhpur Sandstone, India (marine).

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Fig. 7.10-3. (a) Deformed clasts of well-sorted sand. Microbiota presumably provided the flexible bondage between the non-cohesive sand grains. Circa 0.6 Ga Lower Bhander Sandstone, India (marine); (b) Sand curl (arrow) formed through desiccation, analogously to mud curls. Microbial protein and polysacharids presumably made the sand cohesive. Circa 1.8 Ga Waterberg Group, South Africa (aeolian interdune deposit) (section 7.6). 7. l 1.

SEDIMENTATION RATES

P.G. ERIKSSON, EK. BOSE, S. SARKAR AND S. BANERJEE Sedimentation is continuous only locally and over short time periods (<~ several 103 years), mostly in deep marine and prograding deltaic environments fed by perennial rivers (Reading, 1996). Short-term spikes in sedimentation rate reflect event deposits, evaporites and volcaniclastic deposition. Sedimentation rates are either instantaneous to short-term, measured essentially at the depositional system scale, or average/long-term rates. Application of the former to geological time scales gives rates several orders of magnitude larger than those determined over 106 year time scales (Friedman et al., 1992). Erosion and nondeposition, as well as sediment bypassing, result in a rock record with more gaps than strata (Ager, 1981), and explains this apparent paradox (Blatt et al., 1980). Erosion rate largely reflects uplift rate, which may reach 10 m ky-1 over short periods, exceeding short term erosion rates (generally ~< 1 m ky-I); uplift, like sedimentation, is thus episodic (Blatt et al., 1980). Long-term sedimentation rates primarily reflect the rate of subsidence (mostly ~< 1 m ky-1 ) (Blatt et al., 1980), a part of accommodation space creation, rather than sediment supply (reduced by reworking and recycling) or specific environments (section 7.1). Preservation potential of sediments is critical; littoral and continental sediments are generally most susceptible to both reworking and recycling, although rapid creation of accommodation space (due to varying combinations of subsidence and eustasy) (chapter 8) will alleviate this (Reading, 1996). In applying sedimentation rates to the Precambrian record, chronological resolution becomes a problem. While the average clay particle takes c. 50 years to reach the abyssal floor (Gross, 1990), time resolution ~< 105 years can be obtained from Phanerozoic successions using mainly palaeontology or magnetostratigraphy (Reading, 1996). As radiometric The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Fig. 7.11-1. Sedimentation rates in Bubnoff units (mm ky- I ) for a wide range of basins, palaeoenvironments and deposits. (Data from: Emery and Uchupi, 1972; Berger, 1974; Crowell, 1974; Van Houten, 1974; Schwab, 1976; Blatt et al., 1980; Schopf, 1980; Schreiber and HsiL 1980; Schlager, 1981, 1991, 1992; Brooks and Holmes, 1989; Gross, 1990; Eyles et al., 1991; Glaser and Droxler, 1991; Einsele, 1992; Reading, 1996; Altermann and Nelson, 1998).

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dating only provides a resolution of c. 1-5 My for Precambrian rocks, only long-term sedimentation rates can be estimated. Comparing examples of average sedimentation rates measured from relatively thick Recent deposits to those derived from the sub-Recent rock record supports the importance of sediment recycling. Recent rates thus exceed their rock record counterparts, although chemical shelf and clastic pelagic environments exhibit a reasonable degree of equivalence, due to relatively long-lived tectonic and eustatic stability (Fig. 7.11-1) (see Altermann and Nelson, 1998, for discussion of comparable Archaean and modern sedimentation rates). The highest rates are noted in Recent clastic deltaic systems, and generally there is a trend of decreasing rates from Recent continental clastic --+ marine chemical ~ marine clastic settings (Fig. 7.11-1). Recent carbonates show decreasing rates towards deeper water, as do marine clastics to a certain degree. At the basinal scale of preservation, the influence of rapid tectonic uplift and basin floor subsidence becomes paramount, with concomitant very high sedimentation rates, such as in the pull-apart Ridge Basin, California, or along the uplifted Miocene glaciomarine coast of the Gulf of Alaska (Fig. 7.11-1). Recycling is the most important long-term factor limiting estimation of sedimentation rates; as an example, volcanogenic rocks are more resistant than limestones, which may have been recycled up to five times since 3.0 Ga (Windley, 1995). Pelagic oceanic ooze, mostly removed during subduction, has a half life of c. 50-60 My (Worsley et al., 1984) and would not be preserved easily in the Precambrian record. The proportion of sedimentary rocks observed today is thus radically different to original depositional proportions, and Sm-Nd isotopes suggest that the < 2.5 Ga sedimentary mass has been 85-95% cannibalistic (Veizer and Jansen, 1985a). Consequently, the post-2.5 Ga record reflects crustal evolution (chapters 2 and 3). The fact that Archaean greenstone belts still have a high proportion of mafic volcanic rocks (chapter 4) supports an Archaean Earth dominated by the oceanic realms, and the 30% mafic component in the sedimentary mass today is most likely a relic of Archaean sediments still to be dispersed by recycling (Veizer, 1988).

7.12.

COMMENTARY

EG. ERIKSSON AND M.A. MARTINS-NETO The Hadaean-Archaean transition at 4 Ga (see also section 1.2) may have been marked by a change in Earth's cooling regime from heat exchange directly to the atmosphere (i.e., mid-ocean ridges above sea level) to one where the oceans buffered Earth's heat loss (de Wit and Hynes, 1995). These authors suggest that the growth of continental crust (.section 2.8) in the succeeding 4-3.2 Ga period was related to the initiation of synchronous igneous and hydration processes at oceanic constructive plate boundaries. High heat dissipation at these boundaries produced oceanic crust that was hotter and more buoyant than modern counterparts, and a combination of intra-oceanic island arcs, oceanic plateaus and plate tectonic collisional processes led to proto-continents developing (Windley, 1995) (see, however, section 3.6, for discussion of an alternative model). These small cratonic The Plecambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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plates may have made up to 5-10% of the crust (Eriksson, 1995). Generally in the 4-3.2 Ga greenstone belts, komatiitic, tholeiitic and felsic volcanic and volcaniclastic rocks predominate (Fedo et al., 2001) (chapter 4), with thin remnants of passive margin carbonates, BIF, stromatolitic evaporites, pelites and quartzites, as well as subordinate synorogenic turbidites, conglomerates and sandstones, reflecting increasingly stable small continental nuclei (Windley, 1995). The scarcity of Precambrian fossils (sections 6.2 and 6.3) necessitates greater reliance on detailed study of sedimentary structures (e.g., sections 7.2, 7.4 and 7.5) in these ancient sedimentary deposits. However, microbial micro-organisms were important in trapping, binding and precipitating sediments in situ, to build small carbonate platforms (Wright and Altermann, 2000) (sections 6.2 and 6.4). In addition, direct precipitation of calcium carbonate from sea water may have occurred (Grotzinger, 1989). Microbial mats likely also played an important part in clastic sedimentation during the Precambrian (sections 7.10 and 7.11). Archaean gypsum deposits are seen either as evaporites derived from sea water compositions similar to those today (Lowe, 1983; Buick and Dunlop, 1990) or were related to sites of continental runoff (Grotzinger and Kasting, 1993). Archaean seawater (section 3.2) was probably enriched in iron of seafloor hydrothermal affinity beneath the pycnocline (Veizer, 1983a), with sulphidic iron-formations developing at these depths, and oxidic deposits in areas of photosynthetic productivity above the pycnocline (Eriksson et al., 1997). Eriksson (1983)considers Archaean iron-formations to have been analogous to modern starved basin pelagic sediments (section 5.4 discusses iron-formations in detail). Depositional regimes recognised in Mesoarchaean greenstones are debris-flows on high gradient alluvial fans, low sinuosity rivers, shallow water marine environments with wave and tidal action and turbidity currents, some of the latter deposits associated with hummocky cross-strata thereby indicating deposition near storm wave base (Eriksson et al., 1997) (section 7.3). In general, there are indications for shallow marine conditions (e.g., Windley, 1995; Eriksson et al., 1997), bearing in mind that apart from inferred ophiolites (section 3.7) forming part of highly-deformed greenstone belts (section 7.4), no unequivocal ocean floor of Precambrian age has been preserved (e.g., Fedo et al., 2001). Major crustal growth in the 3.2-2.6 Ga period encompassed amalgamation of oceanic terranes at c. 3.3-3.2 Ga (section 3.6 provides a different model) followed by passive continental margins on the earliest stabilised craton, Kaapvaal, as well as Cordilleran- and Himalayan-style collisions (section 3.8) of growing cratons, and ending with continental flood basalts (Windley, 1995) possibly related to a global superplume (sections 3.2 and 3.3) event (e.g., Eriksson et al., 2002b). Increasing cratonic stability led to rift basins and strike-slip basins becoming common (e.g., Mueller and Corcoran, 1998; Smithies et al., 2001) (section 7.3). Many late Archaean greenstone belts (sections 2.4 and 7.4) reflect syntectonic styles of clastic sedimentation, with alluvial fans and immature braided rivers (section 7.8) passing directly into high energy shallow marine settings, in which aggressive weathering (sections 5.10 and 5.11) related to palaeoatmospheric composition produced mature sandstone close to source areas (Corcoran et al., 1998). Others see mature Archaean sandstones as reflecting periods of crustal stability (Donaldson and de Kemp, 1998). Catastrophic influxes of volcaniclastic debris often choked siliciclastic greenstone

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sedimentation systems (Mueller and Corcoran, 1998) (sections 4.1 and 7.3). Contrasting basin types can be observed at c. 3.0 Ga on the two oldest cratons, Pilbara and Kaapvaal. While the c. 3.0-2.94 Ga Mallina basin in Pilbara bears evidence of sediment recycling, it was still a greenstone-type depository (Smithies et al., 2001). In contrast, the more stable Kaapvaal was characterised by the famous auriferous c. 3-2.7 Ga Witwatersrand basin, where a retroarc foreland basin model has been applied for thick siliciclastic fluvial and littoral sediments (section 7.5) which also contain ventifacts (section 7.6), pointing to aeolian erosion processes (Els, 1998; Catuneanu, 2001). Elsewhere, such as on the Zimbabwe craton and in the Slave Province of northern Canada, continental growth continued through greenstone belt evolution (Hofmann et al., 2001; Mueller and Corcoran, 2001). Continental growth rates are thought to have reached a maximum close to the ArchaeanProterozoic boundary (Eriksson, 1995) (section 2.8), and the 2.6-2.4 Ga period was characterised by widespread (global?) orogenic quiescence (Windley, 1995) (see also section 3.9). A Neoarchaean supercontinent may have been extant during this period when large epeiric (section 7.7) basins and passive margin platforms of chemical and or clastic sedimentary affinity formed (Windley, 1995). Condie et al. (2001) discuss an association of a mantle superplume event and supercontinent formation at c. 2.7 Ga (see also section 3.2). Aspler and Chiarenzelli (1998) suggest that two, partly diachronous supercontinents may have existed. The Archaean-Proterozoic boundary is similarly diachronous, with older cratons such as Pilbara and Kaapvaal stabilising earlier (Windley, 1995). Diagnostic geochemical changes observed across this boundary probably reflect the change from mantle buffering of sea water (hydrothermal interaction of seawater and juvenile crust at mid-ocean ridges) to continental buffering as river discharge from the stable continental shelves became predominant (Veizer, 1983a, b; 1988). Analogous geochemical changes are observed across this boundary for continental pelites (Wronkiewicz and Condie, 1990), resulting from more evolved felsic-rich source rocks, and for the basalts in Palaeoproterozoic greenstone belts which reflect greater depths of magma generation (Condie, 1989). Deposition of the earliest large scale carbonate-BIF platformal sequences occurred in this 2.6-2.4 Ga time interval, with the Hamersley Group of Pilbara and the Chuniespoort-Ghaap Groups (lower Transvaal Supergroup; e.g., Eriksson et al., 2001b) of Kaapvaal reflecting cratonic drowning between c. 2630 and 2430 Ma (Nelson et al., 1999) (section 5.4). Their analogous lithostratigraphy has led to suggestions of a "Vaalbara" supercontinent (e.g., Cheney, 1996) (see also section 1.3), a concept not supported by geochronologic or palaeomagnetic data (Altermann and Nelson, 1998; Wingate, 1998; Nelson et al., 1999; Eriksson et al., 2002b). During the same time interval (2.6-2.4 Ga), the thick BIF-carbonate succession of the Minas Supergroup (S~.o Francisco craton, southeastern Brazil) was deposited (Alkmim and Marshak, 1998). From c. 2.4--2.2 Ga, Windley (1995) suggests supercontinental fragmentation as a predominant influence on sedimentation, with rifting and passive margin development being pre-eminent. Partial to full Wilson cycles appear to be pertinent to a number of parts of the postulated Kenorland (e.g., Aspler and Chiarenzelli, 1998) supercontinent, with prime examples being described from the Huronian Supergroup (c. 2.4-2.2 Ga, Superior craton; Young et al., 2001), Karelian Supergroup (c. 2.45-1.9 Ga, Fennoscandian shield;

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Ojakangas et al., 2001a), Hurwitz Group (c. 2.45-1.9 Ga, Hearn domain, Canada; Aspler et al., 2001) and Lake Superior region (c. 2.4-2.2 Ga, Superior craton; Ojakangas et al., 200 l b) (see also section 3.9). While fragmentation of Kenorland proceeded, a "southern" (present-day frame of reference) continent appears to have begun assembly (Eriksson et al., 1999b). Tripartite cycles of glacial diamictites-deltaic mudrocks-fluvial arkoses in the Huronian are related by Young et al. (2001) to self-regulating changes in atmospheric CO2 contents. They infer that high weathering rates at low palaeolatitudes within Kenorland encouraged atmospheric CO2 drawdown, and that reduced weathering during icehouse conditions in combination with volcanic gas emissions resulted in a return to a greenhouse state (sections 5.2, 5.6 and 5.7). Similar deposits can be correlated across large portions of the inferred Kenorland supercontinent, and it is surmised that breakup interfered with this self-regulatory cyclicity in icehouse and greenhouse conditions (Young et al., 2001). It is also thought that significant quantities of free oxygen became available in the atmosphere during the c. 2.4-1.9 Ga period (e.g., Windley, 1995) (see, however, section 5.2), as evidenced by red beds with haematitic grain coatings (e.g., Eriksson and Cheney, 1992), haematite-rich palaeosols (e.g., Wiggering and Beukes, 1990) and supergene oxidic ores developed on BIF (Holland and Beukes, 1990) (sections 5.3-5.5). Major continental crustal growth from c. 2.0-1.7 Ga is interpreted from areas such as the southwestern U.S.A., Western Greenland, the Baltic shield, the Birimian belt in West Africa, and from Brazil (Nelson and DePaolo, 1985; Patchett and Arndt, 1986; Mil6si et al., 1992; Schrank and Silva, 1993; Windley, 1995). Crustal growth processes appear to have been analogous to modern island arcs and accretionary prisms, thus resembling also Archaean greenstone belts, and were associated with widespread post-orogenic granitisation (Windley, 1995). The Laurentia supercontinent is thought to have amalgamated at c. 2.01.7 Ga (Hoffman, 1988; Aspler et al., 2001). A global mantle superplume event (sections 3.2 and 3.3) was likely associated with supercontinental formation at c. 1.9 Ga (Condie et al., 2001). Following c. 1.6 Ga, the supercontinental cycle became well developed, with the classic association of predominant non-marine continental sediments related to assembly phases, and drowned platformal marine deposits reflecting fragmentation and dispersal (e.g., Hoffman, 1989c, 1991; Barley and Groves, 1992; Windley, 1995). Accretion of island arcs continued as well, as is still the case today. In the Neoproterozoic, an association of increased atmospheric oxygen, supercontinent assembly and breakup, and global glaciations reminiscent of the c. 2.4-2.2 Ga event occurred (e.g., Fedo and Cooper, 2001; Dehler et al., 2001; Martins-Neto et al., 2001) (sections 3.10, 3.11, 5.7 and 5.8).

The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geok~gy, Vol. 12 (K.C. Condie, Series Editor) 9 2004 Elsevier B.V. All rights reserved

681

Chapter 8

SEQUENCE STRATIGRAPHY AND THE PRECAMBRIAN

8.1.

INTRODUCTION

A.F. EMBRY, O. CATUNEANU AND P.G. ERIKSSON

History of Sequence Stratigraphy Some of the stratigraphic surfaces which are herein included as part of sequence stratigraphy have been recognised ever since James Hutton first described a subaerial unconformity over 200 years ago. One hundred and thirty years later Joseph Barrell (1917) clearly expressed the concept that nonconformities are generated by changes in base level and succinctly noted that "sedimentation controlled by base level will result in divisions of the stratigraphic series separated by breaks". The publication of Barrell's seminal paper may be regarded as a precursor of modern day sequence stratigraphy. Barrell's concepts lay in limbo for 30 years until Larry Sloss (Sloss et al., 1949) emphasised the utility of using subaerial nonconformities as the boundaries of a stratigraphic unit he termed a sequence. In the 1950s and 1960s Harry Wheeler laid out the basic theoretical framework of sequence stratigraphy in a remarkable set of papers (Wheeler and Murray, 1957; Wheeler, 1958, 1959, 1964). Unfortunately Wheeler introduced a complex jargon (e.g., "degradational vacuity") and the importance of his concepts, which built on Barrell's (1917) work, was not appreciated or utilised by either academic or industrial geologists. In 1977, with the publication of a series of papers on seismic stratigraphy by geologists and geophysicists from Exxon (Payton, 1977), sequence stratigraphy came into the mainstream of stratigraphic thought. This watershed publication made it abundantly clear that the stratigraphic record is punctuated by numerous nonconformities and that it can be divided into a series of units that are partially bound by nonconformities. These units were designated as sequences following Sloss' work, and the term sequence was expanded from a unit bound by nonconformities to one that is bound by both nonconformities and correlative conformities. This was a very important modification because it allowed a sequence to be recognised potentially over an entire basin rather than just on the flanks where the bounding nonconformities are present. The Exxon seismic data demonstrated that sequences are the most practical units to use for stratigraphic subdivision if one wants to use such units to describe and interpret the depositional history of a stratigraphic succession. One unfortunate sideline of these publications is that they strongly advocated the hypothesis that the sequences were primarily the product of eustatic sea level changes without providing any solid evidence to support such an interpretation. This interpretation stood in contrast to that of Sloss (1963), who had always emphasised tectonics as the prime driver

682

Chapter 8: Sequence Stratigraphy and the Precambrian

of sequence generation. This debate of tectonic versus sea level control of the base level shifts that produce nonconformities and their enclosed units, sequences, continues today (Miall and Miall, 2001). The next major contribution to sequence stratigraphy was the publication of S.E.P.M. Special Publication 42 in the spring of 1989 (Wilgus et al., 1988). A series of papers authored by Exxon scientists (Jervey, 1988; Posamentier and Vail, 1988; Posamentier et al., 1988) discussed the recognition and origin of sequences in the rock record, and proposed a sequence stratigraphic model that uses a depositional sequence as its primary stratigraphic unit. On the basin margin, subaerial nonconformities are the designated boundaries of a depositional sequence. Basinwards, beyond the extent of the nonconformities, the boundaries of a depositional sequence are extended along time lines referred to as "correlative conformities". These conformities may be traced with relative confidence on seismic lines, but are virtually impossible to recognise in core, well logs, and most outcrops. The lack of discernable lithological characteristics of the correlative conformity continues to cause problems today. Depositional sequences are divided into three systems tracts on the basis of internal boundaries that coincide with the start of transgression (called the maximum regressive surface herein) and the start of regression (called the maximum flooding surface herein). In ascending order, these systems tracts are called the lowstand, transgressive and highstand systems tracts and the development of each is related to a specific portion of a sinusoidal sea level curve. This model, despite a number of shortcomings, became very popular in the 1990s and is still referred to in many publications. Hunt and Tucker (1992) proposed a fourth systems tract, which has been given a variety of names including falling stage and forced regressive. These authors and subsequent proponents of this systems tract (e.g., Helland-Hansen and Gjelberg, 1994; Plint and Nummedal, 2000) use the oldest portion of the regressive surface of marine erosion and the younger portion of the subaerial unconformity to demarcate its limits. Practical problems related to the conformable portions of the systems tract boundaries do exist, however, due to their lack of discernable lithological characteristics. Van Wagoner et al. (1990) further elaborated upon the Exxon sequence model and provided a number of surface and subsurface examples of its usage. This publication also introduced the parasequence, a small-scale unit consisting of a succession of shallowing-upwards facies. At approximately the same time as S.E.P.M. 42 was published, Galloway (1989), based on Frazier (1974), proposed that maximum flooding surfaces and not subaerial nonconformities be used as the sequence boundary. He called such a unit a genetic stratigraphic sequence and it has also been referred to as an R-T (regressive-transgressive) sequence. In the early 1990s, Embry and Johannessen (1992) proposed a third type of sequence that was termed a transgressive-regressive or T-R sequence. It was similar to the depositional sequence of Exxon in that it used nonconformities for the boundaries on the basin margin. However, unlike a depositional sequence that has a nondescript correlative conformity as the basinwards extension of the sequence boundary, the T-R sequence employs the recognisable boundary at the start of transgression for the basinwards extension of the sequence boundary. This surface is referred to as a maximum regressive surface herein. Only two

8.1. Introduction

683

I

Sequences I Sloss (1949,1963)

DepositionalSequence I (Seismic Stratigraphy) Mitcnum el al. (1977)

IIIII

II~lt

Sequence Stratigraphy

DepositlonalSequenceII Haq et al. (1987) Posamentieret al. (1988)

DepositlonalSequenceIII Van Wagoneret al. (1988,1990) Christie-Blick (1991)

GeneticSequences Galloway (1989) Frazier (1974)

Deposltlonal Sequence IV Hunt and Tucker (1992,1995) Plint and Nummedal (2000)

T-R Sequences Embry (1993,1995) Curray (1964)

Fig. 8.1-1. Family tree of sequence stratigraphy (from Catuneanu, 2002, and modified from Donovan, 2001). The various sequence stratigraphic models mainly differ in the style of conceptual packaging of strata into sequences, i.e., with respect to where the sequence boundaries are picked in the rock record. systems tracts, transgressive and regressive, were recognised for a T-R sequence with the maximum flooding surface forming the boundary between the two. The various sequence models that are currently in use differ from each other mainly in the style of conceptual packaging of the stratigraphic record, using a different timing for systems tract and sequence boundaries in relation to a cycle of base level shifts (Figs. 8.1-1 and 8.1-2). Each sequence model may work best under particular circumstances, and no one model is universally applicable to the entire range of case studies (Catuneanu, 2002). Recent syntheses of the sequence stratigraphic concepts and methods have been published by Galloway and Hobday (1996), Emery and Myers (1996), Miall (1997), Gradstein et al. (1998), Posamentier and Allen (1999) and Catuneanu (2002).

Application of Sequence Stratigraphy to Precambrian Deposits The application of sequence stratigraphy to the Precambrian rock record is still in its infancy. Recent work on Precambrian sequence stratigraphy includes Christie-Blick et al. (1988), Beukes and Cairncross (1991), Krapez (1996, 1997), Catuneanu and Eriksson (1999, 2002), and sections 8.3 and 8.4 in this volume. The application of sequence stratigraphy to the rock record becomes increasingly difficult with older strata because the analysis of stratigraphic surfaces and facies, as well as the age determinations, encounter practical difficulties when applied to such successions. Facies analysis and palaeoenvironmental reconstructions for Precambrian depositional systems are often limited by poor preservation, post-depositional tectonics, diagenetic transformations, and metamorphism. The constraint of Precambrian rocks' ages, based essen-

684

Chapter 8: Sequence Stratigraphy and the Precambrian

Sequencel IEve~ Dep~176 Depositional Depositional Genetic SequenceII SequenceIII SequenceIV Sequence

T-R Sequence

HST

early HST

HST

HST

RST

TST

TST

TST

TST

TST

late LST (wedge)

LST

LST

late LST (wedge)

early LST (fan)

late HST (fan)

FSST

early LST (fan)

HST

early HST (wedge)

HST

HST

end of transgression end of regression end of base level fall onset of base level fall

"•'• ......

sequence boundary within systems tract surface

]

~k

base level fall

~

/

I

RST

end of transgression

Time" regression

i

end of base level fall

Fig. 8.1-2. Position of sequence boundaries, as well as the subdivision into systems tracts, for the sequence models currently in use (from Catuneanu, 2002). Abbreviations: LST--lowstand systems tract; TST--transgressive systems tract; HSTmhighstand systems tract; FSST--falling stage systems tract; RST--regressive systems tract; T-R--transgressive-regressive.

tially on radiometric dating, is also less precise relative to younger sequences, with error margins of at least 1 My. For this reason, no time control can be provided for the correlation of high frequency Precambrian sequences with durations less than 10~ My. As a result, sequence stratigraphic interpretations of early Precambrian successions are not common in the literature. Given the existing limitations, most studies that have been done are low resolution and preliminary interpretations, and generally deal with the basic large-scale subdivision of sedimentary basin-fills. An example of such a study is that of Catuneanu and Eriksson (1999), which provides the fundamental breakdown of the c. 2.6-2.1 Ga Transvaal Supergroup, Kaapvaal craton (first-order sequence) into second-order sequences. This type of study provides the framework for the more detailed sequence stratigraphic analysis of individual large-scale sequences. Where the geometry, sedimentary facies and facies relationships in Precambrian successions are well constrained, thus enabling reliable depositional models to be constructed, it is indeed possible to apply sequence stratigraphy of higher detail, even in the near absence of any geochronological constraints (e.g., Catuneanu and Biddulph, 2001; Catuneanu and Eriksson, 2002) (sections 8.3 and 8.4).

8.2. Conceptsof Sequence Stratigraphy

8.2.

685

CONCEPTS OF SEQUENCE STRATIGRAPHY

O. CATUNEANU, A.E EMBRY AND EG. ERIKSSON

Sequence Stratigraphic Models Fundamentals of sequence stratigraphy At the core of sequence stratigraphic analysis is the sedimentary response to changes in accommodation and supply (Swift and Thorne, 1991). The concept of accommodation (Jervey, 1988) defines the space available for sediments to accumulate. This space can be created or destroyed by fluctuations in base level, and is gradually consumed by sedimentation (Fig. 8.2-1). The interplay of sedimentation and shifting base level generates changes in depositional trends in the rock record, and it is the correlation of these changes that defines the object of study of sequence stratigraphy. Four main events associated with changes in depositional trends are recorded during a complete cycle of base level shifts (Fig. 8.1-2): 1. Onset of base level fallmthis is generally accompanied by a change from sedimentation to erosion/bypass in the fluvial to shallow marine environments; 2. End of base level fallmthis marks a change from degradation to aggradation in the fluvial to shallow marine environments; 3. End of regression~this marks the turnaround point from shoreline regression to subsequent transgression; 4. End of transgression~this marks a change in the direction of shoreline shift from transgression to subsequent regression. These four events, which are in turn controlled by subsidence, eustasy and sediment supply, mark the timing of the formation of all sequence stratigraphic surfaces, as outlined in the next section of this paper. Transgressions, as well as two types of regressions may be defined as a function of the ratio between the rates of base level changes and the sedimentation rates at the shoreline (Posamentier et al., 1992; Fig. 8.2-2). The stratal geometries associated with these basic types of shoreline shifts are presented in Figure 8.2-3. Transgressions occur when accommodation is created more rapidly than it is consumed by sedimentation, i.e., the rates of base level rise outpace the sedimentation rates at the shoreline (Fig. 8.2-2). This results in a retrogradation of facies. The scour surface cut by waves during the shoreline transgression is onlapped by the aggrading and retrograding shoreface deposits (Fig. 8.2-3). Forced regressions occur during stages of base level fall, when the shoreline is forced to regress by the falling base level irrespective of the sediment supply (Fig. 8.2-2). This triggers erosional processes in both the non-marine and shallow marine environments adjacent to the coastline. Fluvial incision is accompanied by the progradation of offlapping shoreface deposits (Fig. 8.2-3). Normal regressions occur in the early and late stages of base level rise, when the sedimentation rates outpace the low rates of base level rise at the shoreline (Fig. 8.2-2). In this lhe Ptecambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

686

Chapter 8: Sequence Stratigraphy and the Precambrian

I Transgressions and Regressions shoreline shifts

,. Baselevel

Iv

I Relative ,~ sealevel

changes base level relative to DATUM

~k,

changes

Sealevel changesl

I

creation and destruction

J Accommodation J

sea level relative to DATUM

base level relative to depositional surface

t I

l

consumption

sea level

DATUM

v relative to the centre of Earth

depositional environment

J SedimentationJ depositional surface relative to DATUM

Fig. 8.2-1. Controls on base level changes, transgressions, and regressions (from Catuneanu, 2002). Note that the relative sea level changes account for the combined effects of eustasy and tectonics. Sediment compaction is included under "Tectonics", as it has the same effect on accommodation as the tectonic subsidence. As the base level is offset relative to the sea level due to the energy of the environment (waves, currents), the base level changes build on, but are not equal to, the relative sea level changes. Sea level changes, relative sea level changes, and base level changes are independent of sedimentation. The interplay between base level changes (combined effect of eustasy, tectonics, compaction, and environmental energy) and sedimentation controls the transgressive or regressive shifts of the shoreline. The DATUM, as defined by Posamentier et al. (1988), is a reference horizon taken beneath the sea floor to monitor the magnitude of vertical tectonics relative to the centre of the Earth. case, the newly created accommodation is totally consumed by sedimentation, aggradation is accompanied by sediment bypass, and a progradation of facies occurs (Fig. 8.2-3). Note that both transgressions and normal regressions may occur during base level rise, as a function of the balance between the rates at which accommodation is created and consumed (Fig. 8.2-2). This would make the transgressive stages shorter in time than the regressive stages (normal and forced), given a symmetrical curve of base level changes. The succession of transgressive and regressive shifts illustrated in Fig. 8.2-2 represents the most complete scenario of stratigraphic cyclicity. In practice, simplified versions of stratigraphic cyclicity may also be encountered, such as: (i) repetitive successions of transgressive and normal regressive facies, where continuous base level rise in the basin outpaces and is outpaced by sedimentation in a cyclic manner; (ii) repetitive successions of forced and normal regressions, where the high sediment input consistently outpaces the rates of base level rise (hence, no transgressions).

8.2. Concepts of Sequence Stratigraphy

687

Fig. 8.2-2. Concepts of transgression, normal regression, and forced regression, as defined by the interplay between base level changes and sedimentation (from Catuneanu, 2002). The sine curve at the top shows the magnitude of base level changes through time. The thicker portions on this curve indicate early and late stages of base level rise, when the rates of base level rise (increasing from zero and decreasing to zero, respectively) are outpaced by the sedimentation rates. The sine curve below shows the rates of base level changes. Note that the rates of base level changes are zero at the end of base level rise and base level fall stages (the change from rise to fall and from fall to rise requires the motion to cease). The rates of base level changes are the highest at the inflection points on the top curve. For simplicity, the sedimentation rates are kept constant during the shown base level fluctuations. Transgressions occur when the rates of base level rise outpace the sedimentation rates.

688

Chapter 8: Sequence Stratigraphy and the Precambrian

Fig. 8.2-3. Shoreline trajectories--normal regressions, forced regressions and transgressions (from Catuneanu, 2002). Abbreviations: FR = forced regression; NR = normal regression.

Sequence stratigraphic surfaces Surfaces that can serve, at least in part, as systems tract boundaries, are sequence stratigraphic surfaces. Systems tracts are linkages of contemporaneous depositional systems, forming the subdivisions of a sequence (Brown and Fisher, 1977). The nomenclature and definition of systems tracts differ among the various sequence models (Figs. 8.1-1 and 8.1-2), but invariably, the timing of each systems tract boundary corresponds to one of the four main events of the base level cycle (Fig. 8.1-2). The timing and diagnostic features of the seven surfaces of sequence stratigraphy are presented in Figures 8.2-4 and 8.2-5. These surfaces are not equally easy to identify in outcrop and subsurface, nor equally useful as time markers in a chronostratigraphic framework. Nevertheless, irrespective of their physical and temporal attributes, each surface may be defined as a distinct stratigraphic contact that marks a specific event of the base level cycle. A succinct presentation of the seven surfaces follows below.

Subaerial unconformity. The importance of subaerial nonconformities as sequencebounding surfaces was emphasised by Sloss et al. (1949). The subaerial unconformity is a surface of erosion or non-deposition created during base level fall by subaerial processes such as fluvial incision, wind degradation, sediment bypass, or pedogenesis. It gradually extends basinwards during the forced regression of the shoreline and reaches its maximum extent at the end of the forced regression (Helland-Hansen and Martinsen, 1996: "seawards, the subaerial unconformity extends to the location of the shoreline at the end of fall"). Criteria for the recognition of subaerial nonconformities in the field have been reviewed by Shanmugam (1988). The subaerial unconformity has a marine correlative con-

689

8.2. Concepts of Sequence Stratigraphy

Base level

Surfaces

Events

. . . . Onset of fall . . . . . ~1- Basal surface of forced regression -End of transgression - ~1- Maximum flooding surface Ravinement surface --End of regression - - .~1- Maximum regressive surface ~1- Correlative conformity* End of fall

-f

}

(1)

t~ m

Subaerial unconformity and Regressive surface of marine erosion

<E

LL

.... v

Rise

I

Fall

Onset of fall -

sensu

~1- Basal surface of forced regression **

Hunt and Tucker (1992)

= correlative conformity sensu Posamentier et al. (1988)

Fig. 8.2-4. Timing of sequence stratigraphic surfaces relative to the main events of the base level cycle (modified from Embry and Catuneanu, 2002). (-A)--negative accommodation.

Surface

Nature of contact

Strata below

Strata above

Temporal attributes

Subaerial unconformity

Scoured or top of palaeosol

Variable (where marine, c-u)

Non-marine

Diachronous hiatus

Correlative conformity

Conformable

Marine, c-u

Marine (c-u on shelf)

Low diachroneity

Regressive surface of marine erosion

Scoured

Shelf, c-u

Shoreface, c-u

High diachroneity

Basal surface of forced regression

Conformable or scoured

Marine, variable (c-u on shelf)

Marine, c-u

Low diachroneity

Maximum regressive surface

Conformable

Variable (where marine, c-u)

Variable (where marine, f-u)

Low diachroneity

Maximum flooding surface

Conformable or scoured

Variable (where marine, f-u)

Variable (where marine, c-u)

Low diachroneity

Ravinement surface

Scoured

Variable (where marine, c-u)

Marine, f-u

High diachroneity

Fig. 8.2-5. Diagnostic features of sequence stratigraphic surfaces (modified from Embry and Catuneanu, 2002). *sensu Hunt and Tucker (1992); **correlative conformity sensu Posamentier et al. (1988). Abbreviations: c-u = coarsening-upwards; f-u = fining-upwards.

690

Chapter 8: Sequence Stratigraphy and the Precambrian

formity whose timing corresponds to the end of base level fall at the shoreline (sensu Hunt and Tucker, 1992; Fig. 8.2-4). Forced regressions require the fluvial systems to adjust to new (lower) graded profiles. A small base level fall at the shoreline may be accommodated by changes in channel sinuosity, roughness and width, with only minor incision (Schumm, 1993; Ethridge et al., 2001). The subaerial unconformity generated by such unincised fluvial systems is mainly related to the process of sediment bypass (Posamentier, 2001 ). A larger base level fall at the shoreline, such as the lowering of the base level below a major topographic break (e.g., the shelf break) results in fluvial downcutting and the formation of incised valleys (Schumm, 1993; Ethridge et al., 2001; Posamentier, 2001). The interfluve areas are generally subject to sediment starvation and soil development. The subaerial unconformity can thus be traced at the top of palaeosol horizons that are correlative to the nonconformities generated in the channel subenvironment (Wright and Marriott, 1993). The subaerial unconformity may be placed at the top of any type of depositional system (fluvial, coastal or marine), but it is always overlain by non-marine deposits. The stratigraphic hiatus associated with the subaerial unconformity is variable, due to differential fluvial incision and the gradual expansion of subaerial erosion in a basinwards direction during the stage of base level fall. A synonymous term for the subaerial unconformity is the regressive surface of fluvial erosion (Plint and Nummedal, 2000).

Correlative conformity. The correlative conformity forms within the marine environment at the end of base level fall at the shoreline (sensu Hunt and Tucker, 1992; Fig. 8.2-4). This approximates the palaeo-sea floor at the end of forced regression, which correlates with the seawards termination of the subaerial unconformity. The correlative conformity was also defined as the palaeo-sea floor at the onset of forced regression (Posamentier et al., 1988), but this choice was criticised because it allows the subaerial unconformity, and the correlative conformity, to be both intercepted in the same vertical section within the area of forced regression (Hunt and Tucker, 1992). In this case, the correlative conformity (sensu Posamentier et al., 1988) does not correlate with the seawards termination of the subaerial unconformity. The correlative conformity turned out to be a problem surface in sequence stratigraphy, surrounded by controversies regarding its timing and physical attributes. The main problem relates to the difficulty of recognising it in most outcrop sections, cores, or wireline logs, although at the larger scale of seismic data one can infer its approximate position as the clinoform that correlates with the basinwards termination of the subaerial unconformity. The latter method of mapping the correlative conformity is limited by the relatively low seismic resolution, which makes it possible that a number of discrete clinoforms may be amalgamated as one seismic horizon. The shallow marine portion of the correlative conformity develops within a conformable prograding package (coarsening-upwards trends below and above), lacking lithofacies and grading contrasts. In the deep marine environment, the correlative conformity is proposed to be mapped at the top of the prograding and coarsening-upwards submarine fan complex (the "basin floor component" of Hunt and Tucker, 1992). The overlying gravity flow

8.2. Concepts of Sequence Stratigraphy

691

deposits tend to display a fining-upwards trend, due to the gradual cut-off of sediment supply during rising base level (Posamentier and Walker, 2002). Beyond these models, the mapping of the end-of-fall surface within deep marine facies is in fact much more difficult because the manifestation of gravity flows, and the associated vertical profiles, depend on a multitude of factors, some of which are independent of base level changes. In addition to this, the idea of correlative changes along strike from coarsening- to fining-upwards trends is based on the assumption that there is a uniform linear source of sediment to the outer shelf, slope, and basin floor. This is generally untrue in most clastic basins, as there is rarely enough clastic sediment influx to the basin to affect deposition in more than a small region at any one time (Frazier, 1974). Considering the natural shifts in sediment accumulation loci, there is little likelihood that changes from coarsening- to fining-upwards are synchronous along strike. The correlative conformity is implied to be a time line, i.e., "the time surface that is correlative with the 'collapsed' unconformity" (Posamentier and Allen, 1999). At the same time, the correlative conformity is also defined in relation to general stacking patterns, at "a change from rapidly prograding parasequences to aggradational parasequences" (Haq, 1991), or at the top of submarine fan deposits (Hunt and Tucker, 1992). The latter definitions imply a diachronous correlative conformity, younging basinwards, with a rate that matches the rate of offshore sediment transport (Fig. 8.2-5; Catuneanu et al., 1998; Catuneanu, 2002).

Basal surface offorced regression. The basal surface of forced regression was introduced by Hunt and Tucker (1992) to define the base of all deposits that accumulate in the marine environment during the forced regression of the shoreline. This corresponds to the correlative conformity of Posamentier et al. (1988), and it approximates the palaeo-sea floor at the onset of base level fall at the shoreline (Fig. 8.2-4). A low diachroneity is, however, recorded in relation to the rates of offshore sediment transport (Catuneanu, 2002). In shallow marine successions, the basal surface of forced regression may be conformable, in which case it poses the same recognition problems as the correlative conformity (coarsening-upwards strata below and above), or it may be reworked by the regressive surface of marine erosion. Where conformable, the basal surface of forced regression may in theory be mapped on seismic lines as the youngest normal regressive clinoform that underlies offlapping forced regressive shoreface lobes. The main pitfall of this approach is that the subaerial unconformity often removes the earliest offlapping sandstone strata, so one cannot always determine when the offlapping deposits actually begin on the seismic section. In the deep marine environment, the basal surface of forced regression is taken at the base of the prograding submarine fan complex (Hunt and Tucker, 1992), as the scour cut by the earliest gravity flows associated with the forced regression of the shoreline. In this case, the basal surface of forced regression separates pelagic sediments below (no grading) from gravity flow deposits above (overall coarsening-upwards). The pitfall of this model is that the first submarine fan strata may not necessarily coincide with the start of base level fall, but may in fact begin at any time during fall, depending on physiography and sediment supply.

692

Chapter 8: Sequence Stratigraphy and the Precambrian

Regressive surface of marine erosion. The regressive surface of marine erosion is a scour cut by waves in the lower shoreface during the forced regression of the shoreline, as the shoreface attempts to preserve its concave-up profile that is in equilibrium with the wave energy (Bruun, 1962; Plint, 1988; Dominguez and Wanless, 1991; Plint and Nummedal, 2000). This surface (the "marine scour" associated with forced regressions in Fig. 8.2-3) underlies sharp-based shoreface deposits (Plint, 1988), and may be separated from the basal surface of forced regression by forced regressive shelf sediments. The landwards portion of the regressive surface of marine erosion is likely to rework the basal surface of forced regression, in which case it becomes a systems tract boundary (figures 21 and 26 in Catuneanu, 2002). Both the underlying and overlying deposits are coarsening-upwards, as being part of a regressive succession. The regressive surface of marine erosion is highly diachronous, with the rate of shoreline forced regression. The formation of the regressive surface of marine erosion requires a shallow gradient of the sea floor, smaller than the average gradient of the shoreface profile (c. 0.3~ This is often the case in shelf settings, where the average gradient of the sea floor is about 0.03 ~. In contrast, slope settings have a steeper sea floor topography (c. 3 ~ relative to what is required by the shoreface to be in equilibrium with the wave energy, and hence no scouring is generated in the lower shoreface during forced regressions. These steep sea floor slopes are prograded by Gilbert-type deltas whose delta front facies are not sharp-based (sensu Plint, 1988). A synonymous term for the regressive surface of marine erosion is the regressive ravinement surface (Galloway, 2001).

Maximum regressive surface.

The maximum regressive surface (Helland-Hansen and Martinsen, 1996) is defined relative to the transgressive-regressive curve, marking the point between regression and subsequent transgression. Hence, this surface separates prograding strata below from retrograding strata above. The change from progradational to retrogradational stacking patterns takes place during the base level rise at the shoreline, when the rates of base level rise start outpacing the sedimentation rates (Fig. 8.2-2). The maximum regressive surface is generally conformable, although the possibility of scouring associated with the change in the direction of shoreline shift at the onset of transgression is not excluded (Loutit et al., 1988; Galloway, 1989). The maximum regressive surface is also known as the transgressive surface (Posamentier and Vail, 1988), initial transgressive surface (Nummedal et al., 1993), or conformable transgressive surface (Embry, 1995). The maximum regressive surface has a low diachroneity along dip that reflects the rates of sediment transport (Catuneanu, 2002). The diachroneity rates may increase substantially along strike, due to the variability in sediment supply (Catuneanu et al., 1998). In a shallow marine succession, the maximum regressive surface is easy to recognise at the top of coarsening-upwards (regressive) deposits. In coastal settings, the maximum regressive surface underlies the earliest estuarine deposits (figures 18 and 37 in Catuneanu, 2002). The contact between estuarine and underlying fluvial facies diverges from the maximum regressive surface beyond the initial length of the estuary at the onset of transgression, becoming time-transgressive in an upstream direction. The extension of the maximum regressive surface into the fluvial part of the basin is much more difficult to pinpoint, but at a

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regional scale it is argued to correspond with an abrupt decrease in fluvial energy, i.e., with a change from amalgamated braided channel-fills to overlying meandering systems (Kerr et al., 1999; Ye and Kerr, 2000). This shift in fluvial styles is attributed to the formation of the low energy estuarine system at the beginning of transgression, which would induce a lowering in fluvial energy upstream. However, much work is still needed to document properly the physical attributes of the non-marine maximum regressive surface.

Maximumflooding surface. The maximum flooding surface (Frazier, 1974; Posamentier et al., 1988; Van Wagoner et al., 1988; Galloway, 1989) is also defined relative to the transgressive-regressive curve, marking the end of shoreline transgression. Hence, this surface separates retrograding strata below from prograding strata above. The presence of prograding strata above identifies the maximum flooding surface as a downlap surface on seismic data. The change from retrogradational to overlying progradational stacking patterns takes place during continued base level rise at the shoreline, when the sedimentation rates start to outpace the rates of base level rise (Fig. 8.2-2). The maximum flooding surface is generally conformable, excepting for the outer shelf and upper slope regions where the lack of sediment supply may leave the sea floor exposed to erosional processes (Galloway, 1989). The maximum flooding surface is also known as the maximum transgressive surface (Helland-Hansen and Martinsen, 1996) or final transgressive surface (Nummedal et al., 1993). The maximum flooding surface has a low diachroneity along dip that reflects the rates of sediment transport (Catuneanu, 2002). As in the case of the maximum regressive surface, the diachroneity rates may increase substantially along strike due to the variability in sediment supply (Catuneanu et al., 1998). In a marine succession, the maximum flooding surface is placed at the top of finingupwards (transgressive) deposits. In an offshore direction, the transgressive deposits may be reduced to a condensed section, or may even be missing. In the latter situation, the maximum flooding surface will be superimposed on and rework the maximum regressive surface. In coastal settings, the maximum flooding surface is placed at the top of the youngest estuarine facies. Criteria for the recognition of the maximum flooding surface in the fluvial portion of the basin have been provided by Shanley et al. (1992), mainly based on the presence of tidal influences in fluvial sandstones. The position of this surface may also be indicated b y a n abrupt increase in fluvial energy, from meandering to overlying braided fluvial systems (Shanley et al., 1992), or by regionally extensive coal seams (Hamilton and Tadros, 1994). Tidal influences in fluvial strata may occur within a few tens of kilometres from the coeval shoreline (Shanley et al., 1992). Farther inland, the maximum flooding surface corresponds to the highest level of the watertable relative to the land surface, which, given a low sediment input and the right climatic conditions, may offer good conditions for peat accumulation at a regional scale.

Ravinement surface. The ravinement surface is a scour cut by waves in the upper shoreface during shoreline transgression (Bruun, 1962; Swift et al., 1972; Swift, 1975; Dominguez and Wanless, 1991; the "wave scour" in Fig. 8.2-3). This erosion may remove as much as 10-20 m of substrate (Demarest and Kraft, 1987), as a function of the wind regime and related wave energy in each particular region. The ravinement surface

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is onlapped during the retrogradational shift of facies by transgressive (fining-upwards) shoreface deposits (coastal onlap), and it may overlie any type of depositional system (fluvial, coastal, or marine). The ravinement surface is highly diachronous, with the rate of shoreline transgression. In a vertical profile that preserves the entire succession of facies, the ravinement surface separates coastal strata below (foreshore and backshore facies in an open shoreline setting, or estuarine facies in a river mouth setting) from shoreface and shelf deposits above. Where the transgressive coastal deposits are not preserved, the ravinement surface may rework the underlying regressive strata and the subaerial unconformity (Embry, 1995). In the latter case, the ravinement surface becomes part of the sequence boundary. Synonymous terms for the ravinement surface include the transgressive ravinement surface (Galloway, 2001), wave-ravinement surface (Swift, 1975), shoreface ravinement (Embry, 1995), and transgressive surface of erosion (Posamentier and Vail, 1988). Within-trend facies contacts. In addition to the seven sequence stratigraphic surfaces described above, facies contacts associated with a strong physical expression may also be recognised within the various systems tracts. Such lithological discontinuities may be caused by abrupt changes in sediment supply during continuous transgressive or regressive trends, and are surfaces of lithostratigraphy or allostratigraphy. For example, some types of "flooding surfaces" qualify as within-trend facies contacts, and these are not proper sequence stratigraphic surfaces as they do not serve as systems tract boundaries (see Catuneanu, 2002, for a discussion). In a sequence stratigraphic approach, facies contacts need to be dealt with only after the framework of sequence stratigraphic surfaces has been constructed. Systems tracts and sequences

Systems tracts and sequences are packages of strata bounded by specific combinations of sequence stratigraphic surfaces. Surfaces assigned as systems tract or sequence boundaries vary with the sequence model that is being employed (Figs. 8.1-1 and 8.1-2), which adds an unnecessary complication to a science that otherwise is very logical and straightforwards. Four out of five models use the subaerial unconformity as the sequence boundary for the non-marine side of the basin (the three types of depositional sequence, and the T-R sequence). These models use, however, different surfaces for the marine portion of the sequence boundary, i.e., the basal surface of forced regression (correlative conformity sensu Posamentier et al., 1988; depositional sequence II), the correlative conformity (sensu Hunt and Tucker, 1992; depositional sequence III and IV), and the maximum regressive surface (T-R sequence). The fifth model (genetic sequence) uses the maximum flooding surface as the sequence boundary for both marine and non-marine portions of the basin (Fig. 8.1-2). For a more comprehensive discussion of the five sequence models, see Catuneanu (2002). Similar to the concept of "sequence", which varies in meaning from one model to another, systems tracts may also change significance among the various models according to what surfaces are used to define their boundaries (Fig. 8.1-2). In this context, the systems tract terminology becomes somewhat meaningless, and a source of confusion. It is

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Fig. 8.2-6. Depositional processes and products of late rise normal regression. The deposits of this stage ("highstand prism") overlie the maximum flooding surface. The bulk of the highstand prism includes fluvial, coastal and shoreface deposits. The shelf and deep marine environments receive mainly fine-grained hemipelagic and pelagic sediments. far more important to identify correctly the types of surfaces in the rock record, and worry less about how to name the packages of strata between them. For this reason, the following discussion of depositional processes and products of a full cycle of base level changes is centred around the types of shoreline shifts (normal regressions, forced regressions, and transgressions) rather than specific systems tracts. Normal regression 1 (late base level rise; Fig. 8.2-6)mits depositional products are often referred to as the highstand systems tract or the highstand wedge of a larger systems tract (Fig. 8.1-2). The shoreline trajectory is defined by a combination of aggradation and progradation processes (Fig. 8.2-3), which result in the formation of a "highstand prism" of fluvial, coastal and shoreface deposits (Fig. 8.2-6). Deltas are far from the shelf edge (as they follow a transgressive episode) and develop diagnostic topset packages of delta plain strata. Strandplains may develop along open shorelines as a result of beach progradation under low rates of base level rise. Shelf edge stability, coupled with a lack of sediment supply to the outer shelf-upper slope area, results in a paucity of gravity flows in the deeper marine environment. Early forced regression (early base level fall; Fig. 8.2-7)mits depositional products are variously included in a lowstand fan, highstand fan, falling stage systems tract or regressive systems tract (Fig. 8.1-2). The shoreline trajectory is defined by progradation and offlap, accompanied by fluvial erosion or bypass upstream (Fig. 8.2-3). The lack of aggradation in the delta plain prevents the formation of a deltaic topset. Elongated beach-upper shoreface

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Fig. 8.2-7. Depositional processes and products of early fall forced regression. Most of the sand that accumulates during this stage is captured within detached and offlapping palaeo-shoreline systems. A significant amount of finer-grained sediment starts to accumulate in the deep marine environment as mudflow deposits. Two sequence stratigraphic surfaces form during this stage: the subaerial unconformity, which gradually expands basinwards as the shoreline regresses; and the regressive surface of marine erosion (RSME) cut by waves in the lower shoreface. The basal surface of forced regression is taken at the base of all forced regressive strata, including the early fall mudflow deposits. In places, this surface may be reworked by the RSME. sand bodies may be abandoned on the subaerially exposed shelf as the base level falls. These palaeo-shoreline sands are generally thin (range of metres) and isolated (i.e., separated by gaps). The degree of detachment depends on the interplay of sediment supply and the rates of base level fall (Posamentier and Morris, 2000). During early fall, the shoreline is still far from the shelf edge, so no riverborn sand is delivered directly to the continental slope. However, lowering of the storm wave base causes instability on the outer shelf, which triggers gravity flow processes into the deep marine environment. These gravity flows mainly include the fine-grained sediment accumulated on the outer shelf-upper slope area during the previous late rise normal regression. Late forced regression (late base level fall; Fig. 8.2-8)--as in the case of the early forced regressive deposits, the products of this stage are variously included in a lowstand fan, highstand fan, falling stage systems tract or regressive systems tract (Fig. 8.1-2). The early forced regressive palaeo-shoreline sands are subject to fluvial and wind degradation, and loose their original linear geometry. As the shoreline approaches the shelf edge, the fluvial sediment is delivered straight to the continental slope, causing major gravity flow events. Additional sediment supply is generated by processes of fluvial incision upstream. The

8.2. Concepts of Sequence Stratigraphy

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Fig. 8.2-8. Depositional processes and products of late fall forced regression. The sediment mass balance changes in favour of the deep sea submarine fans, which capture most of the sand. The subaerial unconformity keeps forming and expanding basinwards until the end of base level fall. Once the shoreline falls below the shelf edge, the regressive surface of marine erosion stops forming, as the sea floor gradient of the slope is steeper than that required by the shoreface profile to be in equilibrium with the wave energy. At the end of base level fall, the top of forced regressive deposits is marked by the correlative conformity (sensu Hunt and Tucker, 1992). lack of accommodation for the fluvial and shoreline systems explains the large volume of turbidites which accumulates during this time in the deep marine environment. As the sediment entry points gradually get closer to the shelf edge during the falling stage, the overall vertical profile of the forced regressive submarine fans is coarsening-upwards, with a transition from mudflows to sandy turbidites (Figs. 8.2-7 and 8.2-8). Normal regression 2 (early base level rise; Fig. 8.2-9)nthe depositional products of this stage are generally labelled as lowstand systems tract or late lowstand wedge deposits (Fig. 8.1-2). Rising base level provides accommodation for both fluvial and coastal aggradation, which results in a net decrease in the volume of deep marine gravity flows. The trapping of sand within the fluvial and shoreline systems also results in a lowering of the sand/mud ratio in the submarine fans (Posamentier and Walker, 2002). Shelf edge deltas continue to prograde on the upper slope, with the development of a topset in response to delta plain aggradation. Increased elevation at the shoreline triggers fluvial aggradation, which starts from the shoreline and gradually expands upstream (fluvial onlap). Early transgression (rapid base level rise; Fig. 8.2-10)rathe depositional products of this stage are part of the transgressive systems tract (Fig. 8.1-2). The shoreline trajectory is defined by a combination of retrogradation and shoreface aggradation (Fig. 8.2-3). Rapid rates of base level rise result in most of the terrigenous sediment being trapped in retro-

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Chapter 8: Sequence Stratigraphy and the Precambrian

Fig. 8.2-9. Depositional processes and products of early rise normal regression. The sediment of this stage is more evenly distributed between the fluvial, coastal, and deep marine systems. Sand is present in amalgamated fluvial channel-fills, beach and delta front systems, as well as in submarine fans. The "lowstand prism" gradually expands landwards via fluvial aggradation and onlap. The top of all early rise normal regressive deposits is marked by the maximum regressive surface.

grading fluvial, estuarine, and barrier beach systems. This transgressive prism gradually expands and onlaps in a landwards direction. Erosion at the shelf edge by wave ravinement processes continues to supply sandy sediment for deep marine gravity flows, although a significant amount of sand is also trapped within backstepping shoreline systems and onlapping healing phase deposits (Fig. 8.2-10; Posamentier and Allen, 1993). Late transgression (rapid base level rise; Fig. 8.2-11 )--the depositional products of this stage are also part of the transgressive systems tract (Fig. 8.1-2). A significant portion of the shelf is now submerged, and tidal currents may lead to the formation of sandy shelf ridges oriented normal to the shoreline (Posamentier, 2002). Elsewhere, the shelf is generally subject to sediment starvation and condensed sections are likely to form (Loutit et al., 1988). Rapid increase in water depth leads to shelf edge instability, which results in the manifestation of gravity flows (Galloway, 1989). Such flows are mud-rich, involving finegrained outer shelf sediments that accumulated far from the sediment entry points. Generally speaking, the sand/mud ratio of the gravity flow deposits accumulated during rising base level (normal regression 2 to transgression) records an overall decrease, from sandy turbidites to mudflows (Figs. 8.2-9-8.2-11; Posamentier and Walker, 2002). This finingupwards trend of the rising stage in the deep marine environment is completed by the accumulation of pelagic/hemipelagic sediments of the late rise normal regression at the top of the submarine fans (Fig. 8.2-6).

8.2. Conceptsof Sequence Stratigraphy

699

Fig. 8.2-10. Depositional processes and products of early transgression. Wave ravinement processes erode the underlying normal regressive shelf edge delta and open shoreline systems, continuing to supply sand for the deep marine turbidity flows. Rapid rates of base level rise trigger a retrogradational shift of facies on the shelf', where most of the riverborn sediment is now trapped in fluvial, estuarine and backstepping shoreline systems. The ravinement surface continues to form during the entire duration of transgression. Wave scouring in the upper shoreface is accompanied by the sedimentation of "healing phase" deposits in the lower shoreface, which partly infill, or heal over, the palaeo-bathymetric profile (Fig. 8.2-3; Posamentier and Allen, 1993). Healing phase deposits onlap the ravinement surface during transgression (coastal onlap), and also trap a significant proportion of the sediment supplied by ravinement erosion. The river mouth settings may become estuaries (shown in the diagram) or backstepping bayhead deltas, depending on the relative rates of transgression between the river mouth and the open shoreline settings.

Sequence Hierarchy Introduction The concept of a hierarchy for sequence boundaries and their associated sequences is not one that has received much attention, but it is one that is crucial for the practical application of sequence stratigraphy. The need for a hierarchy is readily apparent when one considers that there are commonly numerous sequence boundaries, each of which is expressed as either an unconformity or a correlative conformity (maximum regressive surface in the T-R sequence model), in a sedimentary succession. A sequence can be delineated theoretically between any two of the recognised sequence boundaries. Thus, if one recognised 25 sequence boundaries in a succession, 300 different sequences could potentially be delineated. Such a system of sequence delineation borders on the ridiculous but can be avoided

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Fig. 8.2-11. Depositional processes and products of late transgression. Most of the sediment is now trapped in the fluvial to shallow marine transgressive prism, which includes fluvial, estuarine, shoreline, and lower shoreface deposits. Additional sand is incorporated within the shelf ridges generated by tidal currents. Instability at the shelf edge generates mudflows in the deep marine environment. The top of all transgressive deposits is marked by the maximum flooding surface. Where the transgressive deposits are missing (e.g., in the outer shelf-upper slope areas subject to non-deposition and erosion), the maximum flooding surface reworks the maximum regressive surface. The river mouth settings may become estuaries (shown in the diagram) or backstepping bayhead deltas, depending on the relative rates of transgression between the river mouth and the open shoreline settings.

only by the establishment of a sequence boundary hierarchy that contains distinctive groups of boundaries. Two very different methodologies for establishing a hierarchy of sequence boundaries have been proposed. One is based on differing boundary frequency and the other on the relative magnitude of the boundary. There are also two methods for naming the various hierarchical classes of sequence boundaries. One is based on a specific name for each class (e.g., megasequence boundary) and one is based on a numerical ordering (e.g., second-order sequence boundary). Below we discuss these various options and present our preferred methodology for establishing and naming hierarchical classes of sequence boundaries.

Establishing a hierarchy Vail et al. (1977) first discussed a sequence boundary hierarchy in terms of "orders of cycles superimposed on the sea-level curve". Vail et al. (1977) assumed a priori that three distinct orders of sea level variation existed with the largest changes occurring every

8.2. Concepts of Sequence Stratigraphy

701

200-300 million years (first-order), intermediate changes occurring every 10-80 million years and smaller ones occurring every 1-10 million years. Thus sequence boundary hierarchy was tied to boundary frequency although the concept of relative magnitude was also included. This culminated in a publication by Vail et al. (1991) in which six orders of boundaries were defined solely on boundary frequency. The six orders and their characteristic boundary frequencies in this hierarchical scheme are: first-order, 50 My; second-order, 3-50 My; third-order, 0.5-3 My; fourth-order, 0.08-0.5 My; fifth-order, 0.03-0.08 My; and sixth-order, 0.01-0.03 My. Notably Krapez (1996, 1997) has adopted this methodology for the Precambrian. He states "There are no physical criteria with which to judge the rank of a sequence boundary. Therefore, sequence rank is assessed from interpretations of the origin of the strata contained between the key surfaces, and of the period of the processes that formed these strata" (Krapez, 1997, p. 2). The use of boundary frequency for the establishment of a sequence hierarchy is problematic and basically circular. Statistical surveys suggest that there is no evidence for a hierarchy in the actual rock record (e.g., Drummond and Wilkinson, 1996). The frequency of occurrence of a given sequence boundary class (e.g., third-order boundaries) is an interpretation which can be reached only when a clear characterisation of that sequence class has been established. In other words what are the defining characteristics of a third-order boundary such that it can be determined that such boundaries indeed have a 0.5-3 My frequency? Unfortunately neither Exxon scientists nor Krapez (1996, 1997) have provided any clear information on what attributes characterise each of their various classes of sequence boundaries and how the various classes can be distinguished from one another. In addition to this, the dynamics and rates of plate tectonic processes changed from the Precambrian to the Phanerozoic (see also sections 3.2, 3.6 and 3.11), which means that the temporal duration of tectonics-driven stratigraphic cycles is irrelevant in establishing a sequence hierarchy. An alternative practical methodology is to establish a hierarchy of boundaries based on the interpreted relative magnitudes of the base level changes that generated the boundaries in the first place (Embry, 1995). The problem at hand is to find objective scientific criteria which reflect the magnitude of the base level changes represented by a given sequence boundary. Contrary to Krapez's (1996, 1997) claim that there are no physical criteria available for establishing a sequence hierarchy, we think that such criteria can be identified scientifically. It seems reasonable to expect that a base level change of 500 m would result in a sequence boundary that has different attributes to a sequence boundary that was generated by a base level change of 10 m or less. The attributes of a sequence boundary we have chosen to use to estimate the amount of base level change that generated the boundary in the first place, include: (1) the areal extent over which the sequence boundary can be recognised; (2) the areal extent of the unconformable portion of the boundary; (3) the magnitude of base level fall as determined by the thickness of section eroded beneath the unconformable portion of the boundary;

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Chapter 8: Sequence Stratigraphy and the Precambrian

(4) the magnitude of subsequent base level rise as derived from interpretations of the water depth of strata associated with the maximum flooding surface above the unconformable portion of the boundary; (5) the degree of change of the depositional regime across the boundary; (6) the degree of change of the tectonic setting across the boundary. Each of these attributes has to be assessed for each boundary and then those boundaries with similar attributes (i.e., those that were generated by similar base level changes) are put in the same class of boundary. For example, all boundaries with a large areal extent of the unconformity, a thick section of strata eroded by the unconformity, a large deepening on top of the unconformity, a substantial change in depositional regime and a major shift in tectonic setting across the boundary might well be placed in the same boundary class. Such boundaries can be differentiated readily from those with a very small areal extent of the unconformity, minor erosion by the unconformity, minor deepening on top if it and no change in sedimentary and tectonic regimes across the boundary. A possible limitation of this approach relates to the fact that this hierarchy system relies heavily on the preservation of the basin margins (Miall, 1997). The various established classes are ordered in the hierarchy on the basis of the differing amounts of base level shift interpreted to be associated with each class. The class with the attributes that indicate the highest amount of base level shift is placed at the top of the hierarchy and the class with the interpreted least amount of base level shift is placed at the bottom. The naming of each established class in such a relative magnitude hierarchy is discussed below.

Hierarchy nomenclature In terms of naming the various classes of sequence boundaries that have been recognised in a hierarchy, we think the easiest and least confusing method employs a simple numerical ordering. The use of specific names for a given order (e.g., megasequence for the largest magnitude boundary) just adds more jargon to a critically overloaded system and tends to hinder, rather than enhance, communication and understanding. For a given study, the sequence boundaries that are in the class that represents the largest base level changes are designated as first-order boundaries. Thus they would constitute the largest scale boundaries recognised in the study area. We emphasise that a first-order boundary recognised in one study may be very different from a first-order boundary recognised in another, and that there is no such thing as a standard first-order boundary with specific characteristics. First-order boundaries are simply the largest class of sequence boundaries recognised in a study, and one person's first-order boundaries in their study may be very similar to another person's third-order boundaries recognised in another study. This is in contrast to the methodology of Vail et al. (1991) and Krapez (1996, 1997), who infer that a single hierarchical scheme can be applied to all basins, irrespective of age, throughout the world. Such a methodology is derived from the belief in the existence of an orderly set of global base level cycles with set periods and amplitudes, although there is little inductive evidence to support the existence of such cycles above those in the Milankovitch band.

8.2. ConceptsofSequenceStratigraphy (a) BASIN MARGIN

703

BASIN CENTRE

(b)

1s'order

,,,,!r,,er '--3r,!r,,er 2ndorder

4th~rder ,-,-,--,..,,-,..,,-~

4--1d~ror 3 d 3

3rd order

4th order 5th order

Subaerial unconformity Maximumregressive surface

~

d

e

122n rder

er

]V~2--

r

3rd-~order 3+ 2nd order . Iorder 3rd I

3rdorder 2m

Fig. 8.2-12. Hierarchy system based on the magnitude of base level changes that resulted in the formation of bounding surfaces (from Embry, 1993, 1995). (a) Schematic depiction of the five orders of sequence boundaries determined from boundary characteristics which reflect base level changes. (b) Principles of determining the order of a sequence. A sequence cannot contain a sequence boundary with the same or lower order than its highest order boundary; the order of a sequence is equal to the order of its highest order boundary.

Summary Our methodology emphasises the establishment of a hierarchy based on the interpreted magnitude of the base level changes that generated the sequence boundaries (Fig. 8.2-12). If one wants to establish a hierarchy for sequences as well as sequence boundaries, the various sequence boundaries must be ranked first. The order of a sequence is equal to the order of its lowest magnitude boundary (Fig. 8.2-12). Thus a sequence with a fourth-order boundary at the base and a first-order boundary on top is a fourth-order sequence. This brings us back to our original problem of trying to avoid a chaotic and senseless delineation of sequences in a succession with multiple sequence boundaries. With the establishment of a hierarchy as described above, one simple rule now allows recognition of a sensible and orderly succession of sequences. This rule states that a sequence cannot contain within it a sequence boundary that has an equal or greater magnitude than the magnitude of its lowest magnitude boundary. For example, a second-order sequence cannot contain a first- or second-order boundary within it but can contain one or more third- and fourth-order boundaries. This is of utmost importance and is the only way that a chaotic delineation of sequences can be avoided and an orderly one produced.

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Chapter 8: Sequence Stratigraphy and the Precambrian

Discussion and Conclusions

One of the main problems in practicing sequence stratigraphy is the a priori interpretation in terms of the external controls on sedimentation that is often attached to the stratigraphic data observed in the field. This problem is easy to surpass, and all it takes is the basic understanding that base level changes, and implicitly the shifts in depositional trends, can be controlled by any combination of eustatic and tectonic forces, and that the dominance of any of these allogenic mechanisms should be assessed on a case-by-case basis. The interplay of eustasy, tectonics (as well as other implicitly related factors like mantle processes, palaeoclimate and continental freeboard) have been discussed recently for Precambrian basins (e.g., Eriksson, 1999; Eriksson et al., 2001b) (section 7.1). One other obstacle is the proliferation of a complex jargon that includes both terms with different meanings in the light of different models (e.g., see systems tract terminology in Fig. 8.1-2) and different terms with the same meaning (e.g., see synonymous terms for sequence stratigraphic surfaces mentioned earlier in this section). This terminology barrier is in fact trivial, as the sequence models currently in use have a lot of common ground, with the main differences being in the style of conceptual packaging of the same succession of strata. Once these differences are understood, the practitioner has the flexibility of using whatever model works best for the particular circumstances of a specific case study. It is more important to identify the depositional products related to specific types of shoreline shifts, such as normal regressions, forced regressions or transgressions, than the arbitrary choice of where to place a systems tract or sequence boundary. The success of such analyses depends of course on the quality and type of data available for interpretation. For example, seismic data allow one to infer the approximate positions of the potential candidates for the correlative conformity, based on regional stratal stacking patterns. Well logs, core, and most outcrop data only allow the mapping of surfaces that are associated with changes in lithofacies and grading. The latter attributes make the maximum regressive surface easily recognisable within conformable shallow marine successions, in contrast to the correlative conformity (sensu Posamentier et al., 1988, or Hunt and Tucker, 1992). The core of sequence stratigraphy is the recognition and correlation of sequence stratigraphic surfaces, which are defined as surfaces that can serve, at least in part, as systems tract or sequence boundaries. It is less important how we name the packages of strata between specific combinations of sequence stratigraphic surfaces. As mentioned above, it is also important to recognise the type of shoreline shifts, as the distribution of sand (potential reservoirs) between the contemporaneous depositional systems is markedly different as we move into different segments of the base level cycle (Figs. 8.2-6 to 8.2-11). Stratigraphic surfaces may replace (rework) each other, in which case the name of the youngest surface, which imposes its attributes on that particular contact, should be used. Physical attributes are fundamental for both the recognition of surfaces and their hierarchical orders. We often do not know the temporal scale of the sequences we deal with, especially in the Precambrian, so the use of specific names or hierarchical orders for specific temporal scales may become very subjective. In addition to this, the dynamics and rates of plate tectonic processes likely changed from the Precambrian to the Phanerozoic (sec-

8.3. Development and Sequences o f the Athabasca Basin

705

tion 3.6), which means that the temporal duration of tectonics-driven stratigraphic cycles is irrelevant in establishing a sequence hierarchy. It is thus difficult to apply a hierarchy system based on boundary frequency, as proposed by the Exxon research group for the Phanerozoic (Vail et al., 1977, 1991) and as adopted by Krapez (1996, 1997) for the early part of the Precambrian. A practical solution to this problem is to deal with the issue of hierarchy on a case-by-case basis, assigning hierarchical orders to sequences and bounding surfaces based on their relative importance (inferred magnitudes of base level shifts) within each individual basin. This method may prove to be more realistic given the fact that each basin is unique in terms of formation, evolution, and history of base level changes. Sequence stratigraphy can be applied successfully to the analysis of the Precambrian rock record, as demonstrated in several recent case studies, including those in this chapter. Even in the absence of a high resolution time control, sequence stratigraphic models at both larger and smaller scales may be constructed based on a good knowledge of facies architecture and relationships within a study area. Such modelling provides a valuable tool for the reconstruction of the history of sedimentation, shoreline shifts and base level changes during Precambrian time.

8.3.

DEVELOPMENT AND SEQUENCES OF THE ATHABASCA BASIN, EARLY PROTEROZOIC, SASKATCHEWAN AND ALBERTA, CANADA

P. RAMAEKERS AND O. CATUNEANU Introduction

The relatively little disturbed late Palaeoproterozoic Athabasca basin is a lithologically homogeneous intracontinental basin. It comprises four sequences, with about 1800 m of largely quartzose sandstones, with varying amounts of interstitial clay, minor silt and mudstones that are capped by about 500 m of oolitic and stromatolitic dolomite, preserved only in a central circular depression due to a meteor impact of Middle Palaeozoic age (Bell, 1985). This section aims to: (1) place the Athabasca basin in the context of approximately contemporary basins in the Athabasca region; (2) summarise the development of sequences within the basin; (3) relate sequences and drainage systems in the basin to possible source regions and events beyond the basin boundaries; (4) comment on depositional tempos; and (5) comment on basin type and mode of formation. About fifteen percent of the preserved basin has seen intensive exploration with over 30,000 continuously cored drill holes to basement, largely at the shallower margins and in the uranium-rich eastern part. Nevertheless, the bulk of the basin, especially its deeper interior, remains poorly known, with only half a dozen holes penetrating most of the section. The homogeneity of the sandstones makes detailed correlation difficult on lithologic or petrographic grounds. Only three widespread unconformable surfaces can be traced readily regionally, in addition to half a dozen other lithologic boundaries. The lack of biota The Prec(unbrian Earth: Temposand Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

706

Chapter 8: Sequence Stratigraphy and the Precambrian

and radiometrically datable horizons means that the ages of deposition for the various sequences are largely inferred from regional relationships. Thus, understanding of the basin is still largely in a reconnaissance phase.

Geological Setting and Geochronology of the Athabasca Basin The Athabasca Basin straddles the Archaean Rae and Hearne provinces that lie between the stable Archaean Slave, Superior and Wyoming cratons (Hoffman, 1990; Fig. 8.3-1). The cratons are flanked by 2.0-1.68 Ga thrust-fold (Trans-Hudson, Thelon, Forwards, Central Plains, Wollaston) and magmatic (e.g., Taltson, Wathaman, Rimbey, Vulcan, Clearwater) belts. The Clearwater magmatic belt and the Wollaston folded and thrusted continental margin belt are partially overlain by the Athabasca basin. Temporal relationships are shown in Figure 8.3-2. The history of sedimentation in the Athabasca region from 2 to 1.6 Ga is largely the record of crustal contraction between the Slave and Superior cratons, and, from 1.6 to the present, it records events at the margins of Laurentia. The earliest, weakly deformed, post-collisional sedimentary rocks in the region are those of the greenschist grade Thluicho Lake (Scott, 1978) and Nonacho Groups. The Nonacho basin (Aspler and Donaldson, 1985; Fig. 8.3-1) developed as a result of indentation tectonics, in this case sinistral wrench faulting when part of the Rae province south of the Slave craton moved south as a result of the Slave-Rae collision (Gibb, 1978). The Thluicho Lake Group lies along another set of releasing bends along the same faults (e.g., Yatsore-Hill Island fault, Fig. 8.3-3), just north and also underneath the northwestern part of the Athabasca basin. At this time the Slave-Rae-Hearne block was separated from the Superior craton by the Manikiwan ocean (Stauffer, 1984; Symons, 1991). Closing of this ocean led to the development of the Wathaman magmatic belt at about 1.86 Ga along the eastern side of the future Athabasca basin, followed by the Trans-Hudson orogen in the Reindeer-Hearne collision (1.86-1.84 Ga; Machado, 1990). The resulting compression led to dextral movement of the Hearne province along the Grease-Straight fault leading to formation of a magmatic complex in the Clearwater Domain compressional zone (Figs. 8.3-1-8.3-3), in part underlying the modern Athabasca basin. Dextral movement on the Black Bay, St. Louis, and other faults along the western half of the Athabasca basin area likely occurred as part of these events and resulted in deposition of the Martin Group in wrench fault basins. The Martin Group (Beck, 1969; Tremblay, 1972; Langford, 1981; Mazimhaka and Hendry, 1984, 1985) consists of fluvial conglomerates, sandstones, mudstones and volcanics, and is overlain unconformably by the Athabasca Group (Scott, 1978). Further events that may have affected sedimentation in the Athabasca region are the Reindeer-Superior collision Phase 1 (1810-1790 Ma) and Phase 2 (1740-1720 Ma) (Machado, 1990). The ongoing contraction resulted in the formation of the Athabasca basin (Gibb, 1983). The associated crustal scale thrust faults (White and Lucas, 1994; Beaumont et al., 1995; Hajnal et al., 1996), in places with uplifts of up to 10 km, were probably the main sediment sources of the Athabasca Group, the last Proterozoic deposits of the area.

8.3. Development and Sequences of the Athabasca Basin

707

Fig. 8.3-1. Regional setting of the Athabasca Group. Domain boundaries after Hoffman (1990) and Villeneuve et al. (1993).

708

Chapter 8: Sequence Stratigraphy and the Precambrian

Hottah Domain Amundsen Basin 900 __[ Shaler Supergroup 1000 Rae Group

1100

Rae Prov. Thelon Basin

Hearne Province Athabasca Basin

TransHudson Orogen

Other orogens and events Breakupof Rodinia Grenville Orogeny

Erodedbeds? I ThelonFm Dykes U ore age 1~ PrimaryU ore

1200

I CoDoermlnebasalts 1300

" Thelon Fm MainL ore agt:s

Dismal Lakes Group

Mackenziedykes

Be' [__

I

Rifting

Supergroup

Erodedbeds? Main primary U ore ages

New Mexico to SE Ontario Calc-alkaline to alkaline plutons

1400 i1E = --]Atha- I I .. ~ 1--lbasca I ~,,,I ; F-- H-Gr~ I ~ ,,, MSeq.4?I,,,E~I ~ Hor nowI~-~o I ,~ Llerodedl "-'1 ~, Ilbeds? l

1500

1600

40~ Abrupt latitude changes Laurentia 80~

I ~1 .~1 ~J ~:1 I

s

1650 -

1700

1750

1800

1850

I

(.99 ~ I

a_ J~j Athabasca I Kaertok Fm l, .-~(~.,.~._i n--LI ( I~] Sea. Group J Labradori J an. East I Narakayvolc. I-~ o 3+4 I] Orogeny River JForwardOrog. I---- m ~ !' ' l-m ~ ~--.-~ L Athabasca LI ,_, LadyNye I ..u I I Groun l J " 'Fm ~----~ rm [~ Seq.~ [--1 ~P~ Slav~ I-' I Athabasca~ ] ,-~ >o_-~ ~ Reindeer-Superior [--]~E >. ].-. I t"ae -' I IGroup Seq.l?t-~ o "~ .~ ~ Collision2 ~].~:~ >, I con[r lciion , ' 1 ~ cfi / E IE > , ,-, t I Pitzv01.NueltinIntr, F ] ~ l-~ E I =~"o I Athabasca~i~ ~ o2~ / ~ Eo I Gp. Seq.l?q ~s ~ C0ol,nq,upl,ft ~~ ~ault Riv.Fm . . . LI _~1,._ -- pegmatite I. ' . / -o ~ , ~ . ~ 8 emplacementi HombyBayGr I ~ I laker n Martin t ~ ~ Metamorphic[ BigBearFm I ~ El .aKe 14 .., I n , -~ oo o Peakl / ~ -~l ;roup H L~roup --,-6 ~__ ~,~:.~ ]Reindeer n.'o J i~.___] c~ Reame Wathaman GreatBear - ~-' E Upper I..~__1 Batholith WollastonGr . I Arc Batholith Wopmay magmatism Orogen ~ 8,assembly ~

-~~

l

1900

1950

t~_ E = =~ ooE ~_~

NonachoLake t Thluicho Groups

Lower Wollaston Gr. rifting Taltson Foreland?

Fig. 8.3-2. Ages of selected intrusive, volcanic, sedimentary and tectonic events affecting the Athabasca region. Modified after McGlynn and Irving (1981), Hoffman (1988), Machado (1990), Cumming and Krstic (1992), Ross (2000), Kyser et al. (2000) and Santos et al. (2002).

8.3. Development and Sequences of the Athabasca Basin

709

Fig. 8.3-3. Sub-basins and major faults of the Athabasca basin. SHEAR ZONES (SZ): BLSZ, Bayonet Lake; BLKSZ, Black Lake; CBSZ, Cable Bay; CLSZ, Charles Lake; LLSZ, Leland Lake; NFSZ, Needle Falls; PLSZ, Parker Lake; RLSZ, Reilly Lake; VRSZ, Virgin River; FAULTS (F): BF, Bustard; BBF, Black Bay; BRF, Beatty River; CF, Chariot; FF, Fidler; HF, Harrison; SLF, St.Louis; TFS, Tabbernor fault system; YHF, Yatsore-Hill Island; SUB-BASINS (B): MB, Martin; TLB, Thluicho Lake. BMT, basement. Stratigraphic abbreviations as in Table 8.3-2.

Basins similar in size to the Athabasca basin and of roughly comparable age (Fig. 8.3-1) include the Amundsen, west of the Slave province and Wopmay orogen (Upper Hornby Bay Group, 1633 Ma; Bowring and Ross, 1985), the Thelon overlying the Rae province 300 km to the north, and the Sioux Sandstone 1500 km to the south at the southern margin of the Trans Hudson orogen and Superior province. Renewed contraction between the Slave craton and the Hearne province is also documented north of the Athabasca region, between 1750 and 1735 Ma (Henderson et. al., 1990; Rainbird et al., 200 lb, 2002a) and is considered to be a reason for the formation of the Thelon basin (Henderson et al., 1990). The virtually unmetamorphosed clastics of the Martin, Athabasca and Reilly basins overlie metasediments and plutonic rocks of greenschist to granulite grade, indicating a long period of uplift and erosion of the underlying Hudsonian and older rock units. No precise age is available for the deposition of the Martin Group. Evidence reviewed by Scott (1978) suggests an age between 1830 and 1780 (+ 20) Ma. If correlation with the Baker Lake Group (see section 3.5) is valid, an age of about 1830 Ma (Rainbird et al., 2002a) is reasonable.

710

Chapter 8: Sequence Stratigraphy and the Precambrian

Possible sources for rhyolitic shards in the basal units of the Athabasca Group (Pacquet and McNamara, 1985) might include the Pitz volcanics and associated Nueltin intrusives to the northeast (1765 Ma; Peterson and van Breemen, 1999; Peterson et al., 2000) or the poorly studied volcanics present at the northwestern margins of the basin lying just below or within the basal units of the Athabasca Group (Harper, 1996). Apatite cement from the Fair Point Formation (Sequence 1 of the Athabasca Group) and the base of the Wolverine Point "b" unit (Sequence 3 of the Athabasca Group) provides a poorly constrained U-Pb date of 1700-1650 Ma (Cumming et al., 1987). This material in places pseudomorphs volcanic glass shards that were derived from post-Hudsonian volcanic and intrusive suites. Detrital zircons from the Wolverine Point "b" unit give a maximum age of c. 1.66 Ga (Rainbird et al., 2002b). Palaeocurrents from the Wolverine Point and underlying beds show a source to the south, and suggest derivation of the zircons from the Central Plains (1.78-1.68 Ga; Sims and Peterman, 1986), Yavapai (1.79-1.69 Ga; Karlstrom and Bowring, 1988) or the Mazatzal orogens (1.71-1.62 Ga; Karlstrom and Bowring, 1988). Tilting of the basin to the northwest, at this stratigraphic level, suggests crustal loading in that direction and that deposition of this unit was coeval with the Forwards orogen (1.633 Ga; Cook and MacLean, 1995; Bowring and Ross, 1985). A K/Ar age of 1292 -+- 27 Ma has been obtained from illites in the Douglas Formation near the top of the Athabasca Group (Clauer et al., 1985) and provides a minimum age for the top of the Athabasca Group.

Athabasca Group: Depositional Sequences The sediments of the Athabasca Group accumulated in non-marine environments, ranging from fluvial to lacustrine and aeolian (Tables 8.3-1 and 8.3-2), with the possible exception of the uppermost (Carswell) formation. This succession is divided by subaerial nonconformities into four depositional sequences (termed here "Sequences 1-4"). Figure 8.3-4 illustrates north-south and west-east cross-sections through the Athabasca basin that show the sequences, and the major facies distribution. Tables 8.3-1 and 8.3-2 summarise the depositional environment and main lithologies of each of the facies, and show the stratigraphy and facies distribution of each sequence. Sub-basins within the Athabasca basin (Fig. 8.3-3) are apparent in cross-sections and maps of individual sequences (Figs. 8.3-4-8.3-6). The Jackfish basin is restricted to the northwest and formed during the deposition of Sequence 1. The most prominent sub-basin is the Cree basin underlying the eastern two-thirds of the Athabasca basin and formed largely during deposition of Sequence 2, as did the Beatty trough in the southwest. Deposits from Sequences 3 and 4 are now restricted to an area between the Charles Lake and Black Lake shear zones: the Mirror basin (Ramaekers, 1980). This name was erected for the central thick zone shown by seismic work in the Athabasca basin (Hobson and MacAulay, 1969). It has been partitioned by later uplifts along the Bartlett and Patterson highs into the Lillabo trough, a depression at the north end of the Beatty trough and a trough in the central Athabasca basin. The Lillabo trough continued subsiding until after deposition of the Carswell Formation, the youngest preserved unit of the Athabasca Group.

8.3. Developmentand Sequences of the A thabasca Basin

711

Table 8.3-1. Facies, depositional environment, and lithology of the Athabasca Group Faces 12

Depositional environment Playa lakes, sheetflow, braided streams; aeolian influenced

Lithology Well-sorted fine and medium sandstones, with rounded, small, hard intraclasts, mudstones 0-200 cm thick, volcanic ash pseudomorphed by apatite cement

11

Braided streams, sheetflow, minor playa lakes, aeolian influenced

Well-sorted fine and medium sandstones, mudstones 0-50 cm thick

10

Braided streams, sheetflow

Coarse to fine sandstones, more common mudstones 0-20 cm thick, rare 1 layer thick pebble horizons

Braided streams, low palaeoslope

Medium to fine sandstones, abundant large angular clay intraclasts, 1 layer thick pebble beds (Moosonees drainage only)

Braided streams

Coarse to fine sandstones, rare thin mudstones

Braided streams

Coarse to fine sandstones, pebbly sandstones, 1 layer thick pebble beds, minor clay intraclasts, rare thin mudstones (more common in Moosonees drainage)

Braided streams, sheetflow gravel, higher palaeoslope

Coarse to medium sandstones, pebbly sandstones, conglomerates

Braided streams, sheetflow, hyperconcentrated flow

Coarse to medium sandstones, pebbly sandstones

Braided streams, hyperconcentrated flow, sheetflow

Pebbly sandstones, coarse to medium sandstones, thin conglomerates

Hyperconcentrated flow, braided streams, sheetflow, debris flows?

Pebbly sandstones, 1 layer thick pebble beds, granule to medium sandstones

Hyperconcentrated flow, sheetflow, debris flows, braided streams

Pebbly to cobbly sandstones, thin conglomerate beds, minor mudstones

Hyperconcentrated flow, debris fows, braided streams

Cobbly and pebbly conglomerates, minor granule to coarse sandstone, minor mudstones

(Intermittently present at base of FP, MF; mappable locally.) Sheetflow, braided streams, small playa lakes

Pebbly sandstones, sandstones, mudstones

Sequence 1: Fair Point Formation (FP)

Sequence 1 appears to be restricted to the area west of the Clearwater domain and is best developed within the Jackfish basin (Figs. 8.3-3, 8.3-5 and 8.3-6). It overlies metamorphosed folded and thrusted basement of early Palaeoproterozoic to Archaean age. The Fair Point Formation (Ramaekers, 1979, 1980, 1990, 2003; Wilson, 1985) comprises the entire sequence.

712

Chapter 8: Sequence Stratigraphy and the Precambrian

Table 8.3-2. Sequences, lithostratigraphic units, facies, distance to source area, and depositional lithology of the Athabasca Group Sequences

Stratigraphic units

Facies

Carswell Formation (CF) Douglas Formation (DF) Otherside Formation (OF) OFb OFa

Carbonates 11

Locker Lake Formation (LL) LLc LLb LLa Wolverine Point Formation (WP) WPc WPb WPb3 WPb2 WPbl WPa WPa2 WPal

Distance to source Authigenic

Dominant lithology at time of deposition

Quartz arenite, minor sublithic arenite, subarkose, arkose

8 7 Far

Quartz arenite and sublithic arenite

Far

Bimodal: largely arkose with minor quartz arenite

5 4 5 11 12 11 12 7, 8 10

Quartz arenite to minor arkose Far

Quartz arenite, minor subarkose

9 7, 8

More distal

Quartz arenite, minor sublithic arenite

MFb

6

Proximal

Sublitharenite

MFa MFa2 MFal

3 0

Proximal Proximal

Sublitharenite

3 2 1 3 0

Proximal Proximal Very close to source

Lazenby Lake (LzL) Manitou Falls Formation (MF) MFd MFc

Fair Point Formation (FP) FPc FPb FPb2 FPb 1 FPa FPa2 FPal

5 (coarse at base only)

Arkose to subarkose

The Fair Point Formation includes three regionally mappable sandy to cobbly units (FPb 1, FPb2 and FPc). These are underlain by a discontinuous pebbly sand and siltstone unit (FPal) that may grade up into a pebbly sandy unit (FPa2) with lithofacies like those of the FPc. FPb 1 consists largely of coarse conglomerates up to 2 m thick, of debris flow and stacked sheetflow origin, and of pebbly sandstones deposited by hyperconcentrated flows, debris flows and minor braided streams. FPb2 is similar but lacks the massive thick conglomerates. The FPc unit consists largely of pebbly hyperconcentrated flow deposits

Developnlent and Sequences of the Athuhasca Basin

Fig. 8.3-4. North-south and west-east cross-sections through the Athabasca basin showing sequences, facies and structural elements (location of cross-sections are shown in Fig. 8.3-3; stratigraphic abbreviations as in Table 8.3-2).

714 Chapter 8: Sequence Stratigraphy and the Precambrian

Fig. 8.3-5. Depositional sequences of the Athabasca Group: distribution and surface palaeocun-ent directions. Stratigraphic abbreviations as in Table 8.3-2.

8.3. Development and Sequences of the Athabasca Basin 715

Fig. 8.3-6. North-south and westxast cross-sections through the four Athabasca sequences, flattened at the top of each sequence. Stratigraphic abbreviations as in Table 8.3-2.

716

Chapter 8: Sequence Stratigraphy and the Precambrian

interbedded with fining-up braided stream deposits. Sandy units, up to several metres thick, that lack obvious sedimentary structures may be gravity flows. Pebbles in the Fair Point Formation are polymict in marked contrast to the monomict quartz pebbles of higher units and consist of well rounded, subspherical quartzite and gneissic pebbles and subordinate, flatter, dark brown, well-indurated, fine sandstone to mudstone pebbles. The latter may be intraformational, but more likely are reworked from the Martin Group or other underlying units. Dark brown, usually flat and angular finegrained pebbles are referred to by Wilson (1985) as regolith material, but their presence throughout the Fair Point Formation suggests other provenance; it may be volcanic. The Fair Point Formation lags may represent deflation surfaces, and possibly mark discontinuities in deposition, although no ventifacts were noted. Palaeocurrent data from the Fair Point Formation are sparse and come from near its base; they indicate drainage to the northwest and west (Fig. 8.3-5).

Sequence 2: Manitou Falls Formation (MF), Reilly Lake beds Sequence 2 consists of the Manitou Falls Formation (Ramaekers, 1979, 1980, 1981) and overlies lower amphibolite to granulite grade metamorphic rocks, except in the western part of the basin where it unconformably overlies the Fair Point Formation of Sequence 1. There, the abrupt disappearance of the coarser material, a change in clast lithology, and the disappearance of much of the interstitial clay indicate that a different sediment source was tapped by the Manitou Falls Formation. Sequence 2 represents the bulk of the Athabasca Group, especially in the Cree basin (Figs. 8.3-3, 8.3-5 and 8.3-6). Abundant palaeocurrent data indicate that it was deposited by at least four principal drainage systems with provenance from the northeast (Moosonees drainage), east (Ahenakew drainage), southeast (Karras drainage) and northwest (Robert drainage). Palaeocurrents in the coarser basal deposits trend northerly; higher up and towards the centre of the basin they trend more westerly, perhaps indicating that the system is more complex than indicated here. The Manitou Falls Formation is a pebbly to sandy, overall fining-up unit with a maximum thickness of about 900 m. It consists of four members, designated informally MFa to MFd (Table 8.3-2). Pebbles are almost exclusively quartz. Rare sandstone clasts and rare abraded quartz overgrowths indicate that some of the material is recycled. MFa consists of an intermittently distributed mudstone and pebbly sandstone unit (MFal), similar to the basal FPal unit of Sequence 1. This unit is present both overlying the basement and the Fair Point Formation. The bulk of the MFa unit (MFa2) is a pebbly sandstone, deposited by hyperconcentrated flows, and sheetflow similar to FPa2 and FPc of Sequence 1. The extent of the MFa member is poorly known. Its thickness is limited, possibly not much more than the palaeorelief (Fig. 8.3-4). Palaeocurrents, measured in outcrop and open pit mines trend northerly, but there are few data. The MFb unit thickens and coarsens to the east (Fig. 8.3-6). It consists of pebbly braided stream and sheetflow sandstones with 2-20% conglomerates. Two overall fining-up cycles are generally present. The MFc member is similar to MFb but lacks conglomerates and has minor amounts of clay intraclasts. It overlaps MFb at the western margin of the Athabasca basin. The MFd member is characterised by medium- to fine-grained sandstones with abundant clay intr-

8.3. Developmentand Sequences of the Athabasca Basin

717

aclasts, best developed in the eastern Cree basin. The change between MFc and MFd is fairly abrupt, and might indicate minor discontinuities, but more likely indicates a lower palaeoslope in a more distal part of the depositional system. The Reilly Lake beds occur in a single outcrop (Fig. 8.3-5), separated by the Wathaman batholith from the Athabasca basin. Lithologically they resemble the MFa member, but the palaeocurrent directions are to the southwest, in marked contrast to the northerly directions obtained from the MFa in the Athabasca basin. This suggests that the Reilly Lake beds are the remnant of a separate basin. Sequence 3: Lazenby Lake (LzL) and Wolverine Point (WP) Formations Sequence 3 (Figs. 8.3-5 and 8.3-6) comprises the Lazenby Lake (LzL) and Wolverine Point (WP) Formations (Ramaekers, 1979, 1980, 1990, 2003). The base of the sequence is usually marked by a relatively thin conglomeratic layer deposited by hyperconcentrated flows and sheetflows, but in places this coarse layer is underlain by a coarsening-up pebbly sandstone. The conglomerate grades up into pebbly braided stream and sheetflow sandstones. The unit is thickest in the Beatty trough (Fig. 8.3-3) and thins to the north and east, disappearing in the subsurface in the middle of the basin. Palaeocurrent directions are to the north and northeast, and perhaps more westerly at the eastern margin of the unit. The difference in directions of thinning, lithology, and in palaeocurrent trends compared to the underlying MFd, indicate a dramatic change in the organisation of the basin. One quartz pebble ventifact, facetted before burial, was found in the basal unit of the Lazenby Lake Formation. The Lazenby Lake Formation grades upwards into the Wolverine Point Formation, which is characterised by the more common presence of mudstones and claystones. The base of the Wolverine Point is taken at the point where 5-20 cm thick mudstones become more common and occur more regularly. Three informal members are recognised, WPa to WPc. WPa consists of a basal section with fairly regular, < 30 cm thick mudstones, and a sandier upper unit. It was deposited in a sheetwash and braided stream environment, perhaps with intermittent small shallow lakes. WPb is characterised by the presence of thicker (> 30 cm, up to 2 m) and more common mudstones, plus layers of dark, hard and rounded small intraclasts (in contrast to the larger, angular, soft intraclasts of MFd). Within apatitecemented layers, fresh plagioclase and K-feldspar grains are present, in contrast to the rest of the unit where the matrix is often clay-rich; in places WPb shows clay pseudomorphs after sand-sized detrital grains. The unit was deposited in a sheetwash, braided stream and playa lake environment. The presence of very well-sorted fine- to medium-grained sandstones suggests that the sand may have been cycled through an intermediate aeolian stage, but was deposited below the waterline in a fluvial or lacustrine environment. The WPc member consists of--f~ne- to medium-graii~ed, well-soi-ted ~,t~d~tu~ with a few i~udstoi~e beds up to 50 cm thick. Palaeocurrent measurements in the Wolverine Point Formation are few and show variable directions: north, west and northeast, reflecting the low palaeoslope (Fig. 8.3-5). The unit thickens and contains more mudstones to the north (Ramaekers, 1990). In the northeastern part of the Athabasca basin, the WPa member is thin or absent, and the base of

718

Chapter 8: Sequence Stratigraphy and the Precambrian

the WP there overlies the MFd with a thin layer of pebbles derived from the subjacent sandstone.

Sequence 4: Locker Lake (LL), Otherside (OF), Douglas (DF) and Carswell (CF) Formations Sequence 4 (Figs. 8.3-5 and 8.3-6) forms a fining-up series consisting of the pebbly Locker Lake (LL) Formation (Ramaekers, 1979, 1980, 1990), Otherside Formation (Ramaekers, 1979, 1980, 1990), Douglas Formation (Amok, 1974; Ramaekers, 1990) and Carswell Formation (Blake, 1956; Fahrig, 1961; Hendry and Wheatley, 1985; Ramaekers, 1990). The basal contact of the Locker Lake Formation is unconformable, showing some reworking of underlying WPc material and an abrupt return to deposition of coarser material. The underlying Wolverine Point is eroded or was not deposited progressively further to the south. Contacts between the Locker Lake and Otherside Formations are gradational, and marked by the return to a maximum size of less than 8 mm for the contained pebbles. The contacts between the Otherside, Douglas and Carswell Formations are probably gradational, but as they are preserved only within a meteor impact crater this cannot be proven due to the disruption caused by the impact and subsequent slumping. Palaeocurrents from the Locker Lake Formation indicate derivation from the south (Fig. 8.3-5), similar to the drainage pattern of the Bourassa drainage system in Sequence 3. The overlying, finer-grained Otherside Formation shows drainage towards the west, with higher variability than in Sequence 1, and with drainage towards the Lillabo trough in the northwestern part of the Athabasca basin. The Locker Lake Formation is divided into three informal members, LLa to LLc, based on maximum grain size of the contained pebbles, with pebble size in the coarsest LLb unit greater than 16 mm. LLa coarsens up, and the maximum grain size and conglomerate content of the sequence is reached in LLb. Mudstones less than 50 cm thick are present but are increasingly less common upwards in the Locker Lake Formation. Above LLb, grain size decreases progressively, with much of the Otherside Formation (OFb) finer than 2 mm. The Douglas Formation consists of medium- and fine-grained sandstones with common mudstones. These are largely black and organic-rich (Landais and Dereppe, 1985; Wilson et al., 2002) but altered sections show reduction to green and pale red colours and colour patterns similar to those seen in the WPb mudstones. The Carswell Formation includes stromatolitic and oolitic dolostones and mudstones; siliciclastic input is virtually absent. The sequence records changes in depositional environment from sheetwash-dominated (Locker Lake), braided stream (Otherside Formation) to paralic (Douglas Formation) to lacustrine or marine (Carswell Formation). Good tidal indicators (section 7.5) are lacking in the Carswell Formation, the best being the north-south elongation of stromatolite domes (Hendry and Wheatley, 1985). Post-Carswell Formation units Fluid inclusion studies (Pagel, 1975a, b), organic matter maturation studies (Landais and Dereppe, 1985), and illite crystallinity studies (Hoeve et al., 1981) all indicate that the maximum depth of burial of the Athabasca Group was about 4-5 km. As the preserved

8.3. Development and Sequences of the Athabasca Basin

719

section is about 2300 m thick this means that about 2700 m of cover has been removed by erosion. Diabase dykes are not uncommon within the Athabasca basin (Ramaekers, 1980) with dyke complexes found along the southern (Cree Lake) and eastern margins (Moore Lakes). Their presence suggests that part of the missing section may have been volcanics. The three main periods of primary uranium ore formation in the Athabasca basin document extensive and prolonged hydrothermal activity in the basin (Cumming and Krstic, 1992; Fig. 8.3-2), that may have been facilitated by basin deformation accompanying deposition of the later and now eroded units. These periods match times of sediment deposition in the Amundsen basin (Fig. 8.3-2), the times of rifting along the western side of Laurentia with deposition of the Belt Supergroup (Winston, 1990), and the breakup of Rodinia (Hoffman, 1991; Ross et al., 1992; Idnurm and Giddings, 1998b).

Discussion Sequence stratigraphy In intracratonic basins controlled by contractional tectonics, sedimentation is influenced primarily by tectonism in the sediment source areas, within the basin itself, and in the downstream regions beyond the limits of the preserved basin. An additional allogenic control on sedimentation is represented by climate, which modifies the efficiency of weathering, erosion and sediment transport processes. As the preserved sedimentary fill of the Athabasca basin is dominantly non-marine, the sequence stratigraphic terminology of systems tracts and associated surfaces proposed initially for divergent continental margins (e.g., Posamentier et al., 1988) (section 8.2) cannot be applied directly to this basin. Alternative terminology is offered by studies of non-marine depositional sequences in foreland or extensional settings (Boyd et al., 1999; Zaitlin et al., 2000; section 8.4). In the absence of any evidence of what the direction of shift might have been for an age-equivalent shoreline outside the preserved basin, the use of lowstand, transgressive and highstand terminology is inappropriate. Instead, terms such as low and high accommodation systems tracts may be applied to describe the observed changes in energy levels and grain size within each sequence. The four Athabasca sequences all begin with a relatively thin, crudely coarseningupwards set of beds, followed by a series of fining-upwards beds. The coarsening-upwards beds may be discontinuous and mud-rich (FPa in Sequence 1; MFal in Sequence 2), discontinuous and sandy (unnamed coarsening-up beds below the LzL conglomerates in Sequence 3), or sandy and continuous (LLa in Sequence 4). The lower coarsening-upwards part of each sequence may be assigned to a low accommodation systems tract, with fluvial deposits infilling lows and prograding into the developing basin. Following this levelling, the early infilling deposits are overlain by a series of better sorted largely fining-upwards beds that accompany the upwards decrease in energy levels during each major depositional cycle. These deposits form the bulk of each sequence, and may be assigned to a high accommodation systems tract. The underlying assumption behind this systems tract terminology is that following the stages of uplift resulting in the formation of sequence boundaries, the rates of creation

720

Chapter 8: Sequence Stratigraphy and the Precambrian

of accommodation gradually increase from low to high during each depositional cycle. This allows more and more floodplain and associated low energy facies to be deposited as the sequence thickens, given a suitable combination of intrabasinal subsidence and a matching rate of sediment supply. The latter depends on extrabasinal processes such as weathering rate (sections 5.10 and 5.11), uplift rate, water supply, slope (all contributing to erosional rate) in a source area that may be adjacent or very distant. Over time these factors result in the gradual denudation of source areas during the deposition of each sequence, as well as in a decrease in slope gradients during sedimentation (Catuneanu and Elango, 2001; Catuneanu, 2002) and thus may contribute to the frequently observed fining-upwards trends. The concepts of low versus high accommodation systems tracts were developed for Phanerozoic sequences, where vegetation favours the preservation of thick overbank fines and isolated channel-fills under high accommodation conditions. In such environments, the low accommodation systems tract includes amalgamated channel-fills (high sand/mud ratio), whereas the high accommodation systems tract is mainly built by floodplain fines (low sand/mud ratio). The less confined fluvial systems of the vegetationless Precambrian require new criteria more applicable to such conditions. The general lack of overbank fines in the Athabasca basin may be attributed to the dominance of unconfined fluvial systems, where sheetwash facies tend to replace the vegetated overbank deposits of Phanerozoic meandering systems. The lack of fines in a sand-rich vegetationless environment may also be related to a greater aeolian influence, evidenced in the Athabasca Group, as dust storms effectively remove mud from the depositional areas. The removal of the fine-grained sediment fraction also contributes to the generation of texturally supermature beds. The ratio between sand and mud, and the associated fluvial architectural elements, seem therefore to be of less importance when trying to distinguish between low and high accommodation systems tracts in Precambrian deposits. We propose that changes in the overall grading trends, as well as the geometry of fluvial deposits, may provide more useful criteria for the study of Precambrian sequences. The low accommodation systems tract corresponds to the stage of peneplanation in a developing basin, where fluvial deposits prograde and infill an immature landscape. The gradual progradation of coarser facies from outside the basin and the mixing with locally eroded muds, sands and channel bank debris may generate the observed crudely coarsening-upwards trends. Topographic irregularities above the sequence boundary between incised valleys and interfluve areas give a potentially discontinuous geometry to this systems tract, with significant changes in thickness along dip and strike. The high accommodation systems tract has a more predictable and continuous geometry, either sheet- or wedge-like depending on subsidence patterns, and is dominantly aggradational. It corresponds to the stage of decline in the energy level of the fluvial systems, which confers to it an overall fining-upwards trend that may reflect any suitable combination between accommodation, denudation and gradient controls. The thickness of trough cross-beds in cosets may be an indication of low or high accommodation environments. In the Athabasca Group fluvial trough cross-bedding, with cosets of troughs 30-120 cm wide, is very common and characterises channel deposits. In outcrop, well-developed cosets of such cross-beds, in effect climbing trough cross-beds,

8.3. Development and Sequences of the Athabasca Basin

721

may show thicknesses of individual cross-beds of a few mm to 10 cm. Where these are thin, only the toes of the troughs are preserved and the sections shows apparent horizontal bedding to low angle cross-bedding. Outcrops with good horizontal and vertical exposure are common in the Athabasca basin, and interpreting these beds as due to large trough cross-beds is not difficult. However, in core studies such beds may easily be confused with horizontal bedding, low-angle cross-bedding or ripple cross-lamination if the material is well sorted, making correct interpretation of depositional environment impossible. In a degradational or low-accommodation system such bedforms are likely to leave no deposits or cosets of thin lamina resembling horizontal bedding, low-angle cross-bedding, or ripple cross-lamination, each lamina produced by succeeding dunes in the train. In a high accommodation system, especially where the channels spread out into unconfined flow, cosets of thicker trough cross-beds are more likely to be produced. The boundary between the low and high accommodation systems tracts proposed for the Athabasca sequences is reasonably well defined at the bases of FPb (Sequence 1), MFa2 (Sequence 2), LzL conglomerate bed (Sequence 3), and LLb (Sequence 4), but it should by no means be regarded as a single plane, or as a chronostratigraphic horizon. The change from low to high accommodation conditions, according to the criteria defined above, is potentially diachronous across the basin, possibly younging in a distal direction. This makes the fluvial systems tracts very different from the conventional systems tracts defined in marine to non-marine facies transitions, where the timing of systems tract boundaries depend on shoreline shifts and are close to time lines along dip directions (Catuneanu, 2002) (section 8.2). Basin order

In the Jackfish basin three once extensive and thick first-order sequences separated by major nonconformities are present: the Thluicho Lake, Martin and Athabasca Groups. The tectonic events and accompanying erosion separating them were intense enough to fold, uplift and largely remove the underlying sequences. Such low preservation potential may be characteristic of basins in an environment of ongoing crustal contraction. Very little or no sedimentary evidence may be left of once substantial basins. The Kimiwan isotope anomaly of Burwash et al. (2000) in central Alberta (about 1.8 Ga) is interpreted by them as a zone of extension; it may be the sole remains of a basin coeval with the Martin Group. The basal Athabasca unconformity (base of Sequence 1) corresponds to the most prominent change in tectonic style in the Athabasca region, from pull-apart basins along wrench faults involved in escape tectonism as seen in the Thluicho Lake, Nonacho and Martin Groups, to thick-skin compressional and flexural tectonism produced by ongoing compression after the initial orogenic episode. The latter tectonic regime led to the development of broader basins with regionally distributed sequences (e.g., Athabasca and Thelon basins). The basal unconformity of the Athabasca Group therefore qualifies as a first-order sequence boundary, marking the shift from escape tectonism characterised by the prevalence of wrench faulting, to subsidence over larger areas that extended progressively to the east.

722

Chapter 8: Sequence Stratigraphy and the Precambrian

The subtle nonconformities of the Athabasca sequences, disconformable at outcrop scale, likely mask major depositional gaps and shifts in tectonic regime that would be much more obvious in post-Devonian strata. In particular, the unconformity between Sequences 1 and 2, involving a shift from a northeast trending smaller basin and a lithologically immature fill to a much larger east-west trending basin with a mature to supermature fill, may represent a much larger depositional gap than any other present in the basin. More than one orogenic event appears to be involved in the deposition of the Athabasca Group: the Trans-Hudson orogen for Sequences 1 and 2, the Mazatzal and Forwards orogens for Sequence 3 (and 4 ?). The four depositional sequences provide the basic subdivision of the first-order Athabasca sedimentary fill, and therefore can be regarded as second-order sequences. This interpretation may be revised in the future if sufficient evidence can be found to upgrade the status of any one boundary to a first-order level (see section 8.2 and Catuneanu, 2002, for discussions on the issue of sequence hierarchy). The three post-1380 Ma periods of disturbance in the Athabasca basin, inferred from the times of primary unconformity ore generation, match basin formation in rifting events elsewhere along the western margin of North America. This suggests a switch from deposition in an overall compressive regime (Sequences 1-4), to basin formation in a tensional regime, and hence it marks the end of the first-order Athabasca sequence.

Basin development Initial subsidence of the Athabasca basin may be related to continuing crustal contraction following the Trans-Hudson orogen, the associated lateral movement of crustal sections near the indenting cratons, and the shedding of material from zones uplifted due to heating of the lower crust after subduction and their subsequent cooling. The development of the preserved parts of the basin seem to have been related to relatively gentle uplifts and subsidence, similar to the anticlines and monoclines involved in the development of the Laramide Powder River basin. Motion along sub-basin and basin margin faults to the west, northwest, and east, appear to have accompanied the formation of the basin. A number of these faults now lie outside the basin at the present level of unroofing. The depositional sequences related to Trans-Hudson uplifts (Sequences 1 and 2) show a progressive shift of depocentres from west to east, closer to the original orogen. The lowest two Athabasca Group sequences and possibly the correlative Thelon and Amundsen basin strata illustrate a mode of thick-skin tectonics somewhat different from those displayed in the formation of the bulk of the Laramide basins, where the uplifts due to imbricating crustal blocks are expressed at the surface as major thrust faults shedding clastics into a basin, and the crustal blocks seem to have behaved in a more rigid fashion compared to the blocks in the Athabasca region. Lithoprobe sections to the south of the Athabasca region (Ross et al., 2000) show that motion along the imbricating crustal blocks was taken up in part by broad folding above the thrusts. This may have led to more gradual domal uplifts such as the late large anticlines of the Trans-Hudson orogen (Hajnal et al., 1998; Fig. 8.3-1, item 6), more widespread and larger eroding source areas, and slower

8.3. Development and Sequences of the Athabasca Basin

723

rates of erosion, and suggests a different, possibly higher crustal temperature regime within the folding and imbricating blocks. Sedimentation in the Athabasca basin was disrupted by three major stages of uplift and basin reorganisation, which resulted in the formation of the three internal second-order sequence boundaries. At this point it is difficult to assess the magnitude of the stratigraphic hiatuses associated with each of these boundaries, but they were significant enough to generate profound changes in the sedimentation patterns from one sequence to another. Tectonism was clearly the main allogenic control on accommodation, as changes in tilt direction, depocentre locations, and significant changes in grain sizes are recorded across the second-order sequence boundaries (Figs. 8.3-5 and 8.3-6). The importance of the tectonic control is also suggested by the wedge-shaped geometry that characterises most of the preserved sequences, which indicates syndepositional differential subsidence likely caused by uplifts along the basin margins (Flemings and Jordan, 1990). The generally high compositional maturity at the time of deposition of Sequences 2-4, and the paucity of pebbles in Sequences 3 and 4 suggests that the uplifts were slower to develop than during Sequence 1 deposition, that weathering was very intense in the source areas, or that the source area was far away. Conclusions

The Athabasca basin contains the sedimentary record of four largely fining-up sequences, separated by subaerial nonconformities. Energy levels show an overall decrease up sequence. Hyperconcentrated flow deposits dominate Sequence 1, braided stream and sheetflow deposits dominate Sequences 2 and 3, with playa lake deposits present near the top of Sequence 3. Sequence 4 fines up from predominantly braided stream deposits (basal 30%) to paralic fine-grained clastics with carbonaceous mudstones (30%), and terminates with a thick oolitic and stromatolitic dolomite (40%). The Athabasca Group has been identified as a first-order sequence based on the changes in basin-forming mechanisms at its base and top. The underlying Thluicho Lake, Nonacho and Martin Groups accumulated in wrench fault basins associated with escape tectonism. In contrast, the Athabasca Group was related to deposition in broad subsiding areas flanked by major thick-skin antiforms formed in the hanging wall of crustal scale thrusts. PostAthabasca Group Proterozoic history is dominated by extensional tectonism in the region. The four depositional sequences provide the basic subdivision of the first-order Athabasca sedimentary fill, and can therefore be regarded as second-order sequences. The largely non-marine succession of these sequences may be split into a relatively thin low accommodation systems tract that marks initial subsidence and progradation into the developing basin, overlain by the finer-grained sediments of the high accommodation systems tract, that corresponds to aggradation in energy-declining fluvial systems. The Athabasca basin may have lacked the highly active thrusted margin of foreland basins; perhaps the perimeter type of basin of Dickinson et al. (1988), bordered by monoclines rather than active fault scarps, may be more applicable in this case.

Chapter 8: Sequence Stratigraphy and the Precambrian

724

8.4.

THIRD-ORDER SEQUENCE STRATIGRAPHY IN THE PALAEOPROTEROZOIC DASPOORT FORMATION (PRETORIA GROUP, TRANSVAAL SUPERGROUP), KAAPVAAL CRATON

E G. ERIKSSON AND O. CATUNEANU

Introduction The c. 2.7-2.1 Ga Transvaal Supergroup is preserved in three basins on the Kaapvaal craton. Within the largest of these, the Transvaal basin, rocks of the uppermost Pretoria Group form the floor to the c. 2.05 Ga Bushveld Complex (Eriksson and Reczko, 1995; Walraven and Martini, 1995) (Fig. 8.4-1). The Daspoort Formation occurs approximately in the middle of the Pretoria Group and its age is constrained by that of the Bushveld Complex and the c. 2.3 Ga Hekpoort lavas (Fig. 8.4-2). Catuneanu and Eriksson (1999) have applied sequence stratigraphy to the Transvaal Supergroup and identify two unconformity-bounded second-order depositional sequences within the Pretoria Group, namely Rooihoogte-Timeball Hill and Boshoek-Houtenbek. Pretoria Group sedimentation is ascribed to two cycles of rifting followed by thermal subsidence, with the former second-order sequence being related to plate tectonics, and the latter (which includes the Daspoort) to a continental flood basalt event (Hekpoort Formation) (Figs. 8.4-2 and 8.4-3) (Eriksson et al., 2001 c). The Boshoek-Houtenbek second-order sequence preserves a complete succession of systems tracts, comprising basal lowstand (LST), overlain by transgressive (TST), highstand (HST) and falling stage (FSST) tracts (Catuneanu and Eriksson, 1999) (Fig. 8.4-3). The Daspoort Formation itself is the tidally-influenced non-marine portion of this second-order TST, and is bound at the base by a second-order maximum regressive surface (synonymous with the "conformable transgressive surface" appellation used by Catuneanu and Eriksson, 1999; for a detailed discussion of sequence stratigraphic terminology and usage the reader is referred to Catuneanu, 2002) and at the top by a second-order ravinement surface (Fig. 8.4-3). The Daspoort Formation is thus equivalent in age to the lower portion of the transgressive epeiric Silverton Formation (Catuneanu and Eriksson, 1999) and the upper Daspoort ravinement surface is, consequently highly diachronous. This section will discuss the third-order sequence stratigraphy of the Daspoort Formation.

Sedimentology of the Daspoort Formation The Daspoort Formation is characterised by strongly recrystallised fine- to mediumgrained quartzose sandstones, with subordinate coarse-grained, pebbly and arkosic varieties, minor conglomerates, mudrocks and ironstones (Table 8.4-1). Underlying mudrocks and subordinate sandstones of the Strubenkop Formation are arranged in predominantly upwards-coarsening successions, and are thought to reflect predominantly lacustrine conditions with some fluvial inflows (Catuneanu and Eriksson, 1999). The Daspoort Formation is overlain by the argillaceous lithologies of the Silverton Formation, interpreted as essentially substorm wave-base deposits in an epeiric sea (Eriksson et al., 2002a). The Precambrian Earth: Temposand Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

8.4. Palaeoproterozoic Daspoort Formation

725

Fig. 8.4-1. Maps showing the Transvaal Supergroup, South Africa, its three preservational basins, and the outcrops of the Daspoort Formation, Pretoria Group.

726

Chapter 8: Sequence Stratigraphy and the Precambrian

Fig. 8.4-2. Lithostratigraphy, chronology, interpreted tectonic settings and depositional palaeoenvironments, and inferred base-level changes for the Transvaal Supergroup. Wavy lines suggest unconformable contacts. Modified after Catuneanu and Eriksson (1999). Age data from" (1) Armstrong et al. (1991); (2) and (6) Eriksson and Reczko (1995); (3)-(5) and (7) Walraven and Martini (1995); (8) Harmer and Von Gruenewaldt (1991).

8.4. Palaeoproterozoic Daspoort Formation

727

Fig. 8.4-3. Sequence stratigraphic interpretation of the Pretoria Group, Transvaal Supergroup (modified after Catuneanu and Eriksson, 1999). Not to scale; vertical axis suggests both time and thickness; arrows indicate directions of shoreline transgression. Abbreviations: LST = second-order lowstand systems tract; TST = second-order transgressive systems tract; HST = second-order highstand systems tract; FSST = second-order falling stage systems tract.

Despite an essentially sheet-like geometry at the regional scale, the Daspoort tends to thicken from north to south (Fig. 8.4-4), a trend also observed in the much thicker preceding rift deposits of the Boshoek and Dwaalheuwel alluvial-fluvial sandstones and the Hekpoort basaltic andesites (cf. the southwards-deepening half-graben model of Eriksson et al., 1991). The fairly complex thickness patterns observed for the Daspoort thus most

Table 8.4-1. Facies associations of the Daspoort Formation (after Eriksson et al., 1993) Facies association Sandstone

Mudrockironstone

Pebbly sandstone

Description

Interpretation

Finer facies: fine- to medium-grained quartz arenites, lesser coarse-grained sublitharenites and quartz wackes; planar bedding is the predominant structure, with upper flow regime wavy bedding surfaces. and lesser planar and trough cross-bedding, mudcracks, sand waves. ladder ripples and localised erosive surfaces Coarser facies: fine- to medium-grained quartz wackes with thin pebbly interbeds, localised coarse- to very coarse-grained sandstones; structures predominantly planar cross-bedding, lesser trough cross-bedding, channel-fills and planar bedding, and minor localised erosive surfaces, soft sediment deformation structures and herringbone cross-beds

Both facies are interpreted as relatively distal braided river deposits, subject to variable energy conditions, with tidal reworking

Mudrock facies: interlaminated mudstone, siltstone and very fine-grained sandstone, which form beds of a few centimetres to several metres in thickness, and which are interbedded in the finer facies of the sandstone association above. The thickest occurrence is 15 m. Fermginous to very fermginous; planar cross-laminated locally Ironstone facies: occur as interbeds and lenses within the above facies, with thicknesses between 1 and 2 m. Comprises a continuum from femginous quartzites to quartzitic ironstones

The mudrock facies may be at least partly intertidal deposits, based on their lithology alone, although they lack specific tidallyformed structures. These low energy deposits and the ironstone facies together suggest quiet water basinal conditions which may have been either lacustrine (associated with distal parts of the predominant braided river facies in the first association) or shallow marine

Facies include quartzose (often locally recrystallised) sandstones and arkoses, mostly mediumto coarse-grained. pebble-bearing coarse- to very coarse-grained sandstones of similar compositions, and thin conglomerate beds. Pebbles in the coarse sandstones are matrix-supported and consist mainly of sandstone and chert, whereas conglomerate pebbles are quartz, chert. jasper and mudrock. Sedimentary structures in all these rock types are predominantly planar bedding and planar cross-bedding, with lesser channel-fills and trough cross-beds, plus minor erosive surfaces, soft sediment deformation structures, current ripples. upper flow regime wavy bedding and herringbone cross-beds. In most outcrops. upwards-coarsening successions of these facies are observed. Pebbles tend to be well rounded in both conglomerates and sandstones, and are 5-25 mm in diameter. Patchy iron staining and localized pyrite nodules occur within all these facies

Interpreted as braided river deposits; their textural maturity and compositional immaturity support relatively higher energy and shorter transport distances than the predominant sandstone facies. In contrast to the latter association, there is no evidence for tidal reworking, assuming that herringbone structures can also have a fluvial origin cf. (Alam et al., 1982)

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8.4. Palaeoproterozoic Daspoort Formation

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Fig. 8.4-4. Isopach map of the Daspoort Formation (updated with new data, after Eriksson et al., 1993). Authors' data supplemented by thickness measurements from Key (1983), Engelbrecht et al. (1986) and Hartzer (1989).

likely reflect inhomogeneous thermal subsidence above an uneven, rifted floor, compaction within floor and Daspoort lithologies, as well as erosive loss at the upper Daspoort secondorder ravinement surface (cf. Catuneanu, 2002). Both basal and upper contacts are sharp and approximately conformable regionally, with a locally erosive basal contact most often associated with the pebbly sandstone facies association (Table 8.4-1) (Eriksson et al., 1993). A detailed facies and palaeoenvironmental analysis of the Daspoort Formation has been published (Eriksson et al., 1993), and an updated version of these facets of Daspoort sedimentation is presented in Table 8.4-1. The sandstone and the mudrock-ironstone facies associations are ascribed to predominant braided fluvial deposition and basinal sedimentation in the east, respectively, with a measure of tidal reworking of the sandy deposits. An apparently younger (Fig. 8.4-5) and mineralogically and texturally more immature fluvial deposit followed (pebbly sandstone facies association; Table 8.4-1). These younger fluvial deposits appear to have been related to incision of the earlier and more widespread facies associations. Although recrystallisation has hampered study of both palaeocurrents and regional distribution patterns of the three identified facies associations, some definite basin-scale trends can be identified for the preserved Daspoort depository (Figs. 8.4-5 and 8.4-6). These in-

730

Chapter 8: Sequence Stratigraphy and the Precambrian

Fig. 8.4-5. Vertical and lateral arrangement of facies associations (Table 8.4-1) across the preserved Daspoort basin (modified, with new data, from Eriksson et al., 1993). clude: some evidence for upwards-coarsening fluvial deposits; a general fining from west to east of the fluvially deposited sandstone facies association; an intimate spatial association between the finest facies (mudrock, ironstone, fine sandstones), and their preferential occurrence in the east of the preserved basin (cf. Table 8.4-1 with Fig. 8.4-5). In addition, the pebbly sandstone facies association demonstrates a greater degree of incision into the underlying finer fluvial-lacustrine/basinal deposits from west to east (Fig. 8.4-5). Most palaeocurrent data have been obtained from these pebbly sandstones, from the Pretoria region, where weakly bimodal patterns with a predominant southerly mode are recorded (Fig. 8.4-6). Very limited data from the sandstone facies association in the western half of the basin give unimodal southerly, easterly and northerly trends, and more extensive data from the east-southeast of the preserved basin are distinctly polymodal. The latter regions are the same as those where the mudrock-ironstone facies association is best preserved, commonly interbedded with the fine sandstone facies. This relationship, allied to the general west to east fining of the sandstone association as a whole, and to the fact that the eastern sandstones are pyritic (Eriksson et al., 1993), together support the interpretation that the Daspoort fluvial systems passed into a shallow basin towards the

8.4. Palaeoproterozoic Daspoort bbrmation

731

Fig. 8.4-6. Palaeocurrent data recorded from the Daspoort Formation for the sandstone facies association (main map) and for the pebbly sandstone facies association (inset map). Modified after Eriksson et al. (1993).

east-southeast. The Silverton epeiric transgression which began coevally with Daspoort tidally-influenced fluvial deposition and which obtained highstand conditions thereafter (Catuneanu and Eriksson, 1999; Eriksson et al., 2002a), is thought to have had a reverse southeast/east to westerly advance. The Daspoort-Silverton basin thus appears to have been controlled, at least partly, by the approximately east-west strike of the Hekpoort volcanic rift depository (Eriksson et al., 1991 ). Regional palaeoslope was thus approximately from west to east, with lesser S-N and N-S gradients along the E - W grain. Evidence for tidal influence on the fluvial deposits of the sandstone facies association is preserved across the basin, supporting the model that Daspoort fluvial sedimentation rates were slowly overcome by Silverton epeiric basin subsidence rates, as the transgression progressed from east to west. The age-equivalence between Daspoort and lower Silverton Formations explains these tidal influences in the sandstone facies association of the former unit. Incision at the base of the pebbly sandstone facies can be related to the change in accommodation upwards during Daspoort deposition.

732

Chapter 8: Sequence Stratigraphy and the Precambrian

Third-Order Sequence Stratigraphy of the Daspoort Formation The erosive base of the pebbly sandstone facies is the only regional surface of significance within the Daspoort Formation and can be described as a third-order subaerial unconformity. It separates a lower Daspoort unit, comprising sandstone, mudrock and ironstone facies associations, from an upper unit made up of the pebbly association (Fig. 8.4-7). This surface locally incises through the entire lower Daspoort deposit, thus reworking the second-order maximum regressive surface at the base of the Dasport Formation. As the two Daspoort units themselves lack bounding subaerial nonconformities, they cannot be termed sequences, but may be classified as systems tracts (Catuneanu, 2002, for definitions). The two non-marine systems tracts thus defined, one finer- and the other coarser-grained, correspond to the high accommodation and low accommodation systems tracts of Dahle et al. (1997), Boyd et al. (1999) and Zaitlin et al. (2000) (see also section 8.3). Within non-marine successions, low accommodation conditions result in an incised valley-fill type of stratigraphic architecture dominated by multi-storey channel-fills and generally coarser sediments which reflect the lack of floodplain aggradation. The depositional style is progradational, often influenced by the underlying incised valley topography, similar to what is expected from a lowstand systems tract (Boyd et al., 1999). High accommodation conditions (attributed to higher rates of base level rise) result in a simpler stratigraphic architecture that includes thicker- and finer-grained deposits, similar in style to the transgressive and highstand systems tracts. The depositional style is aggradational, with less influence from the underlying topography or structure (Boyd et al., 1999). We thus assign the two finer facies associations of the Daspoort Formation to a thirdorder high accommodation systems tract, characterised by a higher water table, a lower energy regime, and the deposition of finer-grained sediments. This fluvial palaeoenvironment was marked by more floodplain and lacustrine (basinal) deposits than the upper Daspoort, and an eastern depocentre can be defined, approximately coincident with the occurrence of the mudrock and ironstone facies (Fig. 8.4-8). The generally upwards-coarsening character of the sandstone facies association noted earlier (Fig. 8.4-5) may be ascribed to either crevasse-splays (unlikely to uncommon within braided fluvial systems) or fluvial progradation into the lake(s) or eastern basin. Using the same logic, the upper Daspoort (pebbly sandstone facies association) may be assigned to a third-order low accommodation systems tract (Fig. 8.4-7), characterised by amalgamated channel-fills and with no accommodation for floodplain aggradation. Within the depositional palaeoenvironment envisaged above for the Daspoort Formation, encompassing interaction of fluvial with long term (second-order) transgressing epeiric (mainly tidal) marine influences, and with variation in relative sedimentation and base level change rates, it becomes difficult to delineate shoreline shift direction with confidence at shorter time scales. For these reasons, despite the Daspoort Formation having been essentially deposited within a tectonically relatively stable coastline setting, usage of the LST, TST and HST terminology is to be avoided. The low accommodation systems tract applied to the upper Daspoort does, however, have some equivalence to an LST, in that early and slow base level rise led to a restriction of accommodation for floodplain de-

8.4. Palaeoproterozoic Daspoort Formation

733

Fig. 8.4-7. Conceptual diagram showing the two third-order systems tracts of the Daspoort Formation. Not to scale. The high accommodation systems tract (lower Daspoort) includes the sandstone and mudrock-ironstone facies associations, and is interpreted to have formed during a time of relatively high rates of base level rise. The low accommodation systems tract (upper Daspoort) includes the pebbly sandstone facies association, and is interpreted to have formed during a time of relatively low rates of base level rise. The two systems tracts are separated by a subaerial unconformity (third-order base level fall) that locally incises through the entire lower Daspoort deposits. position. The lower Daspoort high accommodation systems tract follows a second-order LST (applied to the underlying Boshoek to Strubenkop Formations; Fig. 8.4-3), and was itself terminated by a relatively short, third-order stage of base level fall, responsible for the formation of the third-order subaerial unconformity separating lower and upper Daspoort systems tracts (Fig. 8.4-7). Thus, the lower Daspoort high accommodation systems tract can be compared to a TST plus HST. Renewed base level rise led to the upper Daspoort low accommodation systems tract, which is overlain by the transgressive Silverton epeiric marine facies (Fig. 8.4-3).

High-Frequency Sequence Stratigraphy Applied to Precambrian Successions Due to the poor preservation, and high levels of deformation, metamorphism and diagenesis common to many Precambrian sedimentary successions, there have been only a limited number of sequence stratigraphic analyses of early Precambrian basins (section 8.1). Most studies, that have been done have been low resolution and preliminary interpretations, as for example, that of Catuneanu and Eriksson (1999) on the 2.7-2.1 Ga Transvaal Supergroup. Error margins implicit even in accurate zircon dating have severely limited application of high-frequency sequence stratigraphy to such basins. This case study demonstrates that where geometry, lithology and sedimentary facies are reconstructed in detail,

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Chapter 8: Sequence Stratigraphy and the Precambrian

Fig. 8.4-8. Palaeogeographic reconstruction of the lower Daspoort depositional setting. Shaded area to the east marks the region of highest subsidence during the lower Daspoort time. This "eastern basin" accumulated fine sandstone, mudrock and ironstone facies under relatively high water table conditions, probably in a fluvial-lacustrine or shallow epeiric basinal environment. The rest of the Transvaal basin accumulated sandstone facies in fluvial systems draining towards the highest subsidence area.

thus enabling reliable depositional models to be constructed, it is indeed possible to apply sequence stratigraphy of higher detail, even in the near-absence of any geochronological constraints. Only one age has been determined within the Pretoria Group, a whole-rock Rb-Sr determination of c. 2.3 Ga for the Hekpoort Formation lavas (e.g., Eriksson and Reczko, 1995; Walraven and Martini, 1995). By rigorous application of the principles of sequence stratigraphy, and through a thorough understanding of the different such models

8.5. Commentary

735

already in use and thereby avoiding confusion among a plethora of terminology and usage (Catuneanu, 2002), higher frequency sequences can be defined. For the Daspoort case study, this was only possible once a basis of the well-defined second-order sequences for the Pretoria Group and entire Transvaal had been defined.

8.5.

COMMENTARY

O. CATUNEANU AND EG. ERIKSSON The base level changes implicit in the relatively new field of sequence stratigraphic studies are dependent upon a wide range of geological variables, discussed in previous chapters of this book. Generation of the continental crust (chapters 2 and 4), and specifically crustal growth rates (section 2.8), as well as the interplay of tectonism and mantle plumes (chapter 3) provide first-order controls on base level. These are complemented by second-order controls from palaeoclimatic (chapter 5), biological (chapter 6) and depositional influences (chapter 7). The supercontinent cycle (sections 3.2, 3.10 and 3.11) and the global glaciation concomitant with Palaeoproterozoic and Neoproterozoic assembly events (sections 5.6-5.8) are related directly to all the above-named variables. Sequence stratigraphy thus draws together the implications and inferences drawn from the many diverse fields of Precambrian geological investigation, and relates genetic processes directly to patterns which can be observed in the rock record. Practical Issues

Sequence stratigraphic models idealise reality in the sense that they provide simplified, theoretical two- or three-dimensional representations of how the architecture of sedimentary facies and stratigraphic surfaces is expected to be in the field. The central assumption of all models is that the predictable stacking pattern of systems tracts and stratigraphic surfaces is controlled mainly by the interplay of base level changes and sedimentation at the shoreline. This interplay controls the direction of shoreline shifts, and implicitly the timing of all systems tracts and bounding surfaces. Under this assumption, the unconformable portion of the depositional sequence boundary (subaerial unconformity) is the time equivalent of the falling stage systems tract, the maximum flooding surface has a predictable position above the subaerial unconformity, and so on (Fig. 8.2-4). Although these expected relationships are valid in most cases, especially in coastal regions, possible deviations from the model predictions should be evaluated carefully. For example, the influence of base level changes at the shoreline on fluvial processes only extends for a limited distance upstream (Shanley and McCabe, 1994). The extent of the base level control depends on the balance between the magnitudes of base level changes, climatic influences, and source area tectonism. There are instances where the role of climate is so dominant that processes of fluvial aggradation and incision are controlled mainly by changes in the balance between river discharge and The Precambrian Earth: Tempos and Events Edited by P.G. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu

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Chapter 8: Sequence Stratigraphy and the Precambrian

sediment load, with a timing that is offset relative to the base level fluctuations at the shoreline (Blum, 1994). The resulting subaerial unconformity therefore will not fit the position predicted by standard sequence models. There are also cases where a subaerial unconformity forms during transgression, in relation to processes of coastal erosion (Leckie, 1994). One other practical problem is the possible lack of preservation of systems tracts, or of portions of systems tracts. In this case, stratigraphic surfaces that are normally expected to be separated by strata may be superimposed. Examples include a ravinement surface that reworks the subaerial unconformity, a regressive surface of marine erosion that reworks the basal surface of forced regression, a maximum flooding surface that reworks the maximum regressive surface, or a subaerial unconformity that reworks the underlying maximum flooding surface. In these situations, the observed surface should be labelled using the name of the younger surface, as the latter overprints the attributes of the original contact. These are all practical problems that the practitioner of sequence stratigraphy may encounter in any case study, whether the succession under investigation is Precambrian or Phanerozoic in age. The sequence stratigraphic analysis obviously becomes more difficult with increasing stratigraphic age, due to additional limitations imposed by poor preservation, post-depositional tectonics, diagenetic transformations, metamorphism, and lack of biostratigraphic support. Nevertheless, where the geometry, sedimentary facies and facies relationships are well constrained, thus enabling reliable depositional models to be constructed, sequence stratigraphy can still be applied, even in the near-absence of geochronological constraints. This has been demonstrated for Precambrian successions in previous publications (Christie-Blick et al., 1988; Beukes and Cairncross, 1991 ; Krapez, 1996, 1997; Catuneanu and Eriksson, 1999, 2002; Catuneanu and Biddulph, 2001), as well as in the case studies included in this chapter (sections 8.3 and 8.4).

Systems Tracts The method of sequence stratigraphy originated, and has been traditionally applied, in basins where coeval fluvial to marine facies transitions are preserved. The observation of the direction and type of palaeo-shoreline shift (i.e., forced regression versus normal regression versus transgression) is crucial in applying the classic systems tract terminology of lowstand, transgressive and highstand deposits. This terminology is inappropriate for overfilled basins dominated by fluvial deposits, or where only the fluvial portion of the basin is preserved, because of the lack of control on the type and direction of shoreline shifts outside of the preserved basin. The study of fluvial depositional sequences, and their subdivision into systems tracts is a challenging and relatively new direction of research in sequence stratigraphy. In the absence of a preserved coeval shoreline, the fluvial succession cannot and should not be separated into lowstand, transgressive and highstand systems tracts. The objective alternative is to look at changes in fluvial styles and architectural elements (see section 7.8) and determine whether or not discernable packages with specific characteristics can be identified within a fluvial sequence. Previous research (e.g., Dahle et al., 1997; Boyd et al., 1999; Zaitlin et al., 2000) introduced the concepts of low versus

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737

high accommodation systems tracts for fluvial deposits, based on the fact that the amounts of available accommodation are always low in the early stages of base level rise, followed by subsequent increases. The application of the low versus high accommodation systems tracts is still in its infancy, and is perhaps more relevant for the study of Precambrian deposits since only portions of the Precambrian-aged basins are usually preserved. Hence, the chance of studying fluvial successions in isolation from their coeval shorelines is higher for increasingly older rocks. The two case studies presented in this chapter (sections 8.3 and 8.4) deal with dominantly fluvial successions, and exemplify the concepts of low and high accommodation systems tracts. This is the first attempt to apply these concepts to the Precambrian rock record, and new criteria are emerging for the recognition of these systems tracts in contrast to their Phanerozoic counterparts. The fluvial systems of the vegetationless Precambrian are dominated by unconfined braided and sheetwash facies (sections 7.1, 7.3, 7.6 and 7.8), which tend to replace the vegetated overbank deposits of Phanerozoic meandering systems. Under these circumstances, the ratio between channel and overbank architectural elements, used to separate the Phanerozoic low and high accommodation systems tracts, does not work as well when applied to Precambrian successions. Instead, changes in depositional trends, overall grading, and the geometry of fluvial deposits, may provide more useful criteria for the study of Precambrian sequences. The low accommodation systems tract corresponds to the stage of peneplanation in a developing basin, where fluvial deposits prograde and infill an immature landscape. This depositional trend may generate crudely coarsening-upwards profiles, and captures the coarsest sediment fraction of the fluvial sequence. Topographic irregularities above the sequence boundary give a potentially discontinuous geometry to this systems tract, with significant changes in thickness along dip and strike. The high accommodation systems tract has a more predictable and continuous geometry, either sheet- or wedge-like depending on subsidence patterns, and is dominantly aggradational. It corresponds to the stage of decline in the energy level of the fluvial systems, which may confer on it an overall fining-upwards profile. The boundary between the two systems tracts is potentially diachronous, younging in a dip direction, as opposed to the classic (lowstand, transgressive, highstand) systems tract boundaries, which are closer to time lines.

Hierarchy Systems One other important issue in sequence stratigraphy is how to design a hierarchy system that can encompass the relative importance of sequences and bounding surfaces in an objective manner. The existing hierarchy systems (section 8.2; Catuneanu, 2002, for a discussion) were proposed initially based on the study of Phanerozoic strata, and subsequently expanded to incorporate the Precambrian deposits as well (Krapez, 1996, 1997). One significant aspect that one has to bear in mind when dealing with the issue of hierarchy is that the span of time of similar tectonic processes-driven cycles changed through time, from the Precambrian to the Phanerozoic, in response to changes in the dynamics of plate tectonic processes (see section 3.6). This fact, generally overlooked, argues that a hierarchy system

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Chapter 8: Sequence Stratigraphy and the Precambrian

based on boundary frequency cannot possibly be applicable universally to the entire rock record. The reasonable alternative is to deal with the issue of hierarchy on a case-by-case basis, assigning hierarchical orders to sequences and bounding surfaces based on their relative importance within each individual basin. This method may prove to be more realistic, especially for ancient Precambrian basins, given the fact that each basin is unique in terms of formation, evolution, and history of base level changes.

The Precambrian Earth: Tempos and Events Edited by EG. Eriksson, W. Altermann, D.R. Nelson, W.U. Mueller and O. Catuneanu Developments in Precambrian Geology, Vol. 12 (K.C. Condie, Series Editor) 9 2004 Elsevier B.V. All rights reserved

739

Chapter 9

TOWARDS A SYNTHESIS RG. E R I K S S O N , O. C A T U N E A N U , D.R. N E L S O N , W.U. M U E L L E R A N D W. A L T E R M A N N

The principal theme of this book is change through time, or tempos and events in the Precambrian (Preface). Each chapter portrays a different part of the Earth's history but there is a unifying theme: Earth's evolution. Chapter 1 explains the celestial origin of our planet and the early development of the Earth into core, mantle, crust and primitive atmosphere. Chapter 2 discusses the generation of continental crust, with the emphasis on granite-greenstone terranes. Chapter 3 builds further upon its predecessor, emphasising the interaction between tectonism and mantle plumes through Precambrian time. Chapter 4 examines the volcanic attributes of the Archaean Earth and how they may have changed, as exemplified by plume-generated komatiites, the constant interaction between arc-plume volcanism and subaqueous caldera formation. Chapter 5 deals with the evolution of Earth's atmosphere and hydrosphere, and chapter 6 with related concepts of the evolution of Precambrian life and bio-geology. Chapter 7 details sedimentation regimes through Precambrian time, while chapter 8 discusses the application of sequence stratigraphy to the Precambrian rock record.

9.1.

EVOLUTION OF THE SOLAR SYSTEM AND THE EARLY EARTH

Investigation of pre-4 Ga Earth history relies largely upon study of the most ancient rocks thus far identified, and upon modelling of the differentiation of Earth's chemical reservoirs (Nelson, section 1.1). As the known preserved rock record dates from 4030 Ma (Stern and Bleeker, 1998; Bowring and Williams, 1999), more than 500 My of Earth's earliest evolution remains essentially speculative. It was only with the identification within meteorites of daughter products from radiogenic decay of long-extinct nuclides (firstly by Reynolds, 1960), that the timing of accretion and differentiation of the early Earth could be investigated (summarised by Nelson, in section 1.2). The short-lived parent nuclides were synthesised during supernova explosions shortly before formation of our solar system; their short half-lives enable precise determination of the chronology of the earliest history of the solar system (section 1.2). Collision and amalgamation of smaller, rocky planetesimals within a protoplanetary disk formed the terrestrial planets, including Earth. As proto-Earth and its Moon grew by these violent accretion processes, earlier differentiation products were largely obliterated; with the growth of embryonic planets the impact

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Chapter 9: Towards a Synthesis

rate decreased and concomitantly, the likelihood of preservation of fragments of the early Earth increased. Current evidence (section 1.2) suggests that short-lived nuclides with atomic masses < 140, together with a part of the heavy elements in our solar system, were synthesised during a core-collapse supernova event at c. 4571 Ma (Lugmair and Shukolyukov, 2001; Gilmour and Saxton, 2001 ). Formation of the Sun and solar system may have been initiated by shock waves from this supernova explosion (probably one of a number of successive such events); injection of short-lived nuclides into a nearby interstellar gas and dust cloud may have triggered its collapse, forming a proto-Sun of radius c. five times its present value, over a time period of < 105 years (Cameron, 1995; Nelson, section 1.2). Progressive collapse from inner to outer parts of the cloud, together with conservation of angular momentum, caused it to spin faster; colliding gas and dust particles orbiting the protoSun in the same direction lost their energy, causing flattening of the cloud, especially near its centre. Gravitational energy was converted to heat during collapse of the nebula. At some time during collapse, the density and temperature became high enough for hydrogen burning to commence, and the proto-Sun began its violent T-Tauri phase (Cameron, 1995; Nelson, section 1.2). More abundant Fe, Ni and silicate-rich components condensed within lower temperature parts of the nebula in its medial to central parts, while volatile elements (e.g., water, ammonia, methane ice) condensed in cold, outer parts of the accretionary disk. Volatiles were possibly carried by the solar wind from inner to outer reaches of the emerging solar system (Shu et al., 1994). Spectroscopy and simulation modelling suggest only a few million years from star formation and large scale accretion of disks into the young solar-type T-Tauri stars. Larger planetesimals may have formed within ~< 2 My of solar system formation (Hutchison et al., 2001). Coagulation consequent upon icy particle collisions within an ice sublimation belt in the cold outer parts of the nebula being more efficient than that between metal or silicate particles, large gas-rich proto-planets (Jupiter and Saturn precursors) formed before the nebula gas dissipated (Cameron, 1995). Collision and amalgamation of chemically refractory dust particles within the inner part of the disk occurred more slowly. Collisions of smaller planetesimals with larger bodies continued to rework early planetary-sized bodies for at least a further 100 My, and triggered large scale melting and magmatic differentiation of silicate components of the larger planetesimals. Dating of meteoritic remnants from these early differentiated planetary bodies indicates that planetesimals of at least 10s-100s of kilometres in diameter underwent internal magmatic differentiation within < 10 My after the supernova event. 187Re-187Os isotopic data from pallasites and iron meteorites (Morgan et al., 1995; Shen et al., 1998; Horan et al., 1998) suggest formation of metallic cores within c. ~< 50 My of formation of the solar system. There is intriguing evidence for hydrothermal alteration processes involving aqueous fluids within planetesimals ~< 2 My after solar system formation. Planetary embryos had thus existed within <~ 5 My of the supernova event that triggered formation of the solar system. Accretion of these embryos within a c. 0.5-2.5 AU range of the Sun (Wetherill, 1994) was largely responsible for formation of the terrestrial plan-

9.1. Evolution of the Solar System

741

ets, with major contributions of more volatile components from the asteroid belt and outer parts of the solar systems during later stages of terrestrial planetary formation. The impact of the Mars-sized planetary embryo "Theia" with the proto-Earth at c. 4550-4540 Ma (i.e., 25-35 My after solar system formation) is widely accepted as having formed the Moon (Giant Impact Hypothesis; see Stewart, 2000 and references therein; Kline et al., 2002; Yin et al., 2002; see also section 5.9). As a result of this collision, shock melting of at least one hemisphere of proto-Earth occurred, together with isostatic readjustment of the remaining planetary mass, followed over c. 1000 years by propagation of a solidification front towards the surface (Solomatov, 2000; section 1.2). A shallow, convecting magma ocean formed above the lower mantle, but within c. 1 My of impact, probably cooled rapidly to a more viscous, low melt fraction (20-30% partial melt) (Abe, 1993). This partially molten magma ocean in the upper mantle may have persisted for up to 200 My after the impact (Abe, 1997). The proto-atmosphere was likely severely depleted by the impact (Chen and Ahrens, 1997). 182Hf- 182W isotopic systematics facilitate understanding of the formation of the Earth's metallic core. Differences in W isotopic compositions between terrestrial silicates and carbonaceous chondrites suggest that Earth's metallic core formed within 35 My of formation of the solar system. However, planetesimals may have undergone significantly faster metal-silicate segregation, at higher rates than that of planetary accretion. The growth rate of proto-Earth's core was thus probably largely limited by the rate of planetary accretion, rather than by the rate of planetary differentiation and metal-silicate segregation. The short-lived nuclides suggest that Earth's accretion and differentiation into metallic core and silicate mantle was completed essentially within c. 20 My of the time of formation of the solar system. Rapid accretion of the terrestrial planets may have enabled capture of early atmospheres from the solar nebula. These early atmospheres may have been lost subsequently from intense ultraviolet luminosity and solar winds generated during the Sun's T-Tauri phase. However, early solar atmospheric gases could have been trapped by dissolution in the magma oceans on the planets prior to nebula dissipation, as supported by the relative abundances and isotopic compositions of the noble gases on Earth. It is possible that solar He, Ne and Ar were incorporated into Earth's magma ocean beneath a massive protoatmosphere of molecular hydrogen and helium prior to dissipation of the solar nebula, followed by acquisition of meteoritic compositions of the heavier noble gases by accretion of planetesimals after nebula-dissipation (Harper and Jacobsen, 1996). Processes responsible for the early differentiation of the Earth into metallic core, silicate mantle, chemically differentiated crust and atmosphere thus overlapped in time and operated in concert with progressive accretion of the planet. Due to the rapid formation of the Earth, its latter-stage collision with a small number of massive planetesimals, and its internal heating by radioactive decay and release of gravitational energy associated with core formation, Earth almost certainly passed through a "magma ocean" stage in its accretion history. The latter may have persisted for several 100 My; a concomitant steam atmosphere may have been lost due to impacts and hydrodynamic escape, with a significant proportion of the present atmosphere derived from

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Chapter 9: Towards a Synthesis

post-accretion comets (Delsemme, 1995; Nelson, section 1.2). Depending on atmospheric prevention of heat loss from the Earth's surface, a thin solid surface layer could have formed over the magma ocean within a period of tens of years (Mukhin and Pinenov, 2002). Early komatiitic and/or high-Mg basaltic terrestrial crusts, derived by very large degrees of partial melting of both shallow and deeper upwelling mantle sources, were probably melted, fragmented and reformed many times over the surface of a chaotically convecting, partially molten mantle (see also section 3.6). Such early Earth crusts were thus shortlived, commonly brecciated by meteorite impacts, and were rapidly recycled back into the mantle; periodic recycling on a global scale, as a consequence of both major impacts and of convective mantle overturn (Nelson, 1998a; see section 3.4) is inferred. Recycling mechanisms returning the earliest thin crust to Earth's mantle remain speculative (see discussion in section 3.6); certainly, higher mantle temperatures (and flood basalt resurfacing) during the Early Archaean will have resulted in substantially thicker and more magnesian oceanic crust than that of today (e.g., Arndt and Chauvel, 1990). Tonalite, trondhjemite and granodiorite melts (chapter 2) formed by partial melting within the early, hydrothermally altered mafic crust, and the oldest identified terrestrial zircons, 4400-4000 Ma old (Nelson, section 1.2), could have crystallised within these felsic differentiates. The chemical, microstructural and U-Pb isotopic characteristics of individual zircon crystals enable interpretation of a record of complex geological histories, and many ~> 4000 Ma zircons so far examined in detail have a wide range of 2~176 dates within each grain. Data from Australia indicate that at least some of the concordant 2~176 dates at c. 4360, 4340, 4320, 4185, 4150, 4005, 3978, 3945 and 3874 Ma determined within these ancient zircons may correspond to events during which temperatures of their host rocks exceeded 850~ (Nelson, section 1.2), possibly during episodic mantle upwelling or overturn episodes (see section 3.4). The oldest rocks known on Earth are 4030 Ma granitic gneisses of the Acasta Gneiss Complex of the Northwest Territories, Canada (Stern and Bleeker, 1998; Bowring and Williams, 1999). The earliest, best-preserved continental crust is comprised of a unique rock association characterised by linear or arcuate belts of predominantly mafic volcanic rocks (or greenstones) in fault contact with voluminous tonalitic, trondhjemitic, graniodioritic and/or granitic (TTG) rocks. This "granite-greenstone" crust is largely unique to the Archaean era (see also chapters 2-4). Pillow lavas within the Isua greenstone belt (sections 2.2 and 2.3) are clear evidence for submarine magmatism (and oceans) at c. 3.7 Ga (Myers, 2001a, b). Although granite-greenstone crust formation within the Pilbara and Kaapvaal cratons was episodic (c. 10-100 My duration; Nelson et al., 1999), their overall chronological patterns are not similar (Nelson et al., 1999), suggesting that early Archaean crustal formation was mostly associated with localised processes rather than global-scale convective overturn (cf. Nelson, section 3.4). During Earth's earlier Precambrian history, impact events continued to be important; the lunar record indicates that the impact rate in the Earth-Moon system exceeded the present rate by about 15 times at 3.8 Ga, declining to about 2 times by 3 Ga (Ryder, 2003). Between 3.8 and 2.5 Ga, it is inferred that more than 350 impact events occurred which were large enough to form global-scale spherule (sand-sized silicate droplets formed by

9. 2. Generation of Continental Crust

743

the melting and vapourisation of terrestrial target rocks during asteroid and comet impacts) layers (Abbott and Hagstrum, section 1.4), although terrestrial impact structures older than c. 2.5 Ga are unlikely to have survived (Simonson et al., section 1.3). Major magmatic (cf. crust-formation) events as well as strong plumes (cf. komatiites; see also sections 4.3, 3.2 and 3.3) may have been related to major impact events (Abbott and Hagstrum, section 1.4). Eleven spherule-rich impact layers have been identified for the Archaean-Palaeoproterozoic period, ten of these are from the Kaapvaal and Pilbara cratons (Simonson et al., section 1.3). These distinctive spherule layers reflect not only a direct impact genesis, but were also subject to reworking by tsunami waves, concomitant high energy currents and large scale aeolian reworking. Archaean spherule layers have been dated at 3470 -+- 2 Ma, c. 3260 Ma and 3243 -4- 4 Ma, with a fourth layer being close to the latter (Byerly et al., 1996, 2002; Lowe et al., 2002). The ages of the Neoarchaean-Palaeoproterozoic layers are less well known, three layers in the Hamerley basin are between c. 2630 and 2490 Ma. The recurrence interval with both groups above is c. 70-77 My (section 1.3). It is inferred that most early Precambrian spherule layers reflect impactors of at least K/T boundary size, and such impacts were probably more closely spaced in time than the c. 70 My interval inferred from the present data base. Early Precambrian spherules appear to have been, on average, more basaltic than Phanerozoic equivalents (Simonson and Harnik, 2000), a viewpoint consistent with continental crustal growth rates (sections 2.8 and 3.6; chapter 2) and a higher incidence of early Precambrian impacts into oceanic crust (sections 1.3 and 1.4). Before 3 Ga, continental crust may have been only about 20% of present volume (covering c. 7% of Earth), and typical Archaean crust (c. 49 km) was thicker than its Phanerozoic (c. 40 km) equivalent; by 2.5 Ga, Archaean continents had attained c. 80% of their modern volume (covering c. 27% of Earth; present-day value is about 41%) (section 1.4).

9.2.

GENERATION OF CONTINENTAL CRUST

The formation of granite-greenstone terranes has invoked a range of diverse ideas about the development of continental crust on Earth. Views on crustal growth rates are equally divergent, some arguing for episodic growth from the Early Archaean (Taylor and McLennan, 1985; McLennan and Taylor, 1991; Condie, 1998), in contrast to Armstrong (1991) who suggests that most of the continental crust was generated prior to 3 Ga, followed by loss due to recycling into the mantle (Dimroth, 1985; Nielsen et al., 2002) (Arndt, section 2.8, for summary). A major global peak in continental crustal growth occurred at 2.7 Ga, demonstrated by the development of voluminous juvenile crust or thermal overprinting of pre-existing crust on every continent. Lesser peaks in this growth rate are only regional in extent: 2.5 Ga (China and India), 2.1 Ga (West Africa and South America) and 1.8-1.9 Ga (North America and Australia) (Arndt, section 2.8). This suggests the possibility of tectonically active periods, related to the peaks, interspersed with periods of rel-

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Chapter 9: Towards a Synthesis

ative tectonic (specifically subduction) stasis, an idea also interpreted from global carbon isotopic records (see section 5.3). Geochemical investigations of Archaean greenstone belts indicate great compositional diversity in volcanic rocks, which in turn support diverse source characteristics, petrogenetic processes, and tectonic settings in the early Earth (Arndt, 1994; Polat and Kerrich, 2001b; Wyman et al., 2002a; Polat et al., 2002; Polat et al., section 2.3). The > 3.7 Ga Isua greenstone belt, the oldest known on Earth (Rosing et al., 1996; Appel et al., 1998; Fedo, 2000) is, as expected, the subject of controversial and diverse interpretations. Detailed mapping (Myers, 200 l a) suggests that several fault-bounded litho-tectonic sequences consist mainly of basaltic and high-MgO basaltic pillow lavas, ultramafic intrusions, chert-BIF, and minor clastic sedimentary rocks (Myers, section 2.2). Polat et al. (section 2.3) recognise two distinct types of mafic to ultramafic volcanic associations within structurally separated sequences of the Isua belt: (1) a "boninitic" one (Polat et al., 2002), and (2) one with geochemical characteristics similar to those of Phanerozoic island arc picrites. Volcanic rocks of boninitic affinity recently reported from the Neoarchaean Abitibi and Frotet-Evans greenstone belts (Superior Province, Canada) suggest that this type of volcanism may have been more widespread in the Archaean than currently recognised (Kerrich et al., 1998; Boliy and Dion, 2002). If the geochemical characteristics of the inferred Isua boninites and picrites have the same geodynamic significance as Phanerozoic counterparts, it is possible that Phanerozoic-like plate tectonic processes were operating as early as 3.8 Ga (e.g., Nutman et al., 2002); their operation in the Neoarchaean is much more generally agreed upon (see also discussion in section 3.6). However, recent field studies of the ancient, highly deformed Isua terrane suggest that what some workers interpret as the oldest intra-oceanic accretionary complex, may not be of primary origin (Myers, section 2.2). The contentious issue of forming large amounts of continental crust is directly related to mantle plumes (superplume events; sections 3.2 and 3.3) and subduction processes (Mueller and Nelson, section 2.1). Mantle plumes provide both under- and overplating (cf. flood basalts), whereas horizontal plate movements related to subduction may generate much of the volcano-plutonic material at active continental margins (where it is often accreted) and mid-ocean ridges (Fisher and Schmincke, 1984). Large-scale granitic diapirism within early Archaean granite-greenstone terranes may have resembled in some way~ plume upwelling within the mantle. There is little evidence of predominant vertical tectonics controlling late Archaean greenstone belt evolution. However, such tectonics (a modification of the "sagduction theme") is supported by some for parts of the Pilbara (Hickman and Van Kranendonk, section 2.6) and Indian cratons (Chardon et al., 1998). Diapirism was possibly more significant in the formation of granite-greenstone crust in the Early Archaean, and may have diminished with time. In contrast, data from the Neoarchaean cratons of North America support coeval plumegenerated komatiites and subduction-related arc volcanism, with good evidence for plumearc interaction over c. 30 My being recorded in the Abitibi greenstone belt (Dostal and Mueller, 1997; Mueller and Mortensen, 2002; Daigneault et al., section 2.4). In the eastern part of the Yilgarn craton, an elongate, c. 2705 Ma rift was filled rapidly by komatiitic

9. 2. Generation of Continental Crust

745

and basaltic lavas; c. 2670 Ma subduction and continent-continent collision terminated rift volcanism and initiated regional deformation, followed by widespread granitoid plutonism (Mueller and Nelson, section 2.1). Comparing the tectonic evolution of Archaean greenstone belts with Phanerozoic counterparts has long been a contentious issue (e.g., Hamilton, 1998). Although plate tectonics is accepted by many workers as the principal influence on Archaean greenstone belt formation and deformation (e.g., Langford and Morin, 1976; de Wit, 1998; see, however, section 3.6 for more detailed discussion), plume activity for komatiite-tholeiitic basalt sequences remains an important process (Tomlinson and Condie, 2001 ). The 2735-2670 Ma Abitibi belt, the largest coherent greenstone belt in the world (Card, 1990) displays strong evidence for arc formation, arc evolution, arc-arc collision, and arc fragmentation with the recognition of strike-slip basins (Mueller et al., 1996); it is thus strikingly similar to modern collisional orogens (Daigneault et al., section 2.4). Exhumation in Archaean greenstone belts can also be explained by plume upwelling, with the possibility of plume-driven extensional structures (containing significant komatiites) forming during terminal stages of arc evolution (Daigneault et al., section 2.4). The genesis of granitoid rocks, a major component of all Archaean terranes, is a significant part of the formation of granite-greenstone terranes. Emplacement of granites, either by uniquely Archaean diapiric processes (Hickman, 1984; Choukroune et al., 1995; Collins et al., 1998; Hickman and Van Kranendonk, section 2.6), or by far-field induced deformation related to a plate tectonic regime, during compression (de Wit et al., 1987a; Bickle et al., 1993), extension (Zegers et al., 1996), or strike-slip deformation (Zegers et al., 2001) (Zegers, section 2.5), is still a hotly debated topic. Although some, mostly Early Archaean cratons (e.g., Pilbara, Kaapvaal and Zimbabwe) are characterised by composite, ovoid granite batholiths surrounded by volcano-sedimentary greenstone sequences, most Late Archaean granite-greenstone terranes consist of elongate alternating belts of granites and greenstones (e.g., the Neoarchaean Yilgarn craton and Superior Province). Within any specific granite-greenstone terrane, there is a secular change from tonalite-trondhjemite-granodiorite (TTG) suites, to granodiorite-granitemonzogranite (GGM) suites (both typically pre- to syntectonic), to the highest-K20 syenite-granite (SG) suites (typically post-tectonic) (Bickle et al., 1989; Feng and Kerrich, 1992; Bickle et al., 1993; Zegers et al., 1998b; Zegers, section 2.5). Field relationships suggest that pre- to syntectonic granites intruded originally as sub-horizontal sheets into ovoid and linear granite-greenstone terranes (de Wit et al., 1987a; Chown et al., 1992; Zegers et al., 1996; Collins et al., 1998; Kloppenburg et al., 2001). The production of large volumes of TTG melt is the first and essential step in forming Archaean continental crust (Zegers, section 2.5). Two subgroups of TTGs have been recognised (Barker and Arth, 1976): (1) a more common, high-Al series, reflecting generation by partial melting of hydrated basalt in the high-pressure garnet stability field (Rapp, 1997; Wyllie et al., 1997) and (2) a low-Al series, generated under lower pressures in the plagioclase stability field (section 2.5). The higher Archaean heat production, estimated at 2-6 times present values (Pollack, 1997) and an oceanic crust analogous to very thick,

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modern oceanic plateaus (Kusky and Kidd, 1992; Condie, 1997a; Polat et al., 1998) were pertinent to formation of TTG (Zegers, section 2.5). TTGs are generally regarded as reflecting partial melting of hydrated basalts, with significant recycling of residual, most probably eclogitic, mafic lower crust back into the lithospheric mantle. Two geodynamic models can be considered: (1) shallow subduction of a thick and hot oceanic lithosphere (Martin, 1986; Drummond and Defant, 1990; Davies, 1992a; Martin and Moyen, 2002); (2) in situ crustal differentiation and delamination (Glikson, 1972; Anderson, 1979; Kr6ner, 1985a; Vlaar et al., 1994; Zegers and van Keken, 2001). Although mantle plumes were probably important in Archaean geodynamics, "plume tectonics" cannot provide a general model tbr the production of TTG melts and for the recycling of eclogites (Zegers, section 2.5). If the subduction and slab-melt model is pertinent, then the first stable continental crust may have comprised a complex and deformed array of accreted oceanic crust/plateaus and volcanic arcs of TTG composition (Calvert and Ludden, 1999). Alternatively, applying the in situ delamination model, a relatively simple and superficially little-deformed initial continental crust, consisting of a refractory mafic and gneissic lower crust, a gneissic middle crust of mixed TTG material, amphibolite and residual material from melt extraction, and an upper crust of TTG laccoliths and the upper parts of the oceanic plateaus, basalts and gabbro sills, might be expected (Zegers, section 2.5). Subsequent to initial Early Archaean crust formation, and prior to final stabilisation, large volumes of granite (GGM and SG, above) formed largely by melting of pre-existing TTG crust. Emplacement probably occurred as subhorizontal sheets (laccoliths), consistent with the generally tabular geometry of Archaean granitic plutons. TTG melt is thought to have migrated as structurally controlled feeder dykes, with final intrusion as flat-lying sheets either during extension (in terms of the eclogite-delamination model) or during compression (applying the flat-subduction model) (Zegers, section 2.5). As successive granite sheets intruded under or directly above existing TTG intrusions, a domal geometry would have formed during the progressive intrusion of granite; such domal geometry may then have been subjected to variable regional deformation, with intense compression and thrusting likely producing a linear fabric (common in most Neoarchaean terranes). Formation of core complexes may also have enhanced the domal geometry (Zegers et al., 2001). Zegers (section 2.5) thus stresses that a domal structure of ovoid granite-greenstone terranes should be interpreted as a lack of intense compressional deformation after granite intrusion, rather than reflecting a unique Archaean diapiric emplacement mechanism. Hickman and Van Kranendonk (section 2.6) note that granitoid-domal structures separated by synformal, keel-like greenstone structures are common in many cratons (e.g., Zimbabwe, Kaapvaal, Dharwar and parts of the Pilbara, Yilgarn and Superior), but emphasise that opinions differ as to their origin. They examine in detail the 3.52-2.83 Ga East Pilbara granite-greenstone terrane of the Pilbara craton, Western Australia, one of the best examples of such an Archaean dome-and-basin pattern. They explain this geometrical pattern as reflecting gravity-driven deformation concomitant upon the overturn of low-density rocks (granitoids) that have been buried under denser rocks (greenstone cover). Their suggested tectonic model is comparable to diapiric models applied

9.3. Tectonism and Mantle Plumes Through Time

747

to other such dome-and-basin terrains (e.g., Macgregor, 1951; Huddleston, 1976; Drury, 1977; Stephansson, 1977; Fyson et al., 1978; Gorman et al., 1978; Glikson, 1979; Gee, 1979; Mareschal and West, 1980; Borradaile, 1982; Anhaeusser, 1984; Bouhallier et al., 1993; Jelsma et al., 1993; Chardon et al., 1996; Choukroune et al., 1997). Since the advent of plate tectonic models in the past 30-odd years, such as that discussed above (Zegers, section 2.5), diapiric models have come to be viewed with some scepticism (see, however, Van Kranendonk et al., 2002, in press). However, it may be that diapirism played a more significant part in formation of some cratons (i.e., Pilbara), whereas in others (e.g., Superior, Yilgarn) evidence of this process is less convincing. Geodynamic models of Early Archaean crustal evolution, discussed above, are thus based either on comparisons with Phanerozoic plate collisions, island arcs or metamorphic core complexes (e.g., de Wit, 1991, 1998; Kr6ner and Layer, 1992; Zegers, 1996; Blewett, 2002; Sugitani et al., 2002), or on solid-state diapirism, crustal delamination, or mantle plume activity (e.g., Campbell and Griffiths, 1992; Choukroune et al., 1995; Collins et al., 1998; Hamilton, 1998; Zegers and Van Keken, 2001). In section 2.7, Nijman and de Vries focus on Archaean sedimentary basin dynamics, based on the premise that basin architecture provides an important link between sedimentary and deformational records of crustal evolution. They postulate generation of the first sedimentary basins as major, possibly ring-shaped, crustal collapse structures, analogous to the coronae on Venus. These earliest (> 3.3 Ga) terrestrial sedimentary basins, filled by volcanic rocks and cherty sediments, were controlled by normal listric growth faults (arranged in non-linear patterns) that linked shallow felsic intrusions with the surficial basin-fills. The inferred extensional tectonics bore no relationship to compression and crustal thickening, nor to present-day geometrical arrangements of granitoid bodies and greenstone belts; instead, Nijman and de Vries (section 2.7) advocate crustal uplift, collapse and basin formation by crustal delamination and related plume activity. In the Pilbara and Kaapvaal cratons, this phase of extensional crustal evolution came to an end at c. 3.3 Ga, to be replaced by compression tectonics and concomitant clastic sedimentation more readily attributable to plate motion; a comparable transition may have taken place 200 My earlier in the Isua greenstone belt (section 2.7).

9.3.

TECTONISM AND MANTLE PLUMES THROUGH TIME

As already noted in section 9.1, Hadaean geological evolution on Earth remains largely speculative. Trendall's (2002) "plughole" model postulates a gradual transition from c. 4.0-2.5 Ga, from whole mantle convection within an Earth dominated by thermal processes to layered mantle convection and the increasing dominance of plate tectonics (section 3.6). The earliest (probably localised) gneissic and sialic protocratonic nuclei possibly began to form at c. 4 Ga, by predominantly thermal-magmatic processes (Trendall, 2002). What is less speculative, is that formation of granite-greenstone crust (section 1.2, chapter 2; section 9.2 above) and the occurrence of komatiitic lavas are essentially Archaean (section 3.4). Although controversial, identified Archaean ophiolites suggest genesis through thickened ocean plateaus together with significantly higher heat flow (Moores,

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2002; section 3.7). Wide acceptance of 2-3 times (modern-Phanerozoic) mantle heat flow in the Archaean supports a more chaotic mantle convection regime, implicit in the Trendall (2002) model; furthermore catastrophic magmatic events were probably significant during crustal growth in the Archaean (Nelson, section 3.4). Many (perhaps most) Archaean cratons are inferred to have been underlain by low Rb-Sr-type metasomatised lithospheric mantle. This may reflect mantle enrichment due to fluid release during shallow subduction, which may have been reasonably common (Zegers, section 2.5), at least in the Neoarchaean (Cousens et al., section 3.5) when plate tectonism had probably become an important geodynamic process. Shallow subduction, applied to an earlier onset of plate tectonics than envisaged in Trendall (2002), could explain, partially, the operation of Archaean plate tectonics (de Wit, 1998). Inferred ophiolites > 1 Ga indicate temporal change in Earth's geothermal gradient, increasing mantle heterogeneity, and thinning of ocean crust which possibly also promoted the onset of conventional plate tectonics (Chiarenzelli and Moores, section 3.7). The interplay of plate tectonics and thermal processes (cf. mantle (super)plumes and their products, large igneous provinces, "LIPs") continued throughout the Precambrian (and later). A superplume is defined variously, as encompassing either buoyant material rising through the mantle irrespective of depth of origin (Ernst et al., section 3.3), or rising from the deep mantle (Condie, section 3.2). The LIP record appears relatively constant, a continental LIP forming at about every 20 My from 2.5 Ga onwards, and there is an inferred weak cyclicity (at c. 170, 330 and 730-600 My), except for possible gaps at 615-720, 2220-2400 and 3000-3300 Ma (section 3.3). Mints and Konilov (section 3.9) also note quiescent within-plate geodynamic processes from 2.44-2.0 Ga. Decreased plume frequency prior to 2.8 Ga may be an artifact of data analysis (Ernst et al., section 3.3). A major change in Earth's evolution occurred in the Neoarchaean: (1) LIPs and their mantle superplume progenitors, increase in frequency at 2.8-2.7 Ga (Condie, 2001 a; Ernst and Buchan, 2002a, b; sections 3.2 and 3.3); (2) catastrophic mantle overturn events became global in scale at c. 2.7 Ga (Nelson, section 3.4); (3) transition to a plate-tectonicallydominated Earth was evidenced by large volumes of granite-greenstone crust formed on the Yilgarn and Superior cratons at 2760-2620 Ma, including major 2705 Ma komatiite eruption (Nelson, 1998), and by evidence for c. 2.7 Ga ophiolites in the Limpopo orogenic belt in southern Africa (section 3.8); (4) the first supercontinent is postulated at c. 2.7 Ga ("Kenorland"; e.g., Aspler and Chiarenzelli, 1998; see also section 3.9); Neoarchaean greenstones with oceanic plateau-type geochemistry appear to have been major contributors to this supercontinent (Condie, 1994b; Tomlinson and Condie, 2001), with pre-2.7 Ga crustal fragments, and, possibly also, ocean arc systems (sections 3.2 and 3.6). There is a close link between the supercontinent cycle and magmatic processes. Precambrian ophiolite complexes, which cluster in time at c. 1-1.5, 1.8-2.3, 2.5-2.7 and at c. 3.4 Ga, may have occurred during the early assembly of supercontinents (Chiarenzelli and Moores, section 3.7). The first such assembly event may reflect the first catastrophic slab avalanche event at c. 2.7 Ga, as plate tectonics became significant and as slabs possibly reached a critical mass at the 660-km mantle discontinuity; slab avalanching into

9.4. Precambrian Volcanism

749

the lower mantle may have triggered the first superplume event (e.g., Peltier et al., 1997; Condie, 1998). Major superplume events would have been associated with supercontinents close to their terminations; the two major such events at c. 2.7 and 1.9 Ga were associated with globally elevated sea levels (chapter 8), peaks in stromatolite (section 6.5) occurrence and diversity, and significant changes in ocean chemistry (Condie, section 3.2; Ohmoto, section 5.2). The c. 2.7 Ga Ventersdorp continental flood basalts are ascribed to a direct plume hit on the Kaapvaal craton, which led to high freeboard (section 7.1) with no evidence for global eustatic rise (Eriksson et al., 2002b). Many Palaeoproterozoic high-grade mobile belts reflect a magmatic origin, with major superplumes affecting Fennoscandia at c. 2.52-2.44 Ga, and a more widespread superplume event at c. 2-1.95 Ga (Mints and Konilov, section 3.9). Formation of the c. 1.2 Ga supercontinent (Rodinia; various appellations and configurations are hotly debated; sections 3.10 and 3.11) was followed by breakup, which began at c. 750 Ma around the Kalahari craton, related to a thermal mantle anomaly (Frimmel, section 3.10). Rapid post-breakup motion of large tectonic plates may result from supercontinental blanketing of mantle heat (Gurnis, 1988; Gurnis and Torsvik, 1994; Honda et al., 2000), augmented by plumes forming 200-400 My after assembly (Meert and Tamrat, section 3.11). The plumes provide a mechanism to enhance post-breakup velocities of plates with thick tectospheres (Gurnis and Torsvik, 1994; Honda et al., 2000). High plate dispersal velocities, such as those inferred for Rodinia, can also be augmented by supercontinental fragments drifting from long wave length geoid highs towards cold spots (cf. geoid lows; Condie, section 3.2) resulting from subduction (Meert and Tamrat, section 3.11).

9.4.

PRECAMBRIAN VOLCANISM, AN INDEPENDENT VARIABLE

Volcanic rocks are significant components of Precambrian greenstone belts, and their genesis is related commonly to plate subduction systems and mantle plumes (Mueller and Thurston, section 4.1). During the Archaean-Palaeoproterozoic period, plume-generated komatiites (Campbell et al., 1989; Abbott and Isley, 2002b) and boninites, adakites and tonalite-trondhjemite-granodiorite (TTG) plutonic suites reflected low-angle subduction tectonics (Chown et al., 1992; Polat et al., 2002; Wyman et al., 2002a; see also sections 2.5, 3.5, 9.2 and 9.3). Komatiites formed subaqueous oceanic plateaus or islands, but also penetrated stable > 2.8 Ga continental crust, and affected oceanic arc sequences, as indicated by interaction with arc volcanism (Mueller, section 4.7). From the Mesoproterozoic, there is a change to high-angle subduction and an absence of komatiite-generated magmas. Precambrian volcanism may be viewed as an independent variable, influenced by mantle and crustal processes, and reflecting geodynamic change in the Earth (chapter 4). Modern explosive arc eruptions and Phanerozoic plume-controlled continental volcanism have Archaean and Proterozoic counterparts (section 4.1). The geodynamic setting of many Archaean grecnstone belts has modern counterparts, but Archaean oceanic assemblages and ophiolites are rare (Williams et al., 1992; Sylvester et al., 1997). Identification

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of Archaean oceanic crust and/or ophiolites is contentious (e.g., Thurston, 1994; Bickle et al., 1994; Kusky and Winsky, 1995; Sylvester et al., 1997), partly because of interpretations of high strain, and of structural contacts with basement as d6collements (Kusky and Kidd, 1992; Kusky and Winsky, 1995) (see detailed discussion of Precambrian ophiolites by Chiarenzelli and Moores, section 3.7). Palaeo-atmospheric composition (chapter 5) would not have affected the eruption mechanism or the transport process, nor would the early Precambrian hydrosphere have changed subaqueous pyroclastic transport mechanisms. Logically, the fundamental difference between the Precambrian and the modern Earth lies in the volume of magma generated at mid-ocean ridges and convergent plate margins. The abundance of aphyric tholeiitic basalts, characterised by large and abundant altered plagioclase spherulites, in numerous Archaean sequences supports the concept that Archaean thermal regimes were different to those of later Eons (Arndt and Fowler, section 4.3.3). The volume of effusive volcanism at oceanic ridges was probably higher during the Archaean (Mueller and Thurston, section 4.1; see also section 3.6). Archaean effusion rates were probably also higher because of the higher geothermal gradient and as a result of rapidly colliding microplates (Bickle, 1978, 1986; Sleep and Windley, 1982; Galer and Metzger, 1998; see also section 2.7; however, cf. with section 3.6). However, mass balance calculations (Dimroth, 1985) suggest rates of magma emplacement for the Abitibi greenstone belt that were similar to those calculated for MesozoicCenozoic arcs; in addition, the c. 80 km spacing between Abitibi arc edifices is similar to that of modern arcs (Windley and Davies, 1978). Arc systems thus appear not to be the best parameter to use for discerning a change in volcanism through time. Inferred higher Archaean-Palaeoproterozoic temperatures (sections 3.6, 9.2 and 9.3) would only cause magma-generation at shallower levels; boninites and adakites may be the response to shallow subducting plates (sections 2.5 and 3.5), as they require high heat flow and rapid subduction of young oceanic crust (Kerrich et al., 1998" Leybourne et al., 1999" Komiya et al., 2002; Wyman et al., 2002a). As Proterozoic arc volcanic assemblages have a higher proportion of volcaniclastic rocks and exhibit evidence for more shallow water and subaerial volcanism than Archaean arc volcanoes, Archaean volcanism may generally have occurred in deeper water than Proterozoic volcanism (Condie, 1994a). This may reflect higher rates of isostatic subsidence due to volcano loading on initially relatively thin and weak lithosphere. The latter thickened and strengthened over time, with an important factor being decreasing thermal regimes in the mantle (Richter, 1985; Thurston and Ayres, section 4.4). A difference in the Precambrian volcanic record relates to a dearth of orogenic andesites in Archaean arc sequences. Abbott and Hoffman (1984) related the inferred hotter Archaean Earth to low-angle subduction of young, hot oceanic crust, resulting in more siliceous melts (Helz, 1976) and bimodal volcanism. A "cooler" mantle and concomitant high-angle subduction favours the formation of orogenic andesites (Gill, 1981). A more profound temporal difference is provided by komatiites (Viljoen and Viljoen, 1969a, b), inferred to originate from mantle plumes (Campbell et al., 1989), and which are abundant in the Archaean and absent in the present (see also sections 3.2-3.4). This has major impli-

9.5. Evolution of the Hydrosphere and Atmosphere

751

cations for Earth's temporal evolution. Plume-generated volcanism was more voluminous during the Archaean and Early Proterozoic, as inferred from superplume events (Nelson, 1998; Abbott and Isley, 2002). Precambrian superplume events between 1.7 and 2.9 Ga (sections 3.2 and 3.3) possibly resurfaced Earth completely (Abbott and Isley, 2002b), with magma volumes ten times larger than Phanerozoic counterparts. Komatiites are subdivided into (1) Al-depleted (Barberton-type) flows, derived from greater depths with a lower degree of melting, and (2) Al-undepleted (Munro-type) flows originating from shallower levels with a higher degree of melting (Dostal and Mueller, section 4.3.2). The latter komatiite type is prominent in younger, c. 2.7 Ga Archaean greenstone belts (e.g., Abitibi; see also section 2.4; and Belingwe in Zimbabwe), and uncommon in pre-3.0 Ga belts. In contrast, Al-depleted komatiites are common in 3.5-3.0 Ga greenstone belts (e.g., Barberton, South Africa; and those of the Pilbara craton, sections 2.5-2.7). Both of these two komatiite types were probably generated by a high degree of melting from a mantle plume (sections 3.2 and 3.3) composed of garnet peridotite (Dostal and Mueller, section 4.3.2). Volcanic rocks in Proterozoic greenstone belts are generally interpreted as subductionrelated, mantle-derived, juvenile island arcs and back-arcs formed within minor spreading centre and plume-related oceanic environments (Thurston and Ayres, section 4.4). Although this inferred tectonic setting is comparable to modern equivalents, three key elements of modern plate tectonics first appear in the Neoproterozoic Pan-African orogen (e.g., Engel et al., 1980; Stern and Abdelsalam, 1998; section 4.4): (1) widespread ophiolites tectonically interspersed with deformed and accreted island arc sequences (Abdelsalam and Stern, 1996), (2) m61anges (Shackleton, 1994), and (3) possible blueschist facies metamorphism (De Souza Filho and Drury, 1998). Proterozoic greenstone belt volcanism exhibits peaks at 2.2-2.1, 1.9 and 1.3 Ga (e.g., Condie, 1994b, 1995), in contrast to the more continuous volcanism found on cratons (e.g., Melezhik and Sturt, 1994). Volcanism in Palaeoproterozoic greenstone belts typically occurred over periods of 30-95 My, to be succeeded by orogenic culmination less than 100-150 My after the first widespread volcanism (Lucas et al., 1996; Nironen, 1997; Hirdes and Davis, 2002). In Neoproterozoic greenstone belts, volcanism lasted for c. 200 My and terminal orogenesis took place 220-300 My after the onset of volcanism (Stern, 1994; Stein and Goldstein, 1996; Blasband et al., 2000; Thurston and Ayres, section 4.4).

9.5.

EVOLUTION OF THE HYDROSPHERE AND ATMOSPHERE

The intimate relationship between life, atmospheric and ocean chemistry on Earth (Ohmoto, section 5.2) allows biogeochemical signatures (e.g., carbon and sulphur isotopes) within the sedimentary record to be used to study Earth's early biosphere and atmosphere (Knoll and Canfield, 1998; Ohmoto, section 5.2). One of the most important parameters is the palaeoredox state of our planet's atmosphere. In the long term, almost all of the organic matter resulting from photosynthesis is decomposed upon exposure to the atmosphere and

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surface water, through several pathways; thus, all atmospheric oxygen is renewed about every 3000 years (section 5.2). The resultant negative feedback mechanism not only partially controls atmospheric oxygen, but also influences concentrations of greenhouse gases (mainly CO2 and CH4; Ohmoto, section 5.2). If long term oxygen production and consumption fluxes are not in balance, accumulation or loss of atmospheric oxygen will occur. Consumption occurs through oxidation of reduced volcanic gases (e.g., as H2, HzS, SO2, CH4 and CO) and of fossil carbon in sediments during pedogenesis (Holland, 1978). Long term oxygen production (over time scales > 3000 years) results from removal of small amounts of organic matter from the decomposition part of the Corg-O2 cycle during burial of marine sediments (Lasaga and Ohmoto, 2002). The same set of (largely equivocal) geological, palaeontological and bio-geochemical data is interpreted in favour of two mutually exclusive models (discussed in detail by Ohmoto, section 5.2) of atmospheric and oceanic chemical evolution (major uncertainty exists whether these data reflect original conditions or subsequent alteration): (1) the Precambrian fossil record; (2) minerals (e.g., uraninite, pyrite) within the early Precambrian record, which are unstable in modern fluvial sedimentary environments; (3) the behaviour of Fe in subaerial and marine environments (especially iron-formations); (4) the sulphur cycle and evolution of sulphur-utilising bacteria, and (5) the carbon cycle. The first, "C-W-H-K" (e.g., Cloud, 1968; Walker, 1977; Holland, 2002; Kasting and Siefert, 2002) model assumes the origin of life under a reducing atmosphere, a small rise of O2 at 3.0 or 2.8 Ga (Rye and Holland, 2000; Kasting and Siefert, 2002) and a major rise at c. 2.0 Ga, enabling emergence of the eukarya. Biogenic methane (with subordinate CO2) is inferred to have counteracted the "faint young Sun" before c. 2.2 Ga, CO2 levels of 300 PAL and < 0.1 ppm O2 at c. 3 Ga are postulated, and anoxic oceans (except for the photic zone) are implied until c. 600 Ma. Except for local evaporitic basins, low SO 2- levels are inferred until c. 2.2 Ga, with a gradual increase until c. 0.8 Ga when a second step-wise increase to modern levels occurred. Before c. 1.8 Ga, global oceans had high Fe 2+ and low HzS, with the reverse relationship from c. 1.8-c. 0.8 Ga; at c. 0.6 Ga the oceans became SOl--rich and HzS-poor (Walker and Brimblecomb, 1985" Bjerrum and Canfield, 2002; section 5.5). The second, "D-O" (e.g., Dimroth and Kimberley, 1976; Lasaga and Ohmoto, 2002) model postulates emergence of oxygenic photosynthesis and possibly of cyanobacteria (however, their fossils only appear in the Neoarchaean) soon after initial differentiation of oceanic and continental lithosphere at c. 4 Ga, followed by a single rise in atmospheric 02 from < 103- to c. 1 PAL soon after 4 Ga; thereafter relatively constant pO2 levels within 50% of PAL are inferred. Furthermore, low CH4 levels, and CO2 as the primary greenhouse gas are implied, as well as oxygenated oceans with essentially constant SO 2- levels since c. 4 Ga; Fe z+ and HzS in normal oceans remained low from c. 4 Ga. Elevated greenhouse gas contents in Earth's early palaeo-atmosphere would have raised weathering rates, while higher geothermal activity and concomitant acidic waters would have promoted aggressive breakdown of rocks (Corcoran and Mueller, section 5.11 ). However, palaeoclimatic reconstruction from Precambrian palaeosols and sedimentary rocks

9.5. Evolution of the Hydrosphere and Atmosphere

753

remains difficult, due to the interplay between chemical weathering and mechanical removal, and due to diagenesis and metasomatism (Nesbitt and Young, section 5.10). Within vegetation-free Precambrian palaeoenvironments, faster erosion rates and lower detritus residence periods before burial reduced the effects of enhanced weathering regimes (Condie et al., 2001 a; section 5.11 ). Iron-formation (IF) is an essentially Precambrian rock type; the oldest (c. 3.8 Ga) banded iron-formations were succeeded by poorly developed Archaean IFs, commonly associated with greenstone belt volcanism (Trendall and Blockley, section 5.4). The global peak in (banded) IF-time distribution at c. 2.5 Ga largely reflects deposits in the Hamersley (Pilbara craton, Australia) and Transvaal (Kaapvaal craton, South Africa) basins; abundant granular IF occurred at c. 1.8 Ga, with a long hiatus preceding local development of small IFs in the Neoproterozoic age (section 5.4). Archaean sea water is thought to have been enriched in (fumarolic) iron only below the pycnocline (Eriksson et al., 1997). Trendall and Blockley (section 5.4) favour an iron- and Eh-stratified ocean model, wherein deep, iron-rich water (Holland, 1973, 1984) welled up into shallow shelf settings of the Palaeoproterozoic Transvaal and Hamersley basins (Klein and Beukes, 1989) or where a pycnocline lay close to the level of these basin floors. This model obviates the use of IFs as a proxy for atmospheric oxygen content, as IF deposition occurred without direct atmospheric influence. The amounts of SiO2 in IF are enormous and its origin remains uncertain (see Klemm, 2000, for a review). The lack of major IF in the mid-Precambrian may be explained within the stratified ocean model by increasing oceanic oxidation as organisms and photosynthesis became more prevalent (section 5.4). The global carbon isotopic curve, based here largely on data from carbonate rocks in Australian Precambrian basins (~13Ccarb-time curve) is flat in the Neoarchaean (at c. 2.6 Ga) and into the early Palaeoproterozoic (Lindsay and Brasier, section 5.3). The two major 613Ccarb oscillatory excursions at c. 2.2-2.3 Ga and at c. 0.65 Ga are separated by essentially flat patterns (section 5.3). These 613Ccarb patterns may largely reflect plate tectonics and the supercontinent cycle (Lindsay and Brasier, 2002; Brasier et al., 2002) rather than atmospheric compositional changes. Major supercontinental assembly events at c. 2.8, 2.0 and 1.0 Ga were associated with mantle instability (sections 3.2-3.4); concomitantly, large intracratonic sag basins accumulated sediments over 200-500 My periods, including major carbonate platform successions (section 5.3). The large positive 613C excursion identified in carbonate rocks at c. 2.2-1.9 Ga is inferred to reflect significant burial of organic carbon (with resultant increase of the 02 production flux) to result in the "Great Oxidation Event" (GOE) of c. 2.3 Ga postulated by Karhu and Holland (1996) (see, however, discussion by Ohmoto, section 5.2). Although an early Archaean evolution of sulphate reducing bacteria is possible, an oxygen-deficient atmosphere at that time (following the "C-W-H-K model") would have resulted only in local examples of signature isotopic fractionations (Lyons et al., section 5.5). As oxygen levels possibly increased in the Palaeoproterozoic (according to the "C-W-H-K model" and in contrast to the "D-O model"), and in association with continental weathering, oceanic sulphate concentrations probably increased to levels where bacterial sulphate reduction was well expressed isotopically; this is supported by abun-

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dant 34S-enriched pyrites and relatively marked sulphur isotope variation in marine sulphates and sulphides (summarised by Lyons et al., section 5.5). There is a striking coincidence, in the Palaeoproterozoic, of the first major Ccarb isotopic excursion (Lindsay and Brasier, section 5.3), major BIF deposition (peak at c. 2.5 Ga; Trendall and Blockley, section 5.4) and evidence for the earliest global glaciation thereafter, at c. 2.4-2.2 Ga (Young, section 5.6). Earth's earliest global-scale glaciogenic deposits occur in passive margin (including epeiric seas) and partly in foreland basin tectonic settings (Young, section 5.6). They show no temporal association with either BIF or cap carbonate rocks, in contrast to their more widespread and stratigraphically complex Neoproterozoic equivalents. The "snowball Earth hypothesis" (SEH) (e.g., Kirschvink, 1992; Hoffman et al., 1998b), although often applied to both Proterozoic glacial intervals, is viewed here as an unlikely postulate, due to a wide variety of strong arguments outlined in sections 5.6 (Young), 5.7 (Williams), 5.8 (Frimmel) and 5.12. Although there is little doubt that Earth experienced major climatic perturbations in the Palaeo- and Neoproterozoic (see also section 5.3), explanation of these two poorly understood global glaciations is further complicated by their being preceded by development of supercontinents (e.g., sections 3.2 and 3.9). High continental freeboard (section 7.1) related to supercontinent assembly would have enhanced weathering regimes (sections 5.10 and 5.11) and thus also CO2 drawdown, and low palaeolatitudinal location would have increased albedo; together these factors may have lead to global cooling (Young, section 5.6). Although not providing a primary cause of global glaciation, the large obliquity hypothesis of Williams (1975, 1993; section 5.7) does offer a viable mechanism for the distribution and nature of the Precambrian glacial environments, and can explain many of the features observed in the Neoproterozoic glacial deposits. Frimmel (section 5.8) correlates the two glaciogenic units within the 770-540 Ma Gariep basin, Namibia, with the Sturtian (750-740 Ma) and Marinoan (590-580 Ma) (or possibly with the 560 Ma Moelv) glaciations, due to their host carbonate rocks having been deposited in either restricted basin conditions or under very shallow water. Anomalously low integrated sedimentation rates for the interval between the two identified glacial events suggest either a c. 100 My-long sea level lowstand (section 8.2) and cold palaeoclimate, or low subsidence rates due to enhanced mid-ocean ridge spreading rates. Frimmel (section 5.8) concludes that the sedimentary and C and Sr isotopic chemostratigraphic (cf., sea water proxies) record for the Gariep basin reflects an interaction of tectonic, eustatic and palaeoclimatic factors; positive 613C excursions are thus not always reliable stratigraphic markers. Episodic growth of the global reduced carbon reservoir supports stepwise oxygenation of the atmosphere resulting from episodic burial of carbon during large scale tectonic cycles (Des Marais, 1994a, 1997; Des Marais et al., 1992; section 5.3). In contrast, flat portions of the global 613Ccarb curve (> c. 2.2-2.3 Ga and in the mid-Proterozoic) reflect low tectonic activity and CO2 in the ocean-atmosphere system being in near-equilibrium with the mass balance of the carbon cycle (Brasier and Lindsay, 1998; Lindsay and Brasier, 2000; section 5.3). Sulphate availability in the early and mid-Proterozoic ocean probably remained low, relative to the Phanerozoic (Lyons et al., section 5.5). Oxygen levels likely

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rose significantly by the Neoproterozoic, thus raising oceanic sulphate and permitting enhanced sulphur isotope fractionation between sulphate and sulphide (to Phanerozoic-type values). This increasing fractionation was probably due to bacterial evolution and related isotopic effects within the oxidative part of the bio-geochemical sulphur cycle (section 5.5). Cyclic tidal rhythmites enable estimation of the Precambrian Earth's rotation and the orbit of the Moon (Williams, section 5.9). Data from c. 600 Ma and c. 750 Ma formations from South Australia indicate a length of day (LOD) of 21.9 and 21.4 hours per day, respectively (section 5.9). These palaeotidal and palaeorotational data indicate no significant expansion of Earth, at least from the late Neoproterozoic (Carey, 1976; Williams, 1998b, 2000). As tidal friction has gradually slowed Earth's rotation through geological time, the combination of lunar and solar torques has resulted in forced nutations (periodic tipping); the latter are subject to resonances (at various frequencies) with the free nutation of Earth's core (Toomre, 1974; Hinderer and Legros, 1988; section 5.9). LOD data suggest that annual resonances took place during the late Neoproterozoic-early Palaeozoic, and there were likely important resonances during the Archaean (Williams, section 5.9). Resonance, and thus also, instability, at the core-mantle boundary released heat through mantle (super)plumes (sections 3.2-3.4); there is thus a relationship between the temporal distribution of superplumes and celestial mechanics (see also chapter 1). It may thus be concluded, that the primary controls on Earth's (Precambrian) geological evolution are the interaction of plate tectonics with mantle superplumes and related thermal processes (chapters 2-4); the synergy of palaeoclimate, eustasy (chapter 8), ocean-atmosphere chemistry (chapter 5), bio-geology (chapter 6), and sedimentation (chapter 7) is directly dependent on these primary controls (e.g., Eriksson et al., 200 l a, b).

9.6.

EVOLUTION OF PRECAMBRIAN LIFE AND BIO-GEOLOGY

A gradual decrease in the flux of impactors affecting the solar system during the first 500-650 My of its history ended between c. 3850 and 3900 Ma (compare with sections 1.2 and 9.1). Between 0 and 6 planet-sterilising (and life-destroying, if life had already developed anywhere) impact events are estimated for the Hadaean and earliest Archaean, up until about 3850 Ma (Maher and Stevenson, 1988; Sleep et al., 1989). Although there is no record of life's origin, it arose on the early Earth under extreme conditions (Westall, section 6.6): hot, no oxygen or only trace amounts, possibly saltier oceans, higher UV flux (sections 5.2-5.5), and a relatively wide range of potential habitats for life existed, from deep water to subaerial (Nisbet, 1995; Nisbet and Sleep, 2001). The limitations of the Archaean rock record make it difficult to estimate the rate of colonisation of the various environments, but the oldest microfossil remains from the Barberton and Pilbara greenstone belts suggest widespread life in shallow water and the intertidal zone, and possibly even in subaerial settings (Buick and Dunlop, 1990; Walsh, 1992; Schopf, 1993; Hofmann et al., 1999; sections 6.2 and 6.6); a strong association between early microbial mats and hydrothermal activity is also noted (Nijman et al., 1999; Rasmussen, 2000; see also sections 2.7 and 6.6).

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The Precambrian record (c. 85% of the history of life) is dominated by prokaryotic (bacterial and cyanobacterial) microbes. Their minute size, incomplete preservation, and simple morphologies that can be mimicked by nonbiologic mineralic microstructures, make discrimination of true microbial fossils problematic (Schopf, section 6.2). As expected, these difficulties are more pertinent to the minuscule Archaean fossil record (number of finds = c. 30), in contrast to the abundant, well-preserved Proterozoic fossils (> 2800 authentic occurrences), many of which have biological affinities to modern taxa (Mendelson and Schopf, 1992; Schopf and Klein, 1992). By c. 3.5 Ga, "microbial life was flourishing and presumably widespread" (Schopf et al., 2002); this appearance of microfossils in the geological record is not preceded by any known bio-geochemical announcement (Altermann, section 6.3). Three major biologic indicators (fossils, organic matter, and isotopic signatures), especially if used together, provide strong evidence for early life. Stabilisation of siliciclastic sediments by microbial mats was apparently important in the Proterozoic, and rare examples are recognised from the Neoarchaean (Altermann, 2002). More than 600 stromatolitic units are known from the Proterozoic. The appearance of eukaryotic cells and, most significantly, of sexual reproduction in the Proterozoic, led to a rapid diversification of life that flourished subsequently in all hydrospheric environments (Altermann, section 6.3). Metazoan fossils appeared in the Mesoproterozoic, the oldest being preserved as trace fossils in 1.1 Ga sandstones in India (Seilacher et al., 1998). The oldest trace fossils interpreted as burrows produced by bilateralian animals occur in late Neoproterozoic shallow marine siliciclastic rocks of the Nama Group, Namibia (Jensen et al., 2000) and overlap in age with the Ediacara fauna. Eight known Early Archaean fossil-bearing deposits, c. 3.2-3.5 Ga and from the Pilbara craton and southeastern (Barberton area) Kaapvaal craton, contain in toto both stromatolites and spheroidal and filamentous microfossils (section 6.2), and other true fossils are reported from various localities (Schopf and Walter, 1983; Lanier, 1986; Klein et al., 1987; Altermann and Schopf, 1995; Kazmierczak and Altermann, 2002). The fossils include microscopic isolated single cells, paired (dividing) cells, ensheathed colonies of coccoidal cells, and both cylindrical and cylindrical-septate filaments, while macroscopic stromatolites include flat-lying, domical, columnar, and conical forms; inferred palaeoenvironments range from shallow marine to hydrothermal (Schopf, section 6.2). Two types of fossil are thus present in the early record, stromatolites and cellularly preserved microorganisms. Schopf (section 6.2) defines a stromatolite as: "an accretionary organosedimentary structure, commonly thinly layered, megascopic, and calcareous, produced by the activities of mat-building communities of mucilage-secreting microorganisms, filamentous and coccoid photoautotrophic prokaryotes such as cyanobacteria". Altermann (section 6.5) discusses in detail, problems in the definition, classification, and palaeoenvironmental and stratigraphic applications of stromatolites. The oldest known stromatolites, from the 3.5 Ga Warrawoona Group of the Pilbara craton, are small and of limited lateral extent. Large carbonate platformal buildups began with the c. 3.0 Ga Pongola Supergroup, Kaapvaal craton; these microbial communities were the first to form large bioherms and to influence local architecture within sedimentary basins (Walter, 1983; Altermann, section 6.3). Stable tectonic terrains related to early proto-

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continental growth may have been even more important in the evolution of these carbonate platforms than their biologic heritage (Grotzinger, 1989, 1994). Soon after, at 2.6 Ga, giant carbonate platforms developed in intracratonic basins on the Kaapvaal and Pilbara cratons, where biogenic activity governed sedimentation (Nelson et al., 1999; Altermann, section 6.3) and where accumulation rates (section 7.11) matched those of modern equivalents (Lanier, 1986; Altermann and Nelson, 1998). Microbial biostromes and bioherms governed the internal basin architecture, water depth and facies distribution in these large epeiric (section 7.7) marine depositories (Altermann, section 6.3). The origin of Precambrian sedimentary carbonate rocks is poorly understood due to diagenetic changes and tectonism, and the role of micro-organisms in the precipitation of Precambrian carbonate minerals is thus difficult to assess (e.g., Fairchild, 1991; Riding, 2000). Many workers support non-biological deposition of Archaean carbonates directly on the sea floor, or accumulation from massive whitings (Grotzinger, 1994; Grotzinger and Knoll, 1995), whereas others see early Archaean stromatolites as microbially generated bio-sedimentary structures, or as resulting from purely inorganic processes (Lowe, 1994b; Grotzinger and Rothman, 1996). Some workers suggest that micro-organisms only played an important role in carbonate sedimentation from the beginning of the Mesoproterozoic, and particularly in the Neoproterozoic (Grotzinger, 1994; Bartley et al., 2000). Structural, textural and mineralogical similarity of Precambrian and modern carbonate rocks implies that they are products of very similar microbiota and calcification processes (Kazmierczak et al., section 6.4). Although direct inorganic precipitation of calcite from Ca-saturated sea water on the Archaean sea floor may have occurred (Sumner and Grotzinger, 1996a), directly precipitated sea floor Ca-cements become rare from the Mesoproterozoic, in parallel with the decrease in stromatolite diversity and abundance (Altermann, section 6.3). The switch from a Na-(carbonate)-rich ocean to a NaCI (halite) ocean at c. 1.8 Ga (when calcitic seafloor precipitates gave way to sulphate evaporites) may have led to the subsequent decline in stromatolites (Grotzinger, 1994). The lack of sulphate evaporites in > 1.8 Ga rocks may reflect low levels of oceanic sulphate due to low oxygen concentration in the atmosphere (Grotzinger and Kasting, 1993; see also sections 5.5, 9.5). In contrast, Kempe and Degens (1985) argue for bicarbonate- and soda-dominated Precambrian oceans (with low chloride concentrations) until c. 1.0 Ga; slow accumulation of NaC1 from hydrothermal leaching of the ocean floor and gradual removal of Na-carbonates to the crust possibly brought on the demise of the "soda ocean" (Altermann, section 6.3). Identification of organic biomarkers, particularly hydrocarbons, is useful in the Proterozoic record as, for example, in the case of the protozoan biomarker tetrahymenol in c. 930 Ma sediments (Summons, 1992). Detection of fossil testate amoebae in the same sequence (Schopf, 1992c; Porter and Knoll, 2000) indicates a minimum age for the Proterozoic emergence of protozoan protists. However, richly petroliferous deposits are unknown in the Archaean, and detection of small quantities of organic biomarkers is thus questionable due to the likelihood of younger contaminants (section 6.2). The many measurements of the isotopic composition of kerogenous components of diverse fossil-bearing shales and cherts (Strauss and Moore, 1.992) include examples as

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old as c. 3.5 Ga. However, such analyses only provide strong evidence for the existence of photoautotrophic primary producers, and cannot enable differentiation between the contributions of oxygenic and anoxygenic photosynthesisers or link the isotopic compositions to particular species (Schopf, section 6.2). Biogenic carbon isotopic ratios from inferred metasediments from the c. 3.8 Ga Isua greenstone belt, Greenland (Schidlowski, 1988; Mojzsis, 1996; Rosing, 1999) are not considered reliable due to misinterpretations of their host rocks (see details in Myers, section 2.2; Altermann, section 6.3). The sudden appearance of eukaryotes in the Proterozoic geological record after almost 1500 My of exclusively prokaryote life matches the sudden emergence of the latter at c. 3.5 Ga (Altermann, section 6.3). Spiral-shaped, megascopic fossils (Grypania) from 2.1 Ga BIF in Michigan were classified by Han and Runnegar (1992) as probable eukaryotic algae based mainly on morphometric arguments (section 6.3). Large coccoid microfossils, usually regarded as eukaryotes, only became abundant in the Mesoproterozoic; the eukaryotic organisms probably arose from prokaryotes, which lived in symbiosis with other eubacteria (Margulis, 1981 ). If Han and Runnegar (1992) are correct, division of the phylogenetic tree separating prokaryotic eubacteria from archaea and eukaryots must have occurred long before 2.1 Ga. The strictly aerobic eukaryotes' emergence would have necessitated oxygen levels of 1-2% PAL (Chapman and Schopf, 1983), which lends some support to models of early atmospheric oxygen (e.g., Ohmoto, 1999; sections 5.2 and 9.5). Acritarchs (eukaryotic algae of unknown biological affinity) appeared at 1.75 Ga, and are the most widespread fossils in Meso- and Neoproterozoic rocks, reaching their maximum diversity at c. 600 Ma, after the Varanger ice age (section 6.3). At c. 1.1 Ga, a rapid diversification of eukaryotic phytoplankton occurred, reaching a maximum at about 900 Ma; a major decline in stromatolite diversity and abundance at 800-700 Ma was related to a decrease in atmospheric CO2 and an increase of 02 (Holland, 1984) with concomitant glaciation (sections 5.6-5.8). The strong diversification of eukaryotic life in the Neoproterozoic enabled, for the first time, biostratigraphic resolution that permits inter-basinal correlations (e.g., section 5.8).

9.7.

SEDIMENTATION REGIMES THROUGH TIME

From c. 4.0-3.2 Ga, a combination of intra-oceanic island arcs, oceanic plateaus and plate tectonic collisional processes is thought by many to have led to the development of proto-continents (Windley, 1995) (see, however, section 3.6, for an alternative model) that may have constituted only 5-10% of present crustal volumes (e.g., Eriksson, 1995). Generally, komatiitic, tholeiitic and felsic volcanic and volcaniclastic rocks predominate in c. 4-3.2 Ga greenstone belts (Fedo et al., 2001) (see also sections 2.2-2.4, 4.4 and 9.2-9.4). Associated with these dominant lithologies, thin remnants of passive margin carbonates, BIF, stromatolitic evaporites, pelites and quartzites, as well as subordinate synorogenic turbidites, conglomerates and sandstones are found; they reflect increasingly stable small continental nuclei (Windley, 1995).

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The scarcity of early Precambrian fossils (e.g., sections 6.2 and 9.6) makes detailed study of sedimentary structures in the ancient rock record critical (e.g., sections 7.2, 7.4 and 7.5). Microbial micro-organisms were important locally in trapping, binding and precipitating sediments in situ, to build small carbonate platforms during the Archaean (Wright and Altermann, 2000) (section 6.4), and microbial mats probably trapped clastic sediments as well (sections 7.10 and 7.11). Archaean gypsum deposits may either have been evaporites derived from sea water compositions similar to modern equivalents (Lowe, 1983; Buick and Dunlop, 1990) or formed at sites of continental runoff (Grotzinger and Kasting, 1993). Archaean seawater (section 5.2) was probably enriched in iron of fumarolic origin beneath the pycnocline (Veizer, 1983a), possibly leading to sulphidic iron-formation deposition at these depths, and oxidic deposits in areas of photosynthetic productivity above the pycnocline (Eriksson et al., 1997) (section 5.4 provides a detailed examination of iron-formation genesis). Eriksson (1983) suggests that Archaean iron-formations were analogous to modern starved basin pelagic ("rain-out") sediments. Depositional regimes interpreted from Mesoarchaean greenstone belts include debrisflows on high gradient alluvial fans, low sinuosity rivers, shallow marine settings with wave and tidal action, and turbidity currents, some associated with hummocky cross-strata and thus suggesting deposition near storm wave base (Eriksson et al., 1997; section 7.3). In general, shallow marine conditions appear to have been prominent within greenstone palaeoenvironments (e.g., Windley, 1995; Eriksson et al., 1997). Apart from inferred (and often contentious) ophiolites (section 3.7) within highly deformed greenstone stratigraphies (section 7.4), no unequivocal ocean floor of Precambrian age has been preserved (Fedo et al., 2001). Significant crustal growth inferred for the 3.2-2.6 Ga period due to amalgamation of oceanic terranes at c. 3.3-3.2 Ga (section 3.6 provides a different model) was followed by the development of passive continental margins on the earliest stabilised craton, the Kaapvaal. In addition, Cordilleran- and Himalayan-style collisions (e.g., section 3.8) of growing cratons occurred, and on many, continental flood basalts (Windley, 1995) of global superplume affinity (sections 3.2 and 3.3) formed during a c. 2.7 Ga event (e.g., Eriksson et al., 2002b). Catastrophic supply of volcaniclastic debris often choked siliciclastic greenstone sedimentation systems (Mueller and Corcoran, 1998; section 7.3). Increasing cratonic stability led to rift basins and strike-slip basins becoming common (e.g., Mueller and Corcoran, 1998; Smithies et al., 2001; section 7.3). Neoarchaean greenstone belts (sections 2.4 and 7.4) commonly exhibit syntectonic styles of clastic sedimentation, with alluvial fans and immature braided rivers (section 7.8) passing directly into high energy shallow marine settings, with aggressive weathering (reflecting palaeoatmosphere composition; sections 5.10 and 5.11) forming mature sandstone close to source areas (Corcoran et al., 1998). In contrast, Donaldson and de Kemp (1998) argue that mature Archaean sandstones reflect periods of crustal stability. The c. 3.0-2.94 Ga Mallina basin, Pilbara craton, provides evidence of sediment recycling, but within an active greenstone-type depository (Smithies et al., 2001). In contrast, the more stable Kaapvaal was characterised by the famous auriferous c. 3-2.7 Ga Witwatersrand basin. Catuneanu (2001) applies a retroarc foreland basin model to these

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thick siliciclastic fluvial and littoral sediments (section 7.5); limited ventifacts (section 7.6) point to localised aeolian erosion processes (Els, 1998). Contemporaneously, on other terranes (Zimbabwe craton, Slave Province), continental growth continued through greenstone belt evolution (Hofmann et al., 2001; Mueller and Corcoran, 2001). As inferred continental growth rates peaked close to the Archaean-Proterozoic boundary (Eriksson, 1995; Arndt, section 2.8), increasingly large cratons were characterised by widespread (global?) orogenic quiescence from c. 2.6-2.4 Ga (Windley, 1995; Mints and Konilov, section 3.9). The development of large epeiric (section 7.7) basins and chemical and clastic passive margin platforms during this quiescent period (Windley, 1995) may reflect the first, Neoarchaean supercontinent ("Kenorland"; e.g., Aspler and Chiarenzelli, 1998). Condie et al. (2001)relate this to a c. 2.7 Ga global mantle superplume event (section 3.2). The Archaean-Proterozoic boundary is diachronous, with older cratons such as Pilbara and Kaapvaal stabilising earlier (Windley, 1995). Diagnostic geochemical changes across this boundary are related to a change from mantle buffering of sea water (hydrothermal interaction of seawater and juvenile crust at mid-ocean ridges) to continental buffering (as river discharge from stable continents became predominant) (Veizer, 1983a, b, 1988). Analogous geochemical changes across this boundary are seen in continental pelites (reflecting more evolved felsic-rich source rocks; Wronkiewicz and Condie, 1990) and in Palaeoproterozoic greenstone basalts (which reflect greater depths of magma generation; Condie, 1989). Earth's earliest large scale carbonate-BIF platformal sequences (Hamersley Group, Pilbara; Chuniespoort-Ghaap Groups, Kaapvaal) developed in this 2.6-2.4 Ga time interval (section 5.4) due to global eustatic rise between c. 2630 and 2430 Ma (Nelson et al., 1999; Eriksson et al., 2001 b). Their analogous lithostratigraphy has led to suggestions of a "Vaalbara" supercontinent (e.g., Cheney, 1996); however, geochronologic and palaeomagnetic data do not support this (Altermann and Nelson, 1998; Wingate, 1998; Nelson et al., 1999; Eriksson et al., 2002b). During the same time interval (2.6-2.4 Ga), the thick BIF-carbonate succession of the Minas Supergroup (Sao Francisco craton, southeastern Brazil) was deposited (Alkmim and Marshak, 1998). Windley (1995) suggests supercontinental fragmentation as a predominant influence on sedimentation from c. 2.4-2.2 Ga, with concomitant rift and passive margin deposition being important. Support for this idea is garnered from interpretation of a number of partial to full Wilson cycles for parts of the postulated Kenorland supercontinent (Aspler and Chiarenzelli, 1998; Young et al., 2001; Ojakangas et al., 2001a, b; Aspler et al., 2001) (see also section 3.9). While fragmentation of Kenorland proceeded, a "southern" supercontinent may have begun assembly (Eriksson et al., 1999b). Young et al. (2001) infer that high weathering rates at low palaeolatitudes within Kenorland encouraged drawdown of atmospheric CO2, thus encouraging global cooling; reduced weathering during global icehouse conditions and volcanic CO2 are thought to have resulted in return to a greenhouse state (sections 5.2, 5.6 and 5.7). Breakup of Kenorland may have interfered with this self-regulatory global icehouse-greenhouse cyclicity (Young et al., 2001). Williams (sections 5.7 and 5.9) discusses celestial mechanics within the solar system as a primary influence on Earth's first global refrigeration event. Significant quantities of free oxygen may

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have become available in the atmosphere during the c. 2.4-1.9 Ga period (e.g., Windley, 1995; see, however, mutually exclusive atmospheric models in sections 5.2 and 9.5). There is evidence of major c. 2.0-1.7 Ga crttstal growth from the southwestern U.S.A., Western Greenland, the Baltic shield, the B irimian belt in West Africa, and from Brazil (Nelson and DePaolo, 1985; Patchett and Arndt, 1986; Mil6si et al., 1992; Schrank and Silva, 1993; Windley, 1995). The Laurentia supercontinent, thought to have amalgamated at c. 2.0-1.7 Ga (Hoffman, 1988; Aspler et al., 2001), coincides with a global mantle superplume event (sections 3.2 and 3.3) at c. 1.9 Ga (Condie et al., 2001). Aeolian ergs first formed from c. 1.8 Ga, after large land masses became common (Simpson et al., section 7.6). From c. 1.6 Ga, the supercontinental cycle and the various sedimentary regimes associated with the Wilson cycle, became well developed (e.g., Hoffman, 1989c, 1991; Barley and Groves, 1992; Windley, 1995). A global event, characterised by a combination of increased atmospheric oxygen, supercontinent assembly and breakup, and global glaciations occurred in the Neoproterozoic, closely resembling the earlier, c. 2.4-2.2 Ga event (e.g., Fedo and Cooper, 2001; Dehler et al., 2001; Martins-Neto et al., 2001 ; sections 3.10, 3.11,5.7 and 5.8).

9.8.

SEQUENCE STRATIGRAPHY THROUGH TIME

Sequence stratigraphic models assume that a predictable stacking pattern of sedimentary facies or systems tracts (i.e., genetic facies associations grouped within various sedimentary geometries or architectures) is controlled essentially by the interaction of base level changes and sedimentation at the shoreline (Catuneanu et al., section 8.2). These base level changes depend upon a wide range of geological variables, discussed in previous chapters of this book. Generation of continental crust (chapters 2 and 4), crustal growth rates (section 2.8), and the interplay of tectonism and mantle plumes (chapter 3) provide first-order controls on base level. Over shorter time intervals, additional controls on stratigraphic cyclicity involve the interplay of a multitude of factors, including palaeoclimatic (chapter 5), biological (chapter 6) and depositional influences (chapter 7). Sequence stratigraphy thus draws together the principles from the many diverse fields of Precambrian geological investigation discussed in this book. The great strength of sequence stratigraphy is that it relates genetic processes directly to patterns that can be observed in the rock record, hence enabling reconstructions of basins' evolution through time. Application of sequence stratigraphic analysis becomes more difficult with increasing stratigraphic age, reflecting poor preservation, post-depositional tectonics, diagenetic and metamorphic changes, and a lack of practical biostratigraphy. However, if the geometry, sedimentary facies and facies relationships of a succession allow reliable depositional models to be constructed, the technique can be applied, even with minimal chronological constraints. Examples of application of sequence stratigraphy to Precambrian successions are few, as yet (Christie-Blick et al., 1988; Beukes and Cairncross, 1991; Krapez, 1996, 1997; Catuneanu and Eriksson, 1999, 2002; Catuneanu and Biddulph, 2001).

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The application of sequence stratigraphy to fluvial depositional sequences (lacking a coeval shoreline) is more difficult, and recent work examines changes in fluvial styles and architectural elements (see sections 7.8 and 8.5) to identify discernable packages with specific characteristics within a fluvial sequence. The application of low and high accommodation systems tracts to fluvial deposits (e.g., Dahle et al., 1997; Boyd et al., 1999; Zaitlin et al., 2000) is particularly relevant to the Precambrian, where basins are often incompletely preserved; the first application of this new concept to the Precambrian record is provided in sections 8.3 and 8.4. A key aspect of sequence stratigraphic analysis is the separation of sequences and bounding surfaces into groups of different relative importance, which defines the concept of hierarchy. The fundamental criterion that should be employed in order to design a hierarchy system is still subject to debate, and the choice is generally between time units (duration of cycles) versus the magnitude of base level changes that result in boundary formation. The reason this debate still continues today is because sequence stratigraphy draws its principles primarily from the Phanerozoic record, which only captures a relatively small fraction of geologic time. Hence, our "window to the world" is rather small when we only look at a glimpse (c. 12%) of Earth history, which does not offer a representative sample to extract the essence of the ground truth. The lesson learnt from Precambrian case studies is that change, rather than constancy, is the norm for geological processes, as documented by many aspects of Earth's evolution. This means that the duration of cycles becomes less relevant relative to the mechanisms of which they are the product, and therefore the physical characteristics of strata and their bounding surfaces are more important than the frequency of their occurrence or change in the rock record.

9.9.

TEMPOS AND EVENTS IN PRECAMBRIAN TIME

The previous eight sections have summarised the inferred rates (cf. "tempos") at which significant geological processes occurred, as well as major "events" (both defined in the Preface) postulated for Precambrian time. These are summarised in a series of figures (Figs. 9.9-1-9.9-4), in an attempt to examine correlations of these events from the highly diverse fields of research covered in the eight main chapters of this book. Immediately apparent from the formation of the solar system, is the rapidity with which proto-Earth became differentiated into metallic core and silicate mantle within about 20 My of the core collapse supernova event (Fig. 9.9-1 ). A major event at c. 4550 Ma, immediately thereafter, was the inferred collision of "Theia" with the proto-Earth to form the Moon; this resulted in shock melting and a terrestrial magma ocean, which cooled over a period of c. 10-100s My, with transient (probably komatiitic) crusts being rapidly recycled into the mantle, as well as being brecciated by common meteorite impacts. The meteorite impact rate, estimated at c. 15 times that at present for 3.8 Ga had decreased to c. 2 times by 3.0 Ga, and several impact spherule layers are known from 3420 Ma to 2490 Ma (Fig. 9.9-1). Although subject to much debate, continental crustal growth rates fluctuated throughout geological time but may have peaked at c. 2.7 Ga. In the Neoarchaean, there is strong evi-

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Fig. 9.9-1. Summary time-chart illustrating the formation of the solar system and the evolution of the early Earth. Based on chapter 1.

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Fig. 9.9-2.

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dence from many parts of the world for active plate tectonics of recognisably Phanerozoicmodern style. For the period prior to c. 2.7 Ga, viewpoints on the mechanisms of formation of cratons and proto-cratons remain divergent: does the generation of continental crust reflect diapirism, or far-field plate tectonic influences including low-angle subduction, or were pre-3.3 Ga basins formed due to extensional collapse related to delamination and mantle plume activity (Fig. 9.9-2)? Prior to c. 2.7 Ga, heat flow values are thought to have exceeded those at present by 2-6 times, and Trendall (2002) has proposed a "plughole" model, which envisages a gradual transition from a thermally dominated Hadaean Earth to one where plate tectonics became predominant (section 3.6) by c. 2.7-2.5 Ga. Figures 9.9-2-9.9-4 document the nature of the changes that Earth underwent at about 2.7 Ga, when the first inferred supercontinent ("Kenorland") formed, preceded by an increasing frequency (from 2.8 Ga) of large igneous provinces (LIPs; cf. mantle plumes; section 3.3) and succeeded by a c. 2.7-2.5 Ga possible ophiolite complex cluster (section 3.7). There is also a postulated, possibly global komatiite eruption event at 2705 Ma (Fig. 9.9-2). This 2.7 Ga "superevent" (event cluster) also appears to encompass the possibility of global catastrophic mantle overturn (section 3.4), and the first catastrophic slab avalanche and related global superplume event (section 3.2; Fig. 9.9-2). A second "superevent" appears pertinent at about 2.2-1.8 Ga, and in the period between this and the 2.7 Ga event-cluster (Figs. 9.9-2-9.9-4), there was a significant measure of tectonic quiescence on the global scale. This quieter period was marked by global eustatic rise (consequent upon the inferred maximum crustal growth at c. 2.7 Ga) and concomitant large epeiric, passive margin basins on many of the cratons then extant; these include the first giant carbonate platforms, on Pilbara and Kaapvaal. Banded iron-formations were a significant part of these depositories and a global BIF peak is noted at c. 2.5 Ga (Fig. 9.9-3). Following this, towards the end of this postulated quiescent period, the first global glaciation (c. 2.4-2.2 Ga) occurred. The second postulated "superevent", at c. 2.2-1.8 Ga, is again one characterised by first-order magmatic and tectonic events which accommodated second-order palaeoatmospheric and oceanic-chemical change, as well as a number of transformations within the biological and sedimentary systems on Earth (Figs. 9.9-2-9.9-4). Formation of two supercontinents (a "southern" supercontinent lasting from c. 2.2-1.8 Ga, and the "northern" Laurentia, from c. 2.0-1.7 Ga) was accompanied, again (as at c. 2.7 Ga), by a peak in greenstone-style volcanism (c. 2.2-2.1 Ga), an apparent ophiolite complex cluster (c. 2.3-1.8 Ga), and by the second postulated global superplume event (c. 1.9 Ga). Bearing in mind the two main (and mutually exclusive) hypotheses on palaeo-atmospheric evolu-

Fig. 9.9-2. Summary time-chart giving a schematic summary of the generation of continental crust on the Precambrian Earth. Note different hypotheses of granite-greenstone crustal evolution (denoted as Ia and b, and II" detailed text discussions are provided in chapter 2), the c. 2.7 Ga Kenorland, c. 2.2-1.8 Ga "southern", c. 2.0-1.7 Ga Laurentia and c. 1.2 Ga Rodinia supercontinents, as well as global superplume events at c. 2.7 Ga and 1.9 Ga. TTG = tonalite, trondhjemite-granodiorite; LIP = large igneous provinces (cf. section 3.3). Based on chapters 2-4.

766

Fig. 9.9-3.

Chapter 9: Towards a Synthesis

9. 9. Tempos and Events in Precambrian Time

767

tion ("D-O" and " C - W - H - K " models--see section 5.2; Fig. 9.9-3), the postulated "great oxidation event" at c. 2.3-2.0 Ga is matched by significant chemical changes in the oceans (with the first stepwise increase in SO 2- at c. 2.2 Ga, and a first major oscillatory excursion in the global C isotope curve at c. 2.3-2.2 Ga). Associated with these is the apparently "sudden" emergence of eukaryotic life at c. 2.1-2.0 Ga (Figs. 9.9-3 and 9.9-4). At about 1.8 Ga, towards the end of this postulated "superevent", granular iron-formations rather than BIF became abundant, the first ergs (aeolian sand seas) developed on a number of cratons, and the previously high Fe 2+ and low H2S oceans gave way to those with low iron and high hydrogen sulphide. It is possible that there was a switch from a soda-ocean to a saline-ocean at c. 1.8 Ga. The first acritarchs appear at c. 1.75 Ga (Fig. 9.9-4). A second period of relative quiescence in terms of global-scale significant change on Earth followed this second postulated "superevent" at c. 2.2-1.8 Ga, and was brought to an end by a third possible event cluster, in the Neoproterozoic (c. 0.8-0.6 Ga; Figs. 9.9-2-9.9-4). Much of the Proterozoic rock record is characterised by linear orogenic belts of both low and high grade granitic rocks, formed during collision of the continents and the reworking of their margins. At c. 0.8 Ga there was a second stepwise increase in oceanic S O ] - , and stromatolites show a major decline at c. 0.8-0.7 Ga, during which BIF enjoyed a short-lived return to marine sedimentary basins. Glaciation accompanied these changes, with the 0.75-0.74 Sturtian, 0.59-0.58 Marinoan and 0.56 Ga Moelv events. Another major global oscillatory excursion occurred in the global C isotopic curve (Fig. 9.9-3) and the acritarchs achieved their maximum development at about 0.66 Ga (Fig. 9.9-4). It is noticeable that the first global glaciation at c. 2.4-2.2 Ga occurred towards the close of the first postulated quiescent period (c. 2.6-2.2 Ga) whereas the Neoproterozoic glaciations occurred within and shortly after the third suggested "superevent" at c. 0.8-0.6 Ga. This may support the idea that CO2 drawdown due to weathering of high-freeboard, exposed larger continental land masses was an important cause of the earlier global glaciation (Young, section 5.6), whereas celestial mechanics played an important part in the younger global glaciation (Williams, sections 5.7 and 5.9). In closing we thus infer that Precambrian time can be divided up into a number of periods, either comprising highly diverse events (varying from global magmatism and plate tectonics to atmospheric, biologic and sedimentary system changes) which appear to cluster at certain periods (c. 2.7 Ga, c. 2.2-1.8 Ga, and c. 0.8-0.6 Ga), or relatively longer intervening periods when there was an apparent relative quiescence on Earth's surface on a global scale. Of course, these latter periods also experienced many changes such as common plate movements and even smaller supercontinents (e.g., the postulated "Columbia"

Fig. 9.9-3. Summary time-chart of atmospheric, oceanic and climatic evolution of the Precambrian Earth. Note two mutually exclusive models for palaeo-atmospheric and oceanic evolution: "DO" = Dimroth-Ohmoto model, and "CWHK" = Cloud-Walker-Holland-Kasting model (see section 5.2 for discussion). Major changes in oceanic and atmospheric chemistry are shown, as well as the schematic history of iron-formation (IF; BIF = banded IF) evolution over Precambrian time, and global glaciation events. PAL -- present atmospheric level. Based on chapter 5.

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EVOLUTION OF LIFE AND SEDIMENTATION SYSTEMS

Fig. 9.9-4. Summary time-chart ofbio-geological evolution and the change in sedimentation systems over Precambrian time. Based on chapters 6-8.

9. 9. Tempos and Events in Precambrian Time

769

supercontinent in the Mesoproterozoic; Rogers and Santosh, 2002), as well as many mantle plumes (sections 3.2 and 3.3) and changes within all natural geological processes. Within this postulated chronological framework the Phanerozoic could also be considered as a period of stasis (cf. Lindsay and Brasier, section 5.3), which once again accommodated several global glaciation events. Most of the Archaean and the preceding Hadaean (prior to the c. 2.7 Ga "superevent") were probably highly unstable periods of Earth's history, during which time the processes responsible for continental crustal growth, plate tectonics, oceanic crust, sedimentation systems, later atmospheric evolution and life itself were becoming established on an Earth increasingly spared from meteorite impacts and catastrophic global mantle overturns. This book also emphasises the interdependence of all geological processes on Earth in establishing what is preserved in the Precambrian rock record. As an example of this, significant global eustatic changes (cf. sequence stratigraphy, chapter 8) would be "events" in the terminology adopted within this book. However, their occurrence and magnitude would be a reflection of the thermal and isostatic state of the continents (i.e., freeboard, section 7.1), the possible impingement of mantle plumes beneath continental and oceanic crust, variability of constructive and destructive plate margin processes, continental weathering and possible glaciation, celestial mechanics of the solar system, and subsidence versus sedimentation (cf. accommodation space) along continental margins. No definition of global eustatic "events" would thus be possible without understanding equally the many other related "events" and "tempos" operating on Earth during the Precambrian.

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REFERENCES

Abbott, D.H., 1996. Plumes and hotspots as sources of greenstone belts. Lithos, 37:113-127. Abbott, D.H. and Hoffmann, S.E., 1984. Archean plate tectonics revisited: 1. Heat flow, spreading rate, and the age of subducting oceanic lithosphere and their effects on the origin and evolution of continents. Tectonics, 3: 429-448. Abbott, D.H. and Isley, A., 2002a. Extraterrestrial influences on mantle plume activity. Earth Planet. Sci. Lett., 205: 53-62. Abbott, D. and Isley, A., 2002b. The intensity, occurrence, and duration of superplume events and eras over geological time. J. Geodynamics, 34: 265-307. Abbott, D., Burgess, L., Longhi, J. and Smith, W.H.F., 1994. An empirical thermal history of the Earth's upper mantle. J. Geophys. Res., 99: 13,835-13,850. Abbott, D.H., Sparks, D., Herzberg, C., Mooney, W., Nikishin, A. and Zhang, Y., 2000. Quantifying Precambrian crustal extraction: The root is the answer. Tectonophysics, 322: 163-190. Abdelsalam, M.G. and Stern, R.J., 1996. Sutures and shear zones in the Arabian-Nubian Shield. J. Afr. Earth Sci., 23: 289-310. Abe, Y., 1993. Physical state of the very early Earth. Lithos, 30: 223-235. Abe, Y., 1997. Thermal and chemical evolution of the terrestrial magma ocean. Phys. Earth. Planet. Interior, 100: 27-39. Abouchami, W., Boher, M., Michard, A. and Albarede, F., 1990. A major 2.1 Ga event of mafic magmatism in West Africa: an early stage of crustal accretion. J. Geophys. Res., 95: 17,605-17,629. Adams, E.W., Grotzinger, J.P., Schr6der, S., A1-Siyabi, H., Amthor, J.E. and McCormick, D.S., 2002. Shape, size and distribution of microbial reefs in a carbonate ramp setting (Nama Group, Namibia). Abstract, 16th International Sedimentological Congress, July 2002, Rand Afrikaans University, Johannesburg, p. 6. Ager, D.V., 1981. The Nature of the Stratigraphical Record, 2nd edn. Halsted Press, New York, 122 pp. Agnor, C.B., Canup, R.M. and Levison, H.F., 1999. On the character and consequences of large impacts in the late stage of terrestrial planet formation. Icarus, 142:219-237. Aigner, T., 1985. Storm Depositional Systems. Lecture Notes in Earth Sciences, 3. Springer-Verlag, Berlin, 174 pp. Aitken, J.D., 1967. Classification and environmental significance of cryptalgal limestones and dolomites, with illustrations from the Cambrian and Ordovician of Southern Alberta. J. Sediment. Petrol., 37(4): 1163-1178. Alam, N.M., Crook, K.A.W. and Taylor, G., 1985. Fluvial herringbone cross-stratification in a modem tributary mouth bar, Coonamble, New South Wales, Australia. Sedimentology, 32: 235-244. Alapiety, T.T., Filen, B.A., Lahtinen, J.J., Lavrov, M.M., Smolkin, V.F. and Voitsekhovsky, S.N., 91990. Early Proterozoic layered intrusions in the northeastern part of the Fennoscandian Shield. Mineral. Petrol., 42: 1-22. Albar6de, E, 1998. The growth of continental crust. Tectonophysics, 296: 1-14. Albarbde, E, Blichert-Toft, J., Vervoort, J.D., Gleason, J.D. and Rosing, M., 2000. Hf-Nd isotope evidence for a transient dynamic regime in the early terrestrial mantle. Nature, 404: 488-490.

772

References

Alexander, C.M.O'D., Boss, A.P. and Carlson, R.W., 2001. The early evolution of the inner solar system: a meteoritic perspective. Sci., 293: 64-68. Alexeyev, N.L., Salnikova, E.V. and Klepinin, S.V., 1999. Tectonic and P-T-time evolution of the Kolvitsa-Umba collision zone. Abstract, SVEKALAPKO Workshop, November 1997, Lammi, Finland, p. 8. Alibert, C. and McCulloch, M.T., 1993. Rare earth element and neodymium isotopic compositions of the banded iron-formations and associated shales from Hamersley, Western Australia. Geochim. Cosmochim. Acta, 57: 187-204. Alibert, C., Kinsley, L. and McCulloch, M.T., 1991. Rare-earth elements and Nd isotope systematics in the banded iron-formations from the Hamersley Basin (Western Australia) and implications for the composition of Early Proterozoic seawater. In: Annual Report, 1990. Australian National University, Research School of Earth Sciences, pp. 16-17. Alkmim, EE and Marshak, S., 1998. Transamazonian orogeny in Southern S~o Francisco Craton Region, Minas Gerais Brazil: evidence for Paleoproterozoic collision and collapse in the Quadrihitero Ferrffero. Precambrian Res., 90: 29-58. Allaart, J.H., 1982. Geologic map of Greenland, 1" 500000. Frederikshfib Isblink--SCndre Stromfjord. GrCnlands Geologiske UndersCgelse, Copenhagen. All~gre, C.J., Manh6s, G. and G6pel, C., 1995. The age of the Earth. Geochim. Cosmochim. Acta, 59: 1445-1456. Allen, J.R.L., 1984. Sedimentary Structures: Their Classification and Physical Basis. Developments in Sedimentology, 30. Elsevier, Amsterdam, 663 pp. Allen, J.R.L., 1985. Principles of Physical Sedimentology. Allen and Unwin, Winchester, MA, 272 pp. Allen, EA., 1981. Wave-generated structures in the Devonian lacustrine sediments of south-east Shetland and ancient wave conditions. Sedimentology, 28: 369-379. Allen, P.A., Bowring, S., Leather, J., Brasier, M., Cozzi, A., Grotzinger, J.E, McCarron, G. and Amthor, J.J., 2002. Chronology of Neoproterozoic glaciations: new insights from Oman. Abstract, 16th International Sedimentological Congress, 8-12 July 2002, Rand Afrikaans Univ., Johannesburg, pp. 7-8. Allen, R.L., Lundstr6m, I., Ripa, M., Simeonov, A. and Christofferson, H., 1996a. Facies analysis of a 1.9 Ga, continental margin, back-arc, felsic caldera province with diverse Zn-Pb-Ag-(Cu-Au) sulfide and Fe oxide deposits, Bergslagen region, Sweden. Econ. Geol., 91: 979-1008. Allen, R.L., Weihed, P. and Svenson, S.-A., 1996b. Setting of Zn-Cu-Au-Ag massive sulfide deposits in the evolution and facies architecture of a 1.9 Ga marine volcanic arc, Skellefte district, Sweden. Econ. Geol., 91: 1022-1053. Alsop, G.I., Brown, J.P., Davison, I. and Gibling, M.R., 2000. The geometry of drag zones adjacent to salt diapirs. J. Geol. Soc. London, 157: 1019-1029. Altermann, W., 2001. The oldest fossils of Africa--a brief reappraisal of reports from the Archean. J. Afr. Earth Sci., 33: 427-436. Altermann, W., 2002. The evolution of life and its impact on sedimentation. In: W. Altermann and EL. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. Spec. Publ. IAS, 33. Blackwell, Oxford, pp. 15-32. Altermann, W. and H~ilbich, I.W., 1990. Thrusting, folding and stratigraphy of the Ghaap Group along the south-western margin of the Kaapvaal Craton. S. Afr. J. Geol., 93(4): 553-566. Altermann, W. and H~ilbich, I.W., 1991. Structural history of the south-western corner of the Kaapvaal craton and the adjacent Namaqua realm: new observations and reappraisal. Precambrian Res., 52: 133-166.

References

773

Altermann, W. and Herbig, H.G., 1991. Tidal flat deposits of the Lower Proterozoic Campbell Group along the southwestern margin of the Kaapvaal craton, Northern Cape Province, South Africa. J. Afr. Earth Sci., 13: 415-435. Altermann, W. and Kazmierczak, J., 2001. The role of coccoid cyanobacteria in the formation of late Archean carbonates. In: K.F. Cassidy, J.M. Dunphy and M.J. VanKranendonk (Eds.), Extended Abstract Volume, 4th International Archaean Symposium, 24-28 September 2001, Perth, Western Australia, pp. 219-221. Altermann, W. and Nelson, D.R., 1998. Sedimentation rates, basin analysis and regional correlations of three Neoarchaean and Palaeoproterozoic sub-basins of the Kaapvaal craton as inferred from precise U-Pb zircon ages from volcaniclastic sediments. Sediment. Geol., 120: 225-256. Altermann, W. and Schopf, J.W., 1995. Microfossils from the Neoarchean Campbell Group, Griqualand West Sequence of the Transvaal Supergroup, and their paleoenvironmental and evolutionary implications. Precambrian Res., 75: 65-90. Altermann, W. and Siegfried, H.P., 1997. Sedimentology and facies development of an Archaean shelf-carbonate platform transition, Kaapvaal Craton, as deduced from a deep borehole at Kathu, South Africa. J. Afr. Earth Sci., 24(3): 391-410. Altermann, W. and Wotherspoon, J.M., 1995. The carbonates of the Transvaal and Griqualand West Sequences of the Kaapvaal craton, with special reference to the Lime Acres limestone deposit. Mineral. Deposita, 30:124-134. Altschuler, Z.S., Dwornik, E.J. and Kramer, H., 1963. Transformation of montmorillonite to kaolinite during weathering. Science, 141: 148-152. Amelin, Y.V., Der-Chuen, L., Halliday, A.N. and Pidgeon, R.T., 1999. Nature of Earth's earliest crust from hafnium isotopes in single detrital zircons. Nature, 399: 252-255. Amelin, Y.V., Heamen, L.M. and Semenov, V.S., 1995. U-Pb geochronology of layered mafic intrusions in the eastern Baltic Shield: implication for the timing and duration of Palaeoproterozoic continental rifting. Precambrian Res., 75(1/2): 31-46. Amemiya, Y. and Nakayama, O., 1984. The chemical composition and metal adsorption capacity of the sheath material isolated from Microo'stis, Cyanobacteria. Japanese J. Limnology, 45: 187-193. Amireh, B.S., Schneider, W. and Abed, A.M., 1994. Evolving fluvial-transitional-marine deposition through the Cambrian sequence of Jordan. Sediment. Geol., 89: 65-90. Amok (Canada) Ltd., 1974. The Carswell circular structure and Cluff Lake uranium orebodies. In: G.R. Parslow (Ed.), Fuels: A Geological Appraisal. Saskatchewan Geol. Soc., Spec. Publ., 2, pp. 47-61. Anbar, A.D. and Knoll, A.H., 2002. Proterozoic ocean chemistry and evolution: a bioinorganic bridge? Science, 297:1137-1142. Anders, M.H. and Sleep, N.H., 1992. Magmatism and extension: the thermal and mechanical effects of the Yellowstone hotspot. J. Geophys. Res., 97: 15,379-15,393. Anderson, D.L., 1979. The upper mantle transition region: eclogite? Geophys. Res. Lett., 6: 433-436. Anderson, D.L., 1982. Hotspots, polar wander, Mesozoic convection and the geoid. Nature, 297: 391-393. Anderson, D.L., 1998. The scales of mantle convection. Tectonophysics, 284: 1-17. Andreoli, M.A.G., Ashwal, L.D., Hart, R.J. and Huizenga, J.M., 1999. A Ni and PGE-enriched quartz norite impact melt complex in the Late Jurassic Morokweng impact structure, South Africa. In: B.O. Dressier and V.O. Sharpton (Eds.), Large Meteorite Impacts and Planetary Evolution. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 339, pp. 91-108.

774

References

Andronikov, A.V. and Foley, S.E, 2001. Trace element and Nd-Sr isotopic composition of ultramafic lamprophyres from the East Antarctic Beaver Lake area. Chem. Geol., 175: 291-305. Anhaeusser, C.R., 1971. Cyclic volcanicity and sedimentation in the evolutionary development of Archaean greenstone belts of shield areas. In: J.E. Glover (Ed.), Symposium on Archaean rocks. Geol. Soc. Australia, Sydney, Spec. Publ., 3, pp. 57-70. Anhaeusser, C.R., 1984. Structural elements of Archaean granite-greenstone terranes as exemplified by the Barberton Mountain Land, southern Africa. In: A. Kr6ner and R. Greiling (Eds.), Precambrian Tectonics Illustrated. E. Schweizerbarts'che Verlagsbuchhandlung, Stuttgart, pp. 57-78. Anhaeusser, C.R., 1985. Archean layered ultramafic complexes in the Barberton Mountainland, South Africa. In: L.D. Ayres, EC. Thurston, K.D. Card and W. Weber (Eds.), Evolution of Archean supracrustal sequences. Geol. Assoc. Can. Spec. Paper, 28, pp. 281-301. Anonymous, 1972. Penrose field conference on ophiolites. Geotimes, 17: 24-25. Appel, EW.U. and La Berge, G.L. (Eds.), 1987. Precambrian Iron-Formations. Theophrastus Publications, Athens, 674 pp. Appel, EW.U., Fedo, C.M., Moorbath, S. and Myers, J.S. 1998. Recognizable primary volcanic and sedimentary features in a low-strain domain of the highly deformed, oldest known (~3.7-3.8 Gyr) Greenstone Belt, Isua, West Greenland. Terra Nova, 10: 57-62. Appel, EW.U., Moorbath, S. and Myers, J.S., 2003. Isuasphaera isua (Pflug) revisited. Precambrian Res. (in press). Appel, EW.U., Rollinson, H.R. and Touret, J.L.R., 2001. Remnants of an early Archean ( > 3.75 Ga) sea-floor, hydrothermal system in the Isua Greenstone Belt. Precambrian Res., 112: 27-49. Arai, S., 1997. Origin of podiform chromitites. J. Asian Earth Sci., 15:303-310. Arkani-Hamed, J. and Jolly, W.T., 1989. Generation of Archean tonalites. Geology, 17: 307-310. Armitage, A.E., 1998. Geology of the Sandhill Zn-Cu-Pb-Ag prospect and economic potential of the Gibson-MacQuoid greenstone belt, District of Keewatin, NWT. Ph.D. Thesis (unpubl.), University of Western Ontario, London, Ontario, 254 pp. Armstrong, R.A., Compston, W., de Wit, M.J. and Williams, I.S., 1990. The stratigraphy of the 3.53.2 Ga Barberton Greenstone Belt revisited: a single zircon ion microprobe study. Earth Planet. Sci. Lett., 101: 90-106. Armstrong, R.A., Compston, W., Retief, E.A., Williams, I.S. and Welke, H.J., 1991. Zircon ion microprobe studies bearing on the age and evolution of the Witwatersrand Triad. Precambrian Res., 53: 243-266. Armstrong, R.L., 1981. Radiogenic isotopes: the case for crustal recycling on a near-steady-state no-continental-growth Earth. Phil. Trans. R. Soc. London Ser. A, 301: 443-472. Armstrong, R.L., 1991. The persistent myth of crustal growth. Aust. J. Earth Sci., 38:613-630. Arnaud, E. and Eyles, C.H., 2001. Giant cross-bedded sandstones in the Neoproterozoic Port Askaig Formation, Scotland: paleogeographic implications. Abstract, Abstracts with Programs, Geological Society of America Annual Meeting, Vol. 33, p. A-75. Arndt, N.T., 1976. Melting relations of ultramafic lavas (komatiites) at 1 atm and high pressure. Carn. Inst. Wash. Y.B., 75: 555-562. Arndt, N.T., 1977a. Mineralogical and chemical variations in two thick, layered komatiitic lava flows. Carn. Inst. Wash. Y.B., 76: 494-502. Arndt, N.T., 1977b. Thick, layered periodotite-gabbro lava flows in Munro Township, Ontario. Can. J. Earth Sci., 14: 2620-2637. Arndt, N.T., 1977c. Ultrabasic magmas and high-degree melting of the mantle. Contrib. Mineral. Petrol., 64:205-211. Arndt, N.T., 1986. Differentiation of komatiite flows. J. Petrol., 27: 279-301.

References

775

Arndt, N.T., 1994. Archean komatiites. In: K.C. Condie (Ed.), Archean Crustal Evolution. Developments in Precambrian Geology, 11. Elsevier, Amsterdam, pp. 11--44. Arndt, N.T., 1999. Why was flood volcanism on submerged continental platforms so common in the Precambrian? Precambrian Res., 97: 155-164. Arndt, N.T. and Chauvel, C., 1990. Crust of the Hadean Earth. Bull. Geol. Soc. Denmark, 39: 145-151. Arndt, N.T. and Fleet, M.E., 1979. Stable and metastable pyroxene crystallization in layered komatiite flows. Amer. Mineral., 64: 856-864. Arndt, N.T. and Nisbet, E.G. (Eds.), 1982a. Komatiites. George Allen and Unwin, London, 526 pp. Arndt, N.T. and Nisbet, E.G., 1982b. What is a komatiite? In: N.T. Arndt and E.G. Nisbet (Eds.), Komatiites. George Allen and Unwin, London, pp. 19-28. Arndt, N.T., Albarbde, E and Nisbet, E.G., 1997. Mafic and ultramafic magmatism. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 233-254. Arndt, N.T., Albarb~de, E, M. Cheadle, M., Ginibre, C., Herzberg, C., Jenner, G., Chauvel, C. and Lahaye, Y., 1998. Were komatiites wet? Geology, 26: 739-742. Arndt, N.T., Lehnert, K. and Vasil'ev, Y., 1995. Meimechites: highly magnesian alkaline magmas from the subcontinental lithosphere? Lithos, 34: 41-59. Arndt, N.T., Naldrett, A.J. and Pyke, D.R., 1977. Komatiitic and iron-rich tholeiitic lavas of Munro Township, northeast Ontario. J. Petrol., 18, 319-369. Arndt, N.T., Teixeira, N.A. and White, W.M., 1989. Bizarre geochemistry of komatiites from the Crixas greenstone belt, Brazil. Contrib. Mineral. Petrol., 101:187-197. Arndt, N.T., Todt, W., Chauvel, C., Tapfer, M. and Weber, K., 1991. U-Pb zircon age and Nd isotopic composition of granitoids, charnockites and supracrustal rocks from Heimefrontfjella, Antarctica. Geol. Rundschau, 80: 759-777. Arnott, R.W.C. and Hand, B.M., 1989. Bedforms, primary structures and grain fabric in the presence of supended sediment rain. J. Sediment. Petrol., 59: 1062-1069. Arora, M., Govil, EK., Charan, S.N., Uday Raj, B., Balaram, V., Manikyamba, C., Chatterjee, A.K. and Naqvi, S.M., 1995. Geochemistry and origin of Archean banded iron formation from the Bababudan Schist Belt, India. Econ. Geol., 90: 2040-2057. Arp, G., Reimer, A. and Reitner, J., 2001. Photosynthesis-induced biofilm calcification and calcium concentrations in Phanerozoic oceans. Science, 292:1701-1704. Artemieva, N.A. and Shuvalov, V.V., 2002. Shock metamorphism on the ocean floor (numerical simulations). Deep Sea Research II, 49: 959-968. Asahara, Y., Ohtani, E. and Suzuki, A., 1998. Melting relations of hydrous and dry mantle compositions and the genesis of komatiites. Geophy. Res. Lett., 25: 2201-2204. Asami M., Suzuki, K. and Grew, E.S., 2002. Chemical Th-U-total Pb dating by electron microprobe analysis of monazite, xenotime and zircon from the Archaean Napier Complex, East Antarctica: evidence for ultra-high-temperature metamorphism at 2400 Ma. Precambrian Res., 114: 249-275. Ashley, G.M., 1990. Classification of large-scale subaqueous bedforms; a new look at an old problem. J. Sediment. Petrol., 60:161-172. Ashton, K.E., Card, C.D., Heaman, L.M. and Coolican, J., 1999. Rae Northeast: analytical update and new field results from the north-central Train Lake domain. Summary of Investigations, Saskatchewan Geological Survey, 1999, Vol. 2 (99-4.2), pp. 3-15. Aspler, L.B. and Chiarenzelli, J.R., 1998. Two Neoarchean supercontinents? Evidence from the Paleoproterozoic. Sediment. Geol., 120: 75-104. Aspler, L.B. and Donaldson, J.A., 1985. The Nonacho Basin (Early Proterozoic), Northwest Territories, Canada: sedimentation and deformation in a strike-slip setting. In: K.T. Biddle and

776

References

N. Christie-Blick (Eds.), Strike-Slip Deformation, Basin Bormation and Sedimentation. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ. 37, pp. 193-210. Aspler, L.B., Chiarenzelli, J.R. and Bursey, T.L., 1994. Ripple marks in quartz arenites of the Hurwitz Group, Northwest Territories, Canada: evidence for sedimentation in a vast, Early Proterozoic, shallow, fresh-water lake. J. Sediment. Res., A64(2): 282-298. Aspler, L.B., Chiarenzelli, J.R., Powis, K.B. and Cousens, B.L., 1998. Progress Report: Precambrian geology, Angikuni Lake area, District of Keewatin, Northwest Territories. Current Research, Geological Survey of Canada, 1998-C, pp. 55-66. Aspler, L.B., Chiarenzelli, J.R., Cousens, B.L. and Valentino, D., 1999. Precambrian geology, northern Angikuni Lake, and a transect across the Snowbird Tectonic Zone, western Angikuni Lake, Northwest Territories (Nunavut). Current Research, Geological Survey of Canada, 1999-C, pp. 107-118. Aspler, L.B., Chiarenzelli, J.R. and Cousens, B.L., 2002a. Fluvial, lacustrine, and volcanic sedimentation during initiation of Baker Lake Basin, Western Churchill Province, Nunavut. Abstracts, G AC-MAC Annual Meeting, p. 4. Aspler, L.B., Cousens, B.L. and Chiarenzelli, J.R., 2002b. Griffin gabbro sills (2.11 Ga), Hurwitz Basin, Nunavut, Canada: long-distance transport of magmas in western Churchill Province crust. Precambrian Res., 117: 269-294. Aspler, L.B., Wisotzek, I.E., Chiarenzelli, J.R., Losonczy, M.E, Cousens, B.L., McNicoll, V.J. and Davis, W.J., 2001. Paleoproterozoic intracratonic basin processes, from breakup of Kenorland to assembly of Laurentia: Hurwitz Basin, Nunavut, Canada. Sediment. Geol., 141/142: 287-318. Attoh, K. and Ekwueme, B.N., 1997. The west African Shield. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 517-528. Awramik, S.M., 1992a. The oldest records of photosynthesis. Photosynthesis Res., 33: 75-89. Awramik, S.M., 1992b. The history and significance of stromatolites. In: M. Schidlowski, S. Golubic, M.M. Kimberley, D.M. McKirdy and EA. Trudinger (Eds.), Early Organic Evolution: Implications for Mineral and Energy Resources. Springer-Verlag, Berlin, pp. 435-449. Awramik, S.M. and Semikhatov, M.A., 1979. The relationship between morphology, microstructure, and microbiota in three vertically intergrading stromatolites from the Gunflint Iron Formation. Can. J. Earth. Sci., 16:2319-2330. Awramik, S.M., Schopf, J.W. and Walter, M.R., 1983. Filamentous fossil bacteria from the Archean of Western Australia. Precambrian Res., 20: 357-374. Awramik, S.M., Schopf, J.W. and Walter, M.R., 1988. Carbonaceous filaments from North Pole, Western Australia: Are they fossil bacteria in Archean stromatolites? A discussion. Precambrian Res., 39: 303-309. Ayer, J. and Davis, D.W., 1997. Neoarchean evolution of differing convergent margin assemblages in the Wabigoon Subprovince: geochemical and geochronological evidence from the Lake of the Woods greenstone belt, Superior Province, Northwestern Ontario. Precambrian Res., 81: 155-178. Ayer, J., Amelin, Y., Corfu, E, Kamo, S., Ketchum, J., Kwok, K. and Trowell, N., 2002. Evolution of the Abitibi greenstone belt based on U-Pb geochronology: autochthonous volcanic construction followed by plutonism, regional deformation and sedimentation. Precambrian Res., 115: 63-95. Ayres, L.D., 1977a. Importance of stratigraphy in Early Precambrian volcanic terranes: cyclic volcanism at Setting Net Lake, Northwestern Ontario. In: W.R.A. Baragar, L.C. Coleman and J.M. Hall (Eds.), Volcanic Regimes in Canada. Geol. Assoc. Can. Spec. Paper, 16, pp. 243-264.

References

777

Ayres, L.D., 1977b. A transition from subaqueous to subaerial eruptive environments in the Middle Precambrian Amisk Group at Amisk Lake, Saskatchewan--a progress report. In: Centre for Precambrian Studies, Annual Report, 1977, Univ. Manitoba, pp. 36-51. Ayres, L.D., 1981. Preliminary stratigraphic investigation of the upper felsic-intermediate component of the Early Proterozoic Amisk Group, Amisk Lake, Saskatchewan. In: Centre for Precambrian Studies, Annual Report, 1980, Univ. Manitoba, pp. 36-46. Ayres, L.D., 1982. A subaqueous to subaerial transition zone in the Early Proterozoic metavolcanic sequence, Amisk Lake, Saskatchewan. In: Centre for Precambrian Studies, Annual Report, 1981, Univ. Manitoba, pp. 49-61. Ayres, L.D. and Corfu, F., 1991. Stacking of disparate volcanic and sedimentary units by thrusting in the Archean Favourable Lake greenstone belt, central Canada. Precambrian Res., 50:221-238. Ayres, L.D. and Peloquin, A.S., 2000. Subaqueous, Paleoproterozoic, metarhyolite dome-flow-cone complex, Flin Flon greenstone belt, Manitoba, Canada. Precambrian Res., 101: 211-235. Ayres, L.D., Van Wagoner, N.A. and Ferreira, W.S., 1991. Voluminous shallow-water to emergent phreatomagmatic basaltic volcaniclastic rocks, Proterozoic (~ 1886 Ma) Amisk Lake composite volcano, Flin Flon greenstone belt, Canada. In: R.V. Fisher and G.A. Smith (Eds.), Sedimentation in Volcanic Settings. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 45, pp. 175-187. Babinski, M., Chemale F., Jr. and Van Schmus, W.R., 1995. The Pb/Pb age of the Minas Supergroup carbonate rocks, Quadrihitero Ferrffero, Brazil. Precambrian Res., 72: 235-245. Bahnemann, K.P., 1972. A review of the structure, stratigraphy and metamorphism of the basement rocks in the Messina District, Northern Transvaal. Ph.D. Thesis (unpubl.), University of Pretoria, Pretoria, South Africa, 156 pp. Bailes, A.H., 1986. Chisel-Morgan Lakes project. In: Report of Field Activities, 1986, Manitoba Energy and Mines, Min. Div., pp. 71-76. Bailes, A.H., 1987. Chisel-Morgan Lakes project. In: Report of Field Activities, 1987, Manitoba Energy and Mines, Min. Div., pp. 70-79. Bailes, A.H. and Galley, A.G., 1996. Setting of Paleoproterozoic, volcanic-hosted massive sulphide deposits, Snow Lake, Manitoba. In: G.F. Bonham-Carter, A.G. Galley and G.E.M. Hall (Eds.), EXTECH I: A Multidisciplinary Approach to Massive Sulphide Research in the Rusty Lake and Snow Lake Greenstone Belts. Geol. Surv. Can. Bull., 426, pp. 105-138. Bailes, A.H. and Galley, A.G., 1999. Evolution of the Paleoproterozoic Snow Lake arc assemblage and geodynamic setting for associated volcanic-hosted massive sulphide deposits, Flin Flon belt, Manitoba, Canada. Can. J. Earth Sci., 36:1789-1805. Bailes, A.H. and Syme, E.C., 1989. Geology of the Flin Flon-White Lake area. Manitoba Energy and Mines, Geol. Rep. GR 87-1, 313 pp. Bailey, R.A., 1989. Geological map of the Long Valley caldera, Mono-Inyo craters volcanic chain, and vicinity eastern California. Misc. Investigation Series, U.S. Geol. Surv., Map 1-1993 (1 : 62500), 11 pp. Bailey, R.C., 1999. Gravity-driven continental overflow and Archaean tectonics. Nature, 398: 413-415. Baker, V.R., 2001. Water and the Martian landscape. Nature, 412: 228-236. Baker, V.R., Carr, M.H., Gulick, V.C., Williams, C.R. and Marley, M.S., 1992. Channels and valley networks. In: H.H. Kieffer, B.M. Jakosky, C.W. Snyder and M.S. Matthews (Eds.), Mars. Univ. Arizona Press, Tucson, pp. 493-522. Balagansky, V.V., Timmerman, M.J., Kislitsyn, R.V., Daly, J.S., Balashov, Yu.A., Gannibal, L.E, Sherstennikova, O.G. and Ryungenen, G.I., 1998. Isotopic age of the rocks of the Kolvitsa belt

778

References

and Umba block (northeastern branch of the Lapland granulite belt), Kola Peninsula. Proccedings of the Murmansk State Technical University, 1(3): 19-32 (in Russian). Balagansky, V.V., Timmerman, M.J., Kozlova, N.E. and Kislitsyn, R.V., 2001. A 2.44 Ga syn-tectonic mafic dyke swarm in the Kolvitsa belt, Kola Peninsula, Russia: implications for early Palaeoproterozoic tectonics in the north-eastern Fennoscandian Shield. Precambrian Res., 105: 269-287. Baldwin, D.A., 1988. Geology of the southern part of the Rusty Lake volcanic belt. Manitoba Energy and Mines, Geol. Rep. GR 86-1, 90 pp. Baldwin, S.L., Lister, G.S., Hill, E.J., Foster, D.A. and McDougall, I., 1993. Thermochronologic constraints on the tectonic evolution of active metamorphic core complexes, D'Entrecasteaux Islands, Papua, New Guinea. Tectonics, 12:611-628. Baragar, W. and Scoates, R., 1981. The circum-Superior belt: A Proterozoic plate margin? In: A. Kr6ner (Ed.), Precambrian Plate Tectonics. Developments in Precambrian Geology, 4. Elsevier, Amsterdam, pp. 297-330. Barghoorn, E.S. and Schopf, J.W., 1965. Microorganisms from the late Precambrian of central Australia. Science, 150: 337-339. Barghoorn, E.S. and Schopf, J.W., 1966. Microorganisms three billion years old from the Precambrian of South Africa. Science, 152: 758-763. Barghoorn, E.S. and Tyler, S.A., 1965. Microorganisms from the Gunflint chert. Science, 147: 563-577. Barker, J.G. and Arth, J.G., 1976. Generation of trondhjemitic-tonalitic liquids and Archaean bimodal trondhjemite-basalt suites. Geology, 4: 596-600. Barley, M.E., 1984. Volcanism and hydrothermal alteration, Warrawoona Group, East Pilbara. In: J.R. Muhling, D.I. Groves and T.S. Blake (Eds.), Archaean and Proterozoic Basins of the Pilbara, Western Australia. Publ. Univ. West. Aust., Geol. Dep. Ext. Serv., 9, pp. 23-36. Barley, M.E., 1987. The Archaean Whim Creek belt, an ensialic fault-bounded basin in the Pilbara Block, Australia. Precambrian Res., 37:199-215. Barley, M.E., 1992. A review of Archean volcanic-hosted massive sulfide and sulfate mineralization in Western Australia. Econ. Geol., 87: 855-872. Barley, M.E., 1997. The Pilbara craton. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 657-664. Barley, M.E. and Groves, D.I., 1992. Supercontinent cycles and the distribution of metal deposits through time. Geology, 20:291-294. Barley, M.E., Dunlop, J.S.R., Glover, J.E. and Groves, D.I., 1979. Sedimentary evidence for an Archaean shallow-water volcanic-sedimentary facies, eastern Pilbara Block, Western Australia. Earth Planet. Sci. Lett., 43: 74-84. Barley, M.E., Pickard, A.L. and Sylvester, EJ., 1997. Emplacement of a large igneous province as a possible cause of banded iron formation 2.45 billion years ago. Nature, 385: 55-58. Barnes, S.-J., 1983. A comparative study of olivine and clinopyroxene spinifex flows from Alexo, Abitibi greenstone belt, Canada. Contrib. Mineral. Petrol., 83: 293-308. Barnes, S.-J., 1985. The petrography and geochemistry of komatiite flows from the Abitibi greenstone belt and a model for their formation. Lithos, 18:241-270. Barnes, S.-J. and Often, M., 1990. Ti-rich komatiites from northern Norway. Contrib. Mineral. Petrol., 105: 42-54. Barnes, S.-J., Gorton, M.E and Naldrett, A. J., 1983. A comparative study of olivine and clinopyroxene flows from Alexo, Abitibi greenstone belt, Ontario, Canada. Contrib. Mineral. Petrol., 83: 293-308.

References

779

Barnett, W., Armstrong, R.A. and de Wit, M.J., 1997. Stratigraphy of the upper Neoproterozoic Kango and lower Palaeozoic Table Mountain Groups of the Cape Fold Belt revisited. S. Afr. J. Geol., 100: 237-250. Barnicoat, A.C., Henderson, I.H.C., Knipe, R.J., Yardley, B.W.D., Napier, R.W., Fox, N.P.C, Kenyon, A.K., Muntigh, D.J., Strydom, D., Winkler, K.S., Lawrence, S.R. and Conford, C., 1997. Hydrothermal gold mineralization in the Witwatersrand basin. Nature, 386: 820-824. Baross, J.A. and Hoffman, S.E., 1985. Submarine hydrothermal vents and associated gradient environment as sites for the origin and evolution of life. Orig. Evol. Life, 15: 327-345. Barr, A.C., Nimmo, E, Pappalardo, R.T. and Gaidos, E., 2002. Shear heating and solid-state convection: implications for astrobiology. Abstract, Lunar and Planetary Sciences Conference 33, Houston, Texas, 2002, abstract no. 1545. Barrell, J., 1917. Rhythms and the measurements of geological time. Geol. Soc. Amer. Bull., 28: 745-904. Barrett, T.J. and Fralick, P.W., 1985. Sediment redeposition in Archean iron formation: examples from the Beardmore-Geraldton greenstone belt, Ontario. J. Sediment. Petrol., 55(2): 205-212. Barrett, T.J. and Fralick, P.W., 1989. Turbidites and iron-formations, Beardmore-Geraldton, Ontario: application of a combined ramp/tan model to Archaean clastic and chemical sedimentation. Sedimentology, 36: 221-234. Barron, E.J., Fawcett, P.J., Peterson, W.H., Pollard, D. and Thompson, S.L., 1995. A "simulation" of Mid-Cretaceous climate. Paleoceanography, 10: 953-962. Barth, M.G., Rudnick, R.L., Horn, I., McDonough, W.E, Spicuzza, M.J., Valley, J.W. and Haggerty, S.E., 2001. Geochemistry of xenolithic eclogites from West Africa, Part I: A link between low MgO eclogites and Archaean crust formation. Geochim. Cosmochim. Acta, 65: 1499-1527. Bartley, J.K., Knoll, A.H., Grotzinger, J.P. and Sergeev, V.N., 2000. Lithification and fabric genesis in precipitated stromatolites and associated peritidal carbonates, Mesoproterozoic Billyakh Group, Siberia. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 67, pp. 59-73. Bartley, J.K., Pope, M., Knoll, A.H., Semikhartov, M.A. and Petrov, P.Yu., 1998. A VendianCambrian boundary succession from the northwestern margin of the Siberian Platform: stratigraphy, paleontology, chemostratigraphy, and correlation. Geol. Mag., 135: 473--494. Barton, J.M., Jr., 1979. The Chemical composition, Rb-Sr isotopic systematics and tectonic setting of certain post-kinematic mafic igneous rocks, Limpopo Mobile Belt, southern Africa. Precambrian Res., 9: 57-80. Barton, J.M., Jr., Holzer, L., Kamber, B., Doig, R., Kramers, J.D. and Nyfeler, D., 1994. Discrete metamorphic events in the Limpopo belt, southern Africa: implications for the application of P - T paths in complex metamorphic terrains. Geology, 22: 1035-1038. Bates, R.L. and Jackson, J.A., (Eds.) 1987. Glossary of Geology. American Geological Institute, Alexandra, Virgina, 788 pp. Bau, M. and M611er, P., 1991. Rare-earth elements of Archean iron-formations: physico-chemical controls on Europium mobilization and precipitation. Terra Abstracts, 3: 195. Bau, M. and M611er, P., 1993. Rare-earth element systematics of the chemically precipitated component in early Precambrian iron-formations and the evolution of the terrestrial atmospherehydrosphere-lithosphere system. Geochim. Cosmochim. Acta, 57: 2239-2249. Bau, M., Romer, R.L., Luders, V. and Beukes, N.J., 1999. Pb, O, and C isotopes in silicified Mooidraai dolomite (Transvaal Supergroup, South Africa): implications for the composition of Palaeoproterozoic seawater and "dating" the increase of oxygen in the Precambrian atmosphere. Earth Planet. Sci. Lett., 174: 43-57.

780

References

Bauld, J., 1981. Geobiological role of cyanobacterial mats in sedimentary environments: production and preservation of organic matter. BMR Journal of Australian Geology and Geophysics, 6:307-317. Baum, S.K. and Crowley, T.J., 2001. GCM response to late Precambrian (-~ 590 Ma) ice-covered continents. Geophys. Res. Lett., 28: 583-586. Beaudoin, Y., 1998. Petrology, geochemistry and Sr/Nd isotopic analysis of lamprophyre dykes: Kaminak Lake area, central Churchill Province, Northwest Territories, Canada. M.Sc. Thesis (unpubl.), The University of Ottawa, Ottawa, 28 pp. Beaumont, C., Ellis, S., Hamilton, J. and Fullsack, P., 1995. A geodynamical model of the transition from subduction to collision tectonics and its application to the eastern Trans-Hudson Orogen. In: Trans-Hudson Orogen Transect, Lithoprobe Report No. 48, pp. 231-243. Beck, L.S., 1969. Uranium deposits of the Athabasca Region, Saskatchewan. Saskatchewan Department of Mineral Resources, Report No. 126, pp. 1-139. Becker, T.W. and O'Connell, R.J., 2001. Predicting plate velocities with mantle circulation models. Geochem. Geophys. Geosys., 2001GC000171. B6dard, L. and Ludden, J.N., 1995. Nd-isotope evolution of Archaean plutonic rocks in southeastern Superior Province. Can. J. Earth Sci., 34: 286-298. Beddoe-Stephens, B. and Mason, I., 1991. The volcanogenetic significance of garnet-bearing minor intrusions within the Borrowdale Volcanic Group, Eckdale area, Cumbria. Geol. Mag., 128(5): 505-516. Bell, K., 1985. Geochronology of the Carswell area, northern Saskatchwan. In: R. Lain6, D. Alonso and M. Svab (Eds.), The Carswell structure uranium deposits, Saskatchewan. Geol. Assoc. Can. Spec. Paper, 29, pp. 33-46. Belyatsky, B.V., Tikhomirova, M. and Savva, E.V., 2000. An age and genesis of the Proterozoic alkaline-ultramafic Tiksheozero massif (Northern Karelia): results of the Pb-Sr-Nd study. In: Isotope Dating of Geological Processes: New Methods and Results, Abstracts of the 1st Russian Conference on Isotope Geochronology. GEOS, Moscow, pp. 63-65. Benn, K., Sawyer, E.W. and Bouchez, J.-L., 1992. Orogen parallel and transverse shearing in the Opatica Belt, Quebec: implications for the structure of the Abitibi Subprovince. Can. J. Earth Sci., 29: 2429-2444. Benn, K., Miles, W., Ghassemi, M. R. and Gillet, J., 1994. Crustal structure and kinematic framework of the north-western Pontiac Subprovince, Qu6bec: an integrated structural and geophysical study. Can. J. Earth Sci., 31:271-281. Bennett, V.C., Nutman, A.P. and McCulloch, M.T., 1993. Nd isotopic evidence for transient, highly depleted reservoirs in the early history of the Earth. Earth Planet. Sci. Lett., 119: 299-317. Bentor, Y.K., 1985. The crustal evolution of the Arabo-Nubian massif with special reference to the Sinai Peninsula. Precambrian Res., 28: 1-74. Beresford, S., Cas, R., Lahaye, Y. and Jane, M., 2002. Facies architecture of an Archean komatiitehosted Ni-suphide ore deposit, Victor, Kambalda, Western Australia: implications for komatiite lava emplacement. J. Volcanol. Geotherm. Res., 118: 57-75. Berger, W.H., 1974. Deep-sea accumulation. In: C.A. Burk and C.L. Drake (Eds.), The Geology of Continental Margins. Springer, New York, pp. 213-241. Bergman, S.C., 1987. Lamproites and other potassium-rich igneous rocks: A review of their occurrence, mineralogy and geochemistry. In: J.G. Fitton and B.J.G. Upton (Eds.), Alkaline Igneous Rocks. Geological Society, London, pp. 103-189. Berhe, S.M., 1990. Ophiolites in northeast and east Africa: implications for Proterozoic crustal growth. J. Geol. Soc. London, 147: 41-57.

References

781

Berhe, S.M., 1997. The Arabian-Nubian Shield. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 761-771. Berkner, L.V. and Marshall, L.C., 1965. The history of growth of oxygen in the earth's atmosphere. In: P.J. Brancazio and A.G.W. Cameron (Eds.), The Origin and Evolution of Atmospheres and Oceans. John Wiley and Sons, Chichester, U.K., pp. 102-126. Bernard-Griffiths, J., Peucat, J.J., Postaire, B., Vidal, P., Convert, J. and Moreau, B., 1984. Isotopic data (U-Pb, Rb-Sr, Pb-Pb and Sm-Nd) on mafic granulites from Finnish Lapland. Precambrian Res., 23: 325-348. Berner, E.K. and Berner, R.A., 1996. Global Environment: Water, Air, and Geochemical Cycles. Prentice-Hall, Upper Saddle River, New Jersey, 376 pp. Berner, R.A., 1984. Sedimentary pyrite formation; an update. Geochim. Cosmochim. Acta, 48: 605-615. Berner, R.A., 2001. Modeling atmospheric 02 over Phanerozoic time. Geochim. Cosmochim. Acta, 65: 685-694. Berner, R.A. and Berner, E.K., 1997. Silicate weathering and climate. In: W.F. Ruddiman (Ed.), Tectonic Uplift and Climate Change. Plenum Press, New York, pp. 353-365. Berner, R.A. and Canfield, D.E., 1989. A model for atmospheric oxygen over Phanerozoic time. Amer. J. Science, 289: 333-361. Berner, R.A. and Petsch, S.T., 1998. The sulfur cycle and atmospheric oxygen. Science, 282: 1426-1427. Berry, W.B.N. and Wilde, P., 1978. Progressive ventilation of the oceans--An explanation for the distribution of the lower Paleozoic black shales. Amer. J. Sci., 278: 257-275. Bertout, C. and M6nard, F., 2001. Future solar systems. C. R. Acad. Sci. Paris, S6r. IV 2: 1057-1068. Bettenay, L.F., Bickle, M.J., Boulter, C.A., Groves, D.I., Morant, P., Blake, T.S. and James, B.A., 1981. Evolution of the Shaw Batholith--an Archaean granitoid-gneiss dome in the eastern Pilbara, Western Australia. Geol. Soc. Aust. Spec. Publ., 7:361-372. Beukes, N.J., 1973. Precambrian iron-formations of Southern Africa. Econ. Geol., 68: 960-1004. Beukes, N.J., 1980. Suggestions towards a classification and nomenclature for iron-formation. Trans. Geol. Soc. S. Afr., 83: 285-290. Beukes, N.J., 1987. Facies relations, depositional environments and diagenesis in a major Early Proterozoic stromatolitic carbonate platform to basinal sequence, Campbellrand Subgroup, Transvaal Supergroup, Southern Africa. Sediment. Geol., 54: 1-46. Beukes, N.J., 1996. Sole marks and combined-flow storm event beds in the Brixton Formation of the siliciclastic Archaean Witwatersrand Supergroup, South Africa. J. Sediment. Res., 66: 567-576. Beukes, N.J. and Cairncross, B., 1991. A lithostratigraphic-sedimentological reference profile for the Late Archaean Mozaan Group, Pongola Sequence: application to sequence stratigraphy and correlation with the Witwatersrand Supergroup. S. Afr. J. Geol., 94(1): 44-69. Beukes, N.J. and Klein, C., 1992. Models for iron-formation deposition. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere. Cambridge University Press, Cambridge, pp. 147-151. Beukes, N.J. and Lowe, D.R., 1989. Environmental control on diverse stromatolite morphologies in the 3000 Myr Pongola Supergroup, South Africa. Sedimentology, 36: 383-397. Beukes, N.J., Dorland, H., Gutzmer, J., Nedachi, M. and Ohmoto, H., 2002. Tropical laterites, life on land, and the history of atmospheric oxygen in the Paleoproterozoic. Geology, 30:491-494. B6ziat, D., Bourges, F., Debat, P., Lompo, M., Martin, F. and Tollon, E, 2000. A Paleoproterozoic ultramafic-mafic assemblage and associated volcanic rocks of the Boromo greenstone belt: fractionates originating from island-arc volcanic activity in the West African craton. Precambrian Res., 101: 25-47.

782

References

Bhat, M.I. and Gosh, S.K., 2001. Geochemistry of the 2.51 Ga old Rampur group pelites, western Himalayas: implications for their provenance and weathering. Precambrian Res., 108: 1-16. Bhattacharya, J.P. and Walker, R.G., 1991. River- and wave-dominated depositional systems of the Upper Cretaceous Dunvegan Formation, northwestern Alberta. Bull. Can. Soc. Petrol. Geol., 39: 165-191. Bhattacharya, J.P. and Walker, R.G., 1992. Deltas. In: R.G. Walker and N.P. James (Eds.), Facies Models: Response to Sea-Level Change. Geol. Assoc. Canada, Ontario, pp. 157-177. Bhattacharyya, A. and Morad, S., 1993. Proterozoic braided ephemeral fluvial deposits; an example from the Dhandraul Sandstone Formation of the Kaimur Group, Son Valley, central India. Sediment. Geol., 84:101-114. Bibikova, E.V., Melnikov, V.F. and Avakyan, K.Kh., 1993. Lapland granulites: petrology, geochemistry, isotopic age. Petrology, 1(2): 181-198 (official English translation). Bickle, M.J., 1978. Heat loss from the Earth: a constraint on Archaean tectonics from the relation between geothermal gradients and the rate of plate production. Earth Planet. Sci. Lett., 40:301-315. Bickle, M.J., 1986. Implications of melting for stabilization of the lithosphere and heat loss in the Archean. Earth Planet. Sci. Lett., 80:314-324. Bickle, M., 1993. Plume origin for komatiites. Nature, 365: 390-391. Bickle, M.J., Bettenay, L.E, Boultier, C.A., Groves, D.I. and Morant, P., 1980. Horizontal tectonic interaction of an Archean gneiss belt and greenstones, Pilbara Block, Western Australia. Geology, 8: 525-529. Bickle, M.J., Morant, P., Bettenay, L.E, Boulter, C.A., Blake, T.S. and Groves, D.I., 1985. Archaean tectonics of the Shaw Batholith, Pilbara Block, Western Australia: structural and metamorphic tests of the batholith concept. In: L.D. Ayers, P.C. Thurston, K.D. Card and W. Weber (Eds.), Evolution of Archean Supracrustal Sequences. Geol. Assoc. Canada, Spec. Pap. 28, pp. 325-341. Bickle, M.J., Bettenay, L.F., Chapman, H.J., Groves, D.I., McNaughton, N.J., Campbell, I.H. and de Laeter, J.R., 1989. The age and origin of younger granitic plutons of the Shaw Batholith in the Archaean Pilbara block, Western Australia. Contrib. Mineral. Petrol., 101:361-376. Bickle, M.J., Bettenay, L.E, Chapman, H.J, Groves, D.I., McNaughton, N.J., Campbell, I.H. and de Laeter, J.R., 1993. Origin of the 3500-3300 Ma calc-alkaline rocks in the Pilbara Archaean: isotopic and geochemical constraints from the Shaw Batholith. Precambrian Res., 60:117-149. Bickle, M.J., Nisbet, E.G. and Martin, A., 1994. Archean greenstone belts are not oceanic crust. J. Geol., 102: 121-138. Bills, B.G., 1994. Obliquity--oblateness feedback: Are climatically sensitive values of obliquity dynamically unstable? Geophys. Res. Lett., 21: 177-180. Billstr6m, K. and Weihed, P., 1996. Age and provenance of host rocks and ores in the Paleoproterozoic Skellefte district, northern Sweden. Econ. Geol., 91: 1054-1072. Bingen, B., Demaiffe, D. and van Breeman, O., 1998. The 616 Ma old Egersund basaltic dyke swarm, SW Norway, and Late Neoproterozoic opening of the Iapetus Ocean. J. Geol., 106: 565-574. Binns, R., 2003. Deep marine pumice from the Woodlark and Manus basins, Papua New Guinea. In: J. White, D. Clague and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph (in press). Birck, J.L. and All~gre, C.J., 1988. Magnesium-chromium isotope systematics and the development of the solar system. Nature, 331: 579-584. Bjerrum, C.J. and Canfield, D.E., 2002. Ocean productivity before about 1.9 Gyr ago limited by phosphorus adsorption onto iron oxides. Nature, 417: 159-162. Black, M., 1933. The algal sediments of Andros Island, Bahamas. Philos. Trans. Roy. Soc. London Ser. B, 122: 169-192.

References

783

Black, R.E, 1952. Growth of ice-wedge polygons in permafrost near Barrow, Alaska. Geol. Soc. Amer. Bull., 63: 1235-1236. Black, R.E, 1982. Patterned-ground studies in Victoria Land. Antarctic J. US, 17: 53-54. Blackburn, C.E., Johns, G.W., Ayer, J.A. and Davis, D.W., 1991. Wabigoon Subprovince. In: P.C. Thurston, H.R. Williams, R.H. Sutcliffe and G.M. Stott (Eds.), Geology of Ontario. Ontario Geol. Surv., Calgary, pp. 303-381. Blackwelder, E., 1926. Pre-Cambrian geology of the Medicine Bow Mountains. Geol. Soc. Amer. Bull., 37: 615-658. Blair, T.C. and McPherson, J.G., 1994. Alluvial fans and their natural distinction from rivers based on morphology, hydraulic processes, sedimentary processes, and facies assemblages. J. Sediment. Res., A64: 450--489. Blake, D.A.W., 1956. Geological notes on the region south of Lake Athabasca and Black Lake, Saskatchewan and Alberta. Geol. Surv. Can. Paper, 55-33: 1-12. Blake, T.S., 2001. Cyclic continental mafic tuff and flood basalt volcanism in the Late Archaean Nullagine and Mount Jope Supersequences in the eastern Pilbara, Western Australia. Precambrian Res., 107: 139-177. Blake, T.S. and Barley, M.E., 1992. Tectonic evolution of the Late Archaean to Early Proterozoic Mount Bruce Megasequence Set, Western Australia. Tectonics, 11: 1415-1425. Blank, C., 2002. Could cyanobacteria have provided the source of oxidants for banded iron formation? Abstract, Annual Meeting, Geol. Soc. Amer., Denver, 34, p. 382. Blank, J.G., Knize, M.G. and Nakafuji, G., 2002. Laboratory studies of shock-induced amino acid polymerisation and implications for extraterrestrial delivery and impact-generated production of biomolecules. Abstract, Lunar and Planetary Sciences Conference 33, Houston, Texas, 2002, abstract no. 2075. Blasband, B., White, S., Brooijmans, P., de Boorder, H. and Visser, W., 2000. Late Proterozoic extensional collapse in the Arabian-Nubian Shield. J. Geol. Soc. London, 157: 615-628. Blatt, H., Middleton, G. and Murray, R., 1980. Origin of Sedimentary Rocks, 2nd edn. Prentice-Hall, Engelwood Cliffs, New Jersey, 782 pp. Blaxland, A.B., 1974. Geochemistry and geochronology of chemical weathering, Butler Hill Granite, Missouri. Geochim. Cosmochim. Acta, 38: 843-852. Bleeker, W., 2003. Archaean tectonics: a review, with illustrations from the Slave craton. In: C.M.R. Fowler, C.J. Ebinger and C.J. Hawkesworth (Eds.), The early Earth, physical, chemical and biological development. Geol. Soc., London, Spec. Publ., 199 (in press). Bleeker, W., Ketchum, J.W.E, Jackson, V.A. and Villeneuve, M.E., 1999. The Central Slave Basement Complex, Part I: its structural topology and autochthonous cover. Can. J. Earth Sci., 36: 1083-1109. Blenkinsop, T., Fedo, C.M., Bickle, M.J., Eriksson, K.A., Martin, A., Nisbet, E.G. and Wilson, J.E, 1993. Ensialic origin for the Ngezi Group, Belingwe greenstone belt, Zimbabwe. Geology, 21: 1135-1138. Blenkinsop, T., Martin, A., Jelsma, H.A. and Vinyu, M.L., 1997. The Zimbabwe craton. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 567-580. Blewett, R.S., 2000. Horizontal vs. vertical tectonics and the dome-and-basin geometry of the Archaean Pilbara Craton, Western Australia: implications for mineral deposits. Abstract, 15th Aust. Geol. Conv., Sydney, July 2000, Geol. Soc. Australia, 59, p. 44. Blewett, R.S., 2002. Archaean tectonic processes: a case for horizontal shortening in the North Pilbara granite-greenstone terrane, Western Australia. Precambrian Res., 113: 87-120.

784

References

Blewett, R.S., Wellman, E, Ratajkoski, M. and Huston, D.I., 2000. Atlas of North Pilbara Geology and Geophysics, 1 : 1.5 Million Scale. Australian Geol. Surv. Org., Record 2000/4, 36 pp. Blichert-Toft, J. and Arndt, N.T., 1999. Hf isotope compositions of komatiites. Earth Planet. Sci. Lett., 171:439--451. Blichert-Toft, J. and Frei, R., 2001. Complex Sm-Nd and Lu-Hf systematics in metamorphic garnets from the Isua supracrustal belt, West Greenland. Geochim. Cosmochim. Acta, 65:3177-3187. Bloomer, 1995. Early arc volcanism and the ophiolite problem: a perspective from drilling in the Western Pacific. In: B. Taylor and J. Natland (Eds.), Active Margins and Marginal Basins of the Western Pacific. American Geophysical Union Monograph Series, 88, pp. 1-30. Bluck, B.J., 1986. Upward coarsening sedimentation units and facies lineages, Old Red sandstone, Scotland. Trans. Roy. Soc. Edinburgh. Earth Sci., 77:251-264. Blum, M.D., 1994. Genesis and architecture of incised valley fill sequences: a late Quaternary example from the Colorado River, Gulf Coastal Plain of Texas. In: P. Weimer and H.W. Posamentier (Eds.), Siliciclastic sequence stratigraphy: recent developments and applications. Amer. Assoc. Petrol. Geol. Mem., 58: 259-283. Blum, M.D. and T6rnqvist, T.E., 2000. Fluvial responses to climate and sea-level change: a review and look forward. Sedimentology, 47 (Suppl. 1): 2-48. Board, W.S., 2002. Tectonothermal evolution of the southern H.U. Sverdrupfjella, western Dronning Maud Land, Antarctica. Ph.D. Thesis (unpubl.), University of Cape Town, 174 pp. Bogdanova, S.V., 1996. High-grade metamorphism of 2.44-2.4 Ga age in mafic intrusions of the Belomorian Belt in the northeastern Baltic Shield. In: T.S. Bruwer (Ed.), Precambrian Crustal Evolution in the North Atlantic Region. Geol. Soc. Spec. Publ., 112, pp. 69-90. Bogdanova, S.V., Bibikova, E.V. and Postnikov, A.V., 1999. A remnant of the Palaeoproterozoic magmatic arc beneath Moscow. In: Yu.V. Kariakin and M.V. Mints (Eds.), Early Precambrian: Genesis and Evolution of the Continental Crust (Geodynamics, Petrology, Geochronology, Regional Geology), Abstracts. GEOS, Moscow, pp. 23-24. Bogdanova, S.V., Gorbatschev, R. and Stephenson, R.A., 200 l a. EUROBRIDGE: Palaeoproterozoic accretion of Fennoscandia and Sarmatia. Tectonophysics, 339: vii-x. Bogdanova, S.V., Page, L.M., Skridlaite, G. and Taran, L.N., 2001 b. Proterozoic tectonothermal history in the western part of the East European Craton: 4~ geochronological constraints. Tectonophysics, 339: 39-66. Boher, M., Abouchami, W., Michard, A., Albarede, E and Arndt, N.T., 1992. Crustal growth in West Africa at 2.1 Ga. J. Geophys. Res., 97: 345-369. B6hm, C.O., Heaman, L.M., Creaser, R.A. and Corkery, M.T., 2000. Discovery of pre-3.5 Ga exotic crust at the northwestern Superior Province margin, Manitoba. Geology, 28: 75-78. Bohor, B.E, Modreski, P.J. and Foord, E.E., 1987. Shocked quartz in the Cretaceous-Tertiary boundary clays; evidence for a global distribution. Science, 236: 705-709. Boily, M. and Dion, C., 2002. Geochemistry of boninite-type volcanic rocks in the Frotet-Evans greenstone belt, Opatica subprovince, Quebec: implications for the evolution of Archaean greenstone belts. Precambrian Res., 115: 349-371. Bond, G.C., Nickeson, P.A. and Kominz, M.A., 1984. Breakup of a supercontinent between 625 and 555 Ma: New evidence and implications for continental histories. Earth Planet. Sci. Lett., 70: 325-345. Borg, G. and Armstrong, R., 2002. Isotopic SHRIMP age dating of zircons from igneous basement and rhyolitic cover rocks at Skorpion, southern Namibia. Abstract, 1l th Quadrennial IAGOD Symposium and Geocongress 2002, July 2002, Windhoek, Geological Society of Namibia (CD-ROM).

References

785

Borradaile, G.J., 1982. Comparison of Archean structural styles in two belts of the Canadian Superior Province. Precambrian Res., 19:179-189. Bosch, P.J.A., Eriksson, P.G. and Snyman, C.P., 1993. The Wolkberg Group in the northern Transvaal: Palaeoenvironment derived from sedimentology and geochemistry. S. Afr. J. Geol., 96(4): 190-204. Bose, P.K. and Chakraborty, P.P., 1994. Marine to fluvial transition: Proterozoic Upper Rewa Sandstone, Maihar, India. Sediment. Geol., 89: 285-302. Bose, P.K., Chakraborty, S. and Sarkar, S., 1999. Recognition of ancient eolian longitudinal dunes: a case study in Upper Bhander Sandstone, Son Valley, India. J. Sediment. Res., 69: 86-95. Bose, P.K., Chaudhuri, A.K. and Seth, A., 1988. Facies, flow and bedform patterns across a stormdominated inner continental shelf: Proterozoic Kaimur Formation, Rajasthan, India. Sediment. Geol., 59: 275-293. Bose, P.K., Sarkar, S., Chakrabarty, S. and Banerjee, S., 2001. Overview of the Meso- to Neoproterozoic evolution of the Vindhyan basin, central India. Sediment. Geol., 141/142: 395-419. Bouhallier, H., Choukroune, P. and Ballevre, M., 1993. Diapirism, bulk homogeneous shortening and transcurrent shearing in the Archaean Dharwar craton, the Holenarsipur area, southern India. Precambrian Res., 63: 43-58. Boulter, C.A., B ickle, M.J., Gibson, B. and Wright, R.K., 1987. Horizontal tectonics pre-dating upper Gorge Creek Group sedimentation, Pilbara Block, Western Australia. Precambrian Res., 36: 241-258. Bouma, A.H., 1969. Methods for the Study of Sedimentary Structures. Wiley Interscience, New York, 458 pp. Bouma, A.H., Brenner, R.L. and Knebel, H.J., 1982. Continental shelf and epicontinental seaways. In: P.A. Scholle and D. Spearing (Eds.), Sandstone Depositional Environments. Amer. Assoc. Petrol. Geol., Tulsa, Oklahoma, pp. 281-327. Bouougri, E. and Porada, H., 2002. Mat-related sedimentary structures in Neoproterozoic peritidal passive margin deposits of the West African craton (Anti-Atlas, Morocco). Sediment. Geol., 153: 85-106. Bourgeois, J. and Johnson, S.Y., 1998. Geologic evidence of earthquakes at the Snohomish Delta, Washington, in the past 1200 yr. Abstracts with Programs, Annual Meeting Geological Society of America, Boston, Massachusetts, 2001, 30, p. 106. Bourgeois, J., Hansen, T.A., Wiberg, P.L. and Kauffman, E.G., 1993. A tsunami deposit at the Cretaceous-Tertiary boundary in Texas. Science, 241: 567-570. Bourgeois, J., Landis, S.P., Mann, C.P. and Pinegina, T., 2001. Sorted stone circles on a tsunami, swept beach, Kamchatka Peninsula, Russia. Abstracts with Programs, Annual Meeting, Geological Society of America, Boston, Massachusetts, 2001, 31, p. 51. Bowring, S.A. and Housh, T., 1995. The Earth's early evolution. Science, 269: 1530-1540. Bowring, S.A. and Ross, G.E., 1985. Geochronology of the Narakay volcanic complex: implications for the age of the Coppermine Homocline and Mackenzie igneous events. Can. J. Earth Sci., 22: 774-781. Bowring, S.A. and Williams, I.S., 1999. Priscoan (4.0(0-4.03 Ga) orthogneisses from northwestern Canada. Contrib. Miner. Petrol., 134: 3-16. Bowring, S.A., Williams, I.S. and Compston, W., 1989. 3.96 Ga gneisses from the Slave Province, NW Territories, Canada. Geology, 17: 971-975. Boyd, ER. and McCallister, R.H., 1976. Densities of fertile and sterile garnet peridotites. Geophys. Res. Lett., 3:509-512.

786

References

Boyd, R., Diessel, C., Wadsworth, J., Chalmers, G., Little, M., Leckie, D. and Zaitlin, B., 1999. Development of a nonmarine sequence stratigraphic model. American Association of Petroleum Geologists Annual Meeting, San Antonio, Texas, USA. Official Program, p. A15. Brack, A., 1998. The Molecular Origins of Life: Assembling Pieces of the Puzzle. Cambridge Univ. Press, Cambridge, 417 pp. Brack, A., 2002a. Origin of life. In: Encyclopedia of Life Sciences. Nature Publishing Group, Vol. 13, pp. 554-560. Brack, A., 2002b. Water, the spring of life. In: G. Horneck and C. Baumstark-Khan (Eds.), Astrobiology: The Quest for the Conditions of Life. Springer, Berlin, pp. 279-288. Bradshaw, J.Y., 1989. Early Cretaceous vein related garnet granulite in Fjordland, southwest New Zealand: a case for infiltration of mantle-derived CO2-rich fluids. J. Geol., 97: 697-717. Brandl, G., 2001. The Limpopo Belt and its environs. Explanation of the Geological map of the Limpopo Belt and its environs (1:500000). Council for Geoscience, Pretoria, South Africa. Brandl, G. and de Wit, M.J., 1997. The Kaapvaal Craton. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Science Publications, Oxford, pp. 581-607. Brasier, M.D. and Lindsay, J.F., 1998. A billion years of environmental stability and the emergence of eukaryotes: new data from Northern Australia. Geology, 26: 555-558. Brasier, M.D. and Shields, G., 2000. Neoproterozoic chemostratigraphy and correlation of the Port Askaig glaciation, Dalradian Supergroup of Scotland. J. Geol. Soc., London, 157:909-914. Brasier, M.D. and Sukhov, S., 1998. The falling amplitude of carbon isotopic oscillations through the Lower to Middle Cambrian; northern Siberia data. Can. J. Earth Sci., 35: 353-373. Brasier, M.D., Rozanov, A.Yu., Zhuravlev, A.Yu., Corfield, R.M. and Derry, L.A., 1994. A carbon isotope reference scale for the Lower Cambrian succession in Siberia: report of IGCP Project 303. Geol. Mag., 131: 767-783. Brasier, M.D., Shields, G., Kuleshov, V. and Zhegallo, E.A., 1996. Integrated chemo- and biostratigraphic calibration of early animal evolution: Neoproterozoic---early Cambrian of SW Mongolia. Geol. Mag., 133: 445-485. Brasier, M., McCarron, G., Tucker, R., Leather, J., Allen, P. and Shields, G., 2000. New U-Pb zircon dates for the Neoproterozoic Ghubrah glaciation and for the top of the Huqf Supergroup, Oman. Geology, 28: 175-178. Brasier, M.D., Green, O.R., Jephcoat, A.P., Kleppe, A.K., Van Kranendonk, M.J., Lindsay, J.F., Steele, A. and Grassineau, N.V., 2002. Questioning the evidence for Earth's oldest fossils. Nature, 416: 76-81. Brassell, S.C., 1992. Biomarkers in sediments, sedimentary rocks and petroleums: Biological origins, geological fate and applications. In: L.M. Pratt, J.B. Comer and S.C. Brassell (Eds.), Geochemistry of Organic Matter in Sediments and Sedimentary Rocks. SEPM Short Course No. 27. Society for Sedimentary Geology, Tulsa, Oklahoma, pp. 29-72. Braterman, P.S. and Cairns-Smith, A.G., 1987a. Photoprecipitation and the banded ironformations--some quantitative aspects. Origins of Life, 17: 221-228. Braterman, P.S. and Cairns-Smith, A.G., 1987b. Iron photoprecipitation and the genesis of the banded iron-formations. In: P.W.U. Appel and G.L. LaBerge (Eds.), Precambrian Iron-Formations. Theophrastus Publications, Athens, pp. 215-245. Braterman, P.S., Cairns-Smith, A.G., Sloper, R.W., Truscott, T.G. and Craw, M., 1984. Photooxidation of iron(II) in water between pH 7.5 and 4.0. J. Chem. Soc., Dalton Transactions: 1441-1445. Breitkopf, J.H., 1988. Iron formations related to mafic volcanism and ensialic rifting in the Southern Margin Zone of the Damara Orogen, Namibia. Precambrian Res., 38:111-130.

References

787

Brenner, R.L., 1980. Construction of process-response models for ancient epicontinental seaway depositional systems using partial analogs. Bull. Amer. Assoc. Petrol. Geol., 64: 1223-1244. Bretz, J.H., 1969. The Lake Missoula floods and the Channeled Scabland. J. Geology, 77: 505-543. Briceno, C., Vivas, A.K., Calvet, N., Hartmann, L., Pacheco, R., Herrera, D., Snyder, J.A. and Andrews, P., 2001. The CIDA-QUEST large-scale survey of Orion OB 1: evidence for rapid disk dissipation in a dispersed stellar population. Science, 291: 93-96. Bridgwater, D., Keto, L., McGregor, V.R. and Myers, J.S., 1976. Archaean gneiss complex of Greenland. In: A. Escher and W.S. Watt (Eds.), Geology of Greenland. GrCnlands Geologiske UndersCgelse, Copenhagen, pp. 18-75. Bridgwater, D., Allaart, J.H., Schopf, J.W., Klein, C., Walter, E.S., Strother, P., Knoll, A.H. and Gorman, B.E., 1981. Microfossil-like objects from the Archaean of Greenland: a cautionary note. Nature, 289:51-53. Bridgwater, D., Scott, D.J., Balagansky, V.V., Timmerman, M.J., Marker, M., Bushmin, A., Alexeyev, N.L. and Daly, J.S., 2001. Age and provenance of early Precambrian metasedimentary rocks in the Lapland-Kola Belt, Russia: evidence from Pb and Nd isotopic data. Terra Nova, 13( 1): 32-37. Broadhurst, F.M., 1988. Seasons and tides in the Westphalian. In: B.M. Besley and G. Kelling (Eds.), Sedimentation in a Synorogenic Basin Complex. Blackie, Glasgow, pp. 264-272. Brocks, J.J., Logan, G.A., Buick, R. and Summons, R.E., 1999. Archean molecular fossils and the early rise of eukaryotes. Science, 285: 1033-1036. Broekhuizen, A., 1998. The geology of the Koedoesrand Formation, north-western Transvaal, and its relationship to the Palala Shear Zone. M.Sc. Thesis (unpubl.), University of Pretoria, South Africa, 89 pp. Brookfield, M.E., 1977. The origin of bounding surfaces in ancient aeolian sandstone. Sedimentology, 24: 303-332. Brooks, C., Ludden, J. and Pigeon, Y., 1982. Volcanism of shoshonite to high-K andesite affinity in an Archean arc environment, Oxford Lake, Manitoba. Can. J. Earth Sci., 19: 55-67. Brooks, E.R., 2000. Geology of a late Paleozoic island arc in the northern Sierra terrane, Field guide to the geology and tectonics of the northern Sierra Nevada. California Department of Conservation, Div. Mines and Geology, Sacramento, California, pp. 53-110. Brooks, G.R. and Holmes, C.W., 1990. Modern configuration of the southwest Florida carbonate slope: development of shelf margin progradation. Marine Geol., 94:301-315. Brosche, P. and Wtinsch, J., 1990. The solar torque--a leak for the angular momentum of the EarthMoon system. In: P. Brosche and J. Stindermann (Eds.), Earth's Rotation from Eons to Days. Springer, Berlin, pp. 141-145. Brown, L.F., Jr. and Fisher, W.L., 1977. Seismic stratigraphic interpretation of depositional systems: examples from Brazilian rift and pull apart basins. In: C.E. Payton (Ed.), Seismic Stratigraphy--Applications to Hydrocarbon Exploration. Amer. Assoc. Petrol. Geol. Mem., 26: 213-248. Brown, M., 1994. The generation, segregation, ascent and emplacement of granite magma--the migmatite-to-crustally-derived-granite connection in thickened orogens. Earth Sci. Rev., 36: 83-130. Brown, S.J.A., Barley, M.E., Krapez, B. and Cas, R.A.F., 2002. The Late Archaean Melita Complex, Eastern Goldfields, Western Australia: shallow submarine bimodal volcanism in a rifted arc environment. J. Volcanol. Geotherm. Res., 115: 303-327. Brun, J.-P., Sokoutis, D. and van den Driessche, J., 1994. Analogue modelling of detachment fault systems and core complexes. Geology, 22: 319-322.

788

References

Bruun, E, 1962. Sea-level rise as a cause of shore erosion. American Society of Civil Engineers Proceedings, Journal of the Waterways and Harbors Division, 88:117-130. Buchan, K.L. and Ernst, R.E., 2003. Diabase dyke swarms and related units in Canada and adjacent regions. Geological Survey of Canada Map 2022A, scale 1:5000000 (in press). Buchan, K.L., Halls, H.C. and Mortensen, J.K., 1996. Paleomagnetism, U-Pb geochronology and geochemistry of Marathon dykes, Superior Province, and comparison with the Fort Frances swarm. Can. J. Earth Sci., 33: 1583-1595. Buchan, K.L., Mertanen, S., Park, R.G., Pesonen, L.J., Elming, S.-A., Abrahamsen, N. and Bylund, G., 2000. Comparing the drift of Laurentia and Baltica in the Proterozoic: the importance of key paleomagnetic poles. Tectonophysics, 319:167-198. Buck, S.G., 1980. Stromatolite and ooid deposits within the fluvial and lacustrine sediments of the Precambrian Ventersdorp Supergroup of South Africa. Precambrian Res., 12:311-330. Buck, S.G. and Minter, W.E.L., 1985. Placer formation by fluvial degradation of an alluvial fan sequence: the Proterozoic Carbon Leader Placer, Witwatersrand Supergroup, South Africa. J. Geol. Soc. London, 142: 757-764. Buczynski, C. and Chafetz, H.S., 1991. Habit of bacterially induced precipitates of calcium carbonate and the influence of medium viscosity on mineralogy. J. Sediment. Petrol., 61: 226-233. Buhlmann, A.L., Cavell, E, Burwash, R.A., Creaser, R.A. and Luth, R.W., 2000. Minette bodies and cognate mica-clinopyroxenite xenoliths from the Milk River area, southern Alberta: Records of a complex history of the northernmost part of the Archean Wyoming Craton. Can. J. Earth Sci., 37: 1629-1650. Buick, R., 1984. Carbonaceous filaments from North Pole, Western Australia, Are they fossil bacteria in Archean stromatolites? Precambrian Res., 24: 157-172. Buick, R., 1988. Carbonaceous filaments from North Pole, Western Australia, Are they fossil bacteria in Archean stromatolites? A reply. Precambrian Res., 39:311-317. Buick, R., 1992. The antiquity of oxygenic photosynthesis: evidence from stromatolites in sulfatedeficient Archean Lakes. Science, 255: 74-77. Buick, R. and Dunlop, J.S.R., 1990. Evaporitic sediments of Early Archaean age from the Warrawoona Group, North Pole, Western Australia. Sedimentology, 37: 247-277. Buick, R., Dunlop, J.S.R. and Groves, D.I., 1981. Stromatolite recognition in ancient rocks: an appraisal of irregularly laminated structures in early Archaean chert-barite unit from North Pole, Western Australia. Alcheringa, 5: 161-181. Buick, R., Des Marais, D.J. and Knoll, A.H., 1995a. Stable isotope compositions of carbonates from the Mesoproterozoic Bangemall Group, northwestern Australia. Chem. Geol., 123: 153-171. Buick, R., Groves, D.I. and Dunlop, J.S.R., 1995b. Abiological origin of described stromatolites older than 3.2 Ga: comment and reply. Comment. Geology, 23: 191. Buick, R., Thornett, J.R., McNaughton, N.J., Smith, J.B., Barley, M.E. and Savage, M., 1995c. Record of emergent continental crust "-3.5 billion years ago in the Pilbara craton of Australia. Nature, 375: 574-577. Bumby, A.J., 2000. The geology of the Blouberg Formation, Waterberg and Soutpansberg Groups in the area of Blouberg Mountain, Northern Province, South Africa. Ph.D. Thesis (unpubl.), University of Pretoria, Pretoria, 325 pp. Bumby, A.J., Eriksson, EG., van der Merwe, R. and Brtimmer, J.J., 2001a. Shear-zone controlled basins in the Blouberg area, Northern Province, South Africa: syn- and post-tectonic sedimentation relating to c. 2.0 Ga reactivation of the Limpopo Belt. J. Afr. Earth Sci., 33:445-461.

References

789

Bumby, A.J., Eriksson, EG., van der Merwe, R. and Maier, W., 200lb. The stratigraphic relationship between the Waterberg Group and the Soutpansberg Group (Northern Province, South Africa): evidence from the Blouberg area. S. Afr. J. Geol., 104: 105-216. Burbidge, E.M., Burbidge, G.R., Fowler, W.A. and Hoyle, E, 1957. Synthesis of the elements in stars. Rev. Mod. Phys., 29: 547-650. Burdett, J.W., Arthur, M.A. and Richardson, M., 1989. A Neogene seawater sulfate isotope age curve from calcareous pelagic microfossils. Earth Planet. Sci. Letters, 94:189-198. Burke, K.D., 1997. Foreword. In" M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Science Publications, Oxford, pp. v-vii. Burkhard, D.J.M., 2003. Thermal interaction between lava lobes. Bull. Volcanol., 65: 136-143. Bume, R.V. and Moore, L.S., 1987. Microbialites" Organosedimentary deposits of benthic microbial communities. Palaios, 2/3: 241-254. Bums, S.J. and Matter, A., 1993. Carbon isotopic record of the latest Proterozoic from Oman. Eclogae Geologicae Helvetiae, 86: 595-607. Burwash, R.A., Chacko, T., Muehlenbachs, K. and Bouzidi, Y., 2000. Oxygen isotope systematics of the Precambrian basement of Alberta: implications for Paleoproterozoic and Phanerozoic tectonics in northwestern Alberta. Can. J. Earth Sci., 37" 1611-1628. Busby, C.J. and Ingersoll, R.V. (Eds.), 1995. Tectonics of Sedimentary Basins. Blackwell Science, - Cambridge, 579 pp. Busby-Spera, C.J., 1984. Large volume rhyolite ash-flow eruptions and submarine caldera collapse in the Lower Mesozoic Sierra Nevada, California. J. Geophys. Res., 89:8417-8427. Busso, M., Gallino, R. and Wasserburg, G.J., 1999. Nucleosynthesis in asymptotic giant branch stars: relevance for galactic enrichment and solar system formation. Annu. Rev. Astron. Astrophys., 37: 239-309. Button, A., 1973. A regional study of the stratigraphy and development of the Transvaal Basin in eastern and northeastern Transvaal. Ph.D. Thesis (unpubl.), Univ. of the Witswatersrand, Johannesburg, 352 pp. Button, A. (Rapporteur), 1982. Sedimentary iron deposits, evaporites and phosphorites--state of the art report. In: H.D.Holland and M.Schidlowski (Eds.), Mineral Deposits and the Evolution of the Biosphere. Dahlem Konferenzen, Springer-Verlag, Berlin, pp. 259-273. Byerly, G.R., 1999. Komatiites of the Mendon Formation: late-stage ultramafic volcanism in the Barberton greenstone belt. In" D.R. Lowe and G.R. Byerly (Eds.), Geologic Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer. Spec. Paper, 329: pp. 189-212. Byerly, G.R. and Lowe, D.R., 1994. Spinel from Archaean impact spherules. Geochim. Cosmochim. Acta, 58: 3469-3486. Byerly, G.R., Kr6ner, A., Lowe, D.R., Todt, W. and Walsh, M.W., 1996. Prolonged magmatism and time constraints for sediment deposition in the early Archaean Barberton greenstone belt: evidence from the Upper Onverwacht and Fig Tree Groups. Precambrian Res., 78: 125-138. Byerly, G.R., Lowe, D.R. and Walsh, M.M., 1986. Stromatolites from the 3,300-3,500 Myr Swaziland Supergroup, Barberton Mountain Land, South Africa. Nature, 319:489-491. Byerly, G.R., Lowe, D.R., Wooden, J.L. and Xie, X., 2002. An Archaean impact layer from the Pilbara and Kaapvaal cratons. Science, 297: 1325-1327. Cabrol, N.A. and Grin, E.A., 1999. Distribution, classification, and ages of Martian impact crater lakes. Icarus, 142:160-172. Caldeira, K. and Rampino, M.R., 1991. The mid-Cretaceous superplume, carbon dioxide, and global warming. Geophys. Res. Lett., 18: 987-990.

790

References

Callaghan, C.C., Eriksson, EG. and Snyman, C.E, 1991. The sedimentology of the Waterberg Group in the Transvaal, South Africa: An overview. J. Afr. Earth Sci., 13: 121-139. Calver, C.R. and Lindsay, J.E, 1998. Ediacarian sequence and isotope stratigraphy of the Officer Basin, South Australia. Australian J. Earth Sci., 45: 513-532. Calvert, A.J. and Ludden, J.N., 1999. Archean continental assembly in the southeastern Superior Province of Canada. Tectonics, 18: 412-429. Calvert, A., Sawyer, E.W., Davis, W.J. and Ludden, J.N., 1995. Archaean subduction inferred from seismic images of a mantle suture in the Superior Province. Nature, 375: 670-674. Camacho, A. and Fanning, C.M., 1995. Some isotopic constraints on the evolution of the granulite and upper amphibolite facies terranes in the eastern Musgrave Block, central Australia. Precambrian Res., 71: 155-181. Cameron, A.G.W., 1957. Stellar evolution, nuclear astrophysics and nucleogenesis. Chalk River Rep. AECL (Atomic Energy of Canada): CRL-41. Cameron, A.G.W., 1995. The first ten million years in the solar nebula. Meteoritics, 30: 131-161. Cameron, A.G.W., 1998. Two cradles for the heavy elements. Nature, 391:228-231. Cameron, A.G.W. and Truman, J.W., 1977. The supernova trigger for formation of the solar system. Icarus, 30:447-461. Cameron, A.G.W., H6flich, P., Myers, P.C. and Clayton, D.D., 1995. Massive supernovae, Orion Gamma Rays and the formation of the solar system. Astrophys. J., 447: L53-L57. Cameron, A.G.W., Vanhala, H. and H6flich, P., 1997. Some aspects of triggered star formation. In: T.J. Bernatowicz and E.K. Zinner (Eds.), Astrophysical Implications of the Laboratory Study of Presolar Materials. Proceedings, AIP conference, St. Louis, Missouri, The American Institute of Physics, Woodbury, New York, pp. 665-693. Campbell, I.H., 2001. Identification of ancient mantle plumes. In: R.E. Ernst and K.L. Buchan (Eds.), Mantle Plumes: Their Identification through Time. Geol. Soc. Amer. Spec. Paper, 352: pp. 5-21. Campbell, I.H. and Griffiths, R.W., 1990. Implications of mantle plume structure for the evolution of flood basalts. Earth Planet. Science Lett., 99: 79-83. Campbell, I.H. and Griffiths, R.W., 1992. The changing nature of mantle hotspots through time: implications for the chemical evolution of the Mantle. J. Geol., 92: 497-523. Campbell, I.H. and Hill, R.I., 1988. A two-stage model for the formation of the granite-greenstone terrane of the Kalgoorlie-Norseman area, Western Australia. Earth Planet. Sci. Lett., 90:11-25. Campbell, I.H., Griffiths, R.W. and Hill, R.I., 1989. Melting in an Archean mantle plume: heads it's basalts, tails it's komatiites. Nature, 339: 697-699. Canfield, D.E., 1996. Evolution of the sulfur cycle. In: S.H. Bottrell (Ed.), Proceedings, 4th International Symposium on the Geochemistry of the Earth's Surface (GES-4), University of Leeds, pp. 23-27. Canfield, D.E., 1998. A new model for Proterozoic ocean chemistry. Nature, 396: 450-453. Canfield, D.E., 2001. Isotope fractionation by natural populations of sulfate-reducing bacteria. Geochim. Cosmochim. Acta, 65:1117-1124. Canfield, D.E. and Raiswell, R., 1999. The evolution of the sulfur cycle. Amer. J. Sci., 299: 697-723. Canfield, D.E. and Teske, A., 1996. Late Proterozoic rise in atmospheric oxygen concentration inferred from phylogenetic and sulphur-isotope studies. Nature, 382: 127-132. Canfield, D.E. and Thamdrup, B., 1994. The production of 34S-depleted sulfide during bacterial disproportionation of elemental sulfur. Science, 266:1973-1975. Canfield, D.E., Habicht, K.S. and Thamdrup, B., 2000. The Archean sulfur cycle and the early history of atmospheric oxygen. Science, 288:658-661.

References

791

Canning, J.C., Henney, R, Morrison, M.A., Van Calsteren, RW.C., Gaskarth, J.W. and Swarbrick, A., 1998. The Great Glen Fault: a major vertical crustal boundary. J. Geol. Soc. London, 155: 425-428. Cant, D.J. and Hein, EJ., 1986. Depositional sequences in ancient shelf sediments: some contrasts in style. In: R.J. Knight and J.R. McLean (Eds.), Shelf Sands and Sandstones. Can. Soc. Pet. Geol. Mem., 11, pp. 303-312. Canup, R.M. and Angor, C.B., 2000. Accretion of the terrestrial planets and the Earth-Moon system. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 113-129. Canup, R.M. and Asphaug, 2001. Origin of the Moon in a giant impact near the end of the Earth's formation. Nature, 412:708-712. Canup, R.M. and Righter, K., 2000. Origin of the Earth and Moon. Univ. Arizona Press, Houston, 555 pp. Car, D. and Ayres, L.D., 1991. A thick dacitic debris flow sequence, Lake of the Woods greenstone terrane, central Canada: resedimented products of Archean vulcanian, plinian, and dome-building eruptions. Precambrian Res., 50: 239-260. Card, K.D., 1990. A review of the Superior Province of the Canadian Shield, a product of Archean accretion. Precambrian Res., 48: 99-156. Carey, S., 1997. Influence of convective sedimentation on the formation of widespread tephra fall layers in the deep sea. Geology, 25: 839-842. Carey, S.W., 1976. The Expanding Earth. Elsevier, Amsterdam, 488 pp. Carey, S.W., 1996. Earth Universe Cosmos. Univ. Tasmania Press, Hobart, 231 pp. Carlisle, D.B. and Braman, D.R., 1991. Nanometre-size diamonds in the Cretaceous/Tertiary boundary clay of Alberta. Nature, 352: 708-709. Carlson, C., 1993. Recognition of an Early Proterozoic ophiolite allochthon in the Birimian Nangodi greenstone belt, NE Ghana. Geol. Soc. Amer. Abstr. Progr., 25: p. 18. Carlson, C., 1997. Paleoproterozoic intraoceanic accretionary tectonics recorded in the approximately 2.1 Ga Birimian Nangodi greenstone belt, NE Ghana. Eos, 78: 720. Carlson, R.W. and Lugmair, G.W., 2000. Timescales of planetesimal formation and differentiation based on extinct and extant radioisotopes. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 25--44. Carr, M.H., 1996. Water on Mars. Cambridge Univ. Press, Cambridge, 229 pp. Cas, R.A.F and Wright, J.V., 1987. Volcanic Successions, Modern and Ancient. A geological Approach to Processes, Products and Successions. George Allen and Unwin, London, 528 pp. Cas, R., Self, S. and Beresford, S., 1999. The behaviour of the fronts of komatiite lavas in medial and distal settings. Earth Panet. Sci. Lett., 172: 127-139. Cas, R.A.E, Beresford, S.W. and Appel, P., 2001. Reassessing the earliest known paleoenvironments on Earth; the 3.7-3.8 Ga Isua Greenstone Belt, Greenland. In: K.E Cassidy, J.M. Dunphy and M.J. Van Kranendonk (Eds.), Extended Abstract volume, 4th Intern. Archaean Symp. Perth, 2001. AGSO--Geoscience Australia, Record 2001/37, pp. 37-38. Cashman, K.V. and Fiske, R.S., 1991. Fallout of pyroclastic debris from submarine volcanic eruptions. Science, 253:275-280. Cattell, A. and Arndt, N.T., 1987. Low- and high-alumina komatiites from a late Archean sequence, Newton Township, Ontario. Contrib. Mineral. Petrol., 97: 218-227. Catuneanu, O., 2001. Flexural partitioning of the Late Archaean Witwatersrand foreland system, South Africa. Sediment. Geol., 141/142:95-112.

792

References

Catuneanu, O., 2002. Sequence stratigraphy of clastic systems: concepts, merits, and pitfalls. J. Aft. Earth Sci., 35(1): 1--43. Catuneanu, O. and Biddulph, M.N., 2001. Sequence stratigraphy of the Vaal Reef facies associations in the Witwatersrand foredeep, South Africa. Sediment. Geol., 141/142:113-130. Catuneanu, O. and Elango, H.N., 2001. Tectonic control on fluvial styles: the Balfour Formation of the Karoo Basin, South Africa. Sediment. Geol., 140: 291-313. Catuneanu, O. and Eriksson, P.G., 1999. The sequence stratigraphic concept and the Precambrian rock record: an example from the 2.7-2.1 Ga Transvaal Supergroup, Kaapvaal craton. Precambrian Res., 97: 215-251. Catuneanu, O. and Eriksson, P.G., 2002. Sequence stratigraphy of the Precambrian RooihoogteTimeball Hill rift succession, Transvaal Basin, South Africa. Sediment. Geol., 147:71-88. Catuneanu, O., Willis, A.J. and Miall, A.D., 1998. Temporal significance of sequence boundaries. Sediment. Geol., 121: 157-178. Caudwell, C., 1987. l~tude exp6rimentale de la formation de micrite et de sparite dans le stromatolites d'eau douce h Rivularia. Bull. Soc. G6ol. France, 3(8): 299-306. Chacko, T., Creaser, R.A. and Poon, D., 1994. Spinel + quartz granites and associated metasedimentary enclaves from the Taltson magmatic zone. Alberta, Canada: a view into a root zone of a high temperature S-type granite batholith. Mineral. Mag. 8A: 161-162. Chacko, T., De, S.K., Creaser, R.A. and Muehlenbachs, K., 2000. Tectonic setting of the Talston magmatic zone at 1.9-2.0 Ga: a granitoid-based perspective. Can. J. Earth Sci., 37: 1597-1609. Chadwick, B., Vasudev, V.N. and Hegde, G.V., 2000. The Dharwar craton, southern India, interpreted as the result of late Archaean oblique convergence. Precambrian Res., 99:91-111. Chadwick, B., Claeys, P. and Simonson, B.M., 2001. New evidence for a large Palaeoproterozoic impact: spherules in a dolomite layer in the Ketilidian orogen, South Greenland. J. Geol. Soc. London, 158:331-340. Chafetz, H.S., 1994. Bacterially induced precipitates of calcium carbonate and lithification of microbial mats. In: W.E. Krumbein, D.M. Paterson and L.J. Stal (Eds.), Biostabilization of Sediments. BIS, Oldenburg, pp. 149-163. Chafetz, H.S. and Buczynski, C., 1992. Bacterially induced lithification of microbial mats. Palaios, 7: 277-293. Chakraborty, C. and Bose, P.K., 1990. Internal structures of sandwaves in a tide-storm interactive system: Proterozoic Lower Quartzite Formation, India. Sediment. Geol., 67: 133-142. : Chakraborty, C. and Bose, P.K., 1992. Rhythmic shelf storm beds: Proterozoic Kaimur Formation, India. Sediment. Geol., 77: 259-268. Chakraborty, T., 1999. Reconstruction of fluvial bars from the Proterozoic Manchaeral Quartzite, Pranhita-Godavari Valley, India. In: N.D. Smith and J. Rogers (Eds.), Fluvial Sedimentology IV. IAS Spec. Publ., 28. Blackwell, Oxford, pp. 451-466. Chakraborty, T. and Chakraborty, C., 2001. Eolian-aqueous interactions in the development of a Proterozoic sand-sheet: Shikaoda Formation, Hosangabad, India. J. Sediment. Res., 71 : 107-117. Chamberlain, C.P. and Sonder, L.J., 1990. Heat-producing elements and the thermal and baric patterns of metamorphic belts. Science, 250: 763-769. Chambers, J.E. and Wetherill, G.W., 1998. Making the terrestrial planets: N-body integrations of planetary embryos in three dimensions. Icarus, 136: 304-327. Chambers, L.A., Trudinger, P.A., Smith, J.W. and Bums, M.S., 1975. Fractionation of sulfur isotopes by continuous cultures of Desulfovibrio desulfuricans. Can. J. Microbiol., 21 : 1602-1607.

References

793

Champagne, C., Pilote, E and Mueller, W., 2002. Volcanologie physique et g6ochimie des komatiites arch6ennes: Groupe de Malartic, Abitibi, Qu6bec. Minist~re des Ressources naturelles, Qu6bec, MB 2002-04, 69 pp. (with map). Champion, D.C. and Smithies, R.H., 2000. The geochemistry of the Yule Granitoid Complex, East Pilbara granite-greenstone terrane; evidence for early felsic crust. Western Australia Geol. Survey, Annu. Rev., 1999-2000: 42-48. Champion, D.C. and Smithies, R.H., 2001. Archaean granites of the Yilgarn and Pilbara cratons, Western Australia. In: K.F. Cassidy, J.M. Dunphy and M.J. Van Kranendonk, (Eds.), Extended Abstract volume, 4th Intern. Archaean Symposium. Perth, 2001. AGSOmGeoscience Australia, Record 2001/37, pp. 134-136. Champion, D.C. and Smithies, R.H. (in preparation). The role of TTGs in crustal recycling during the Archaean: the less you see the more you had. Chan, M.A., Kvale, E.P., Archer, A.W. and Sonett, C.P., 1994. Oldest direct evidence of lunar-solar tidal forcing encoded in sedimentary rhythmites, Proterozoic Big Cottonwood Formation, central Utah. Geology, 22:791-794. Chandler, J.F., Reasenberg, R.D. and Shapiro, I.I., 1993. New bound on G. Bull. Amer. Astr. Soc., 25: 1233. Chapman, C.R. and Morrison, D., 1994. Impacts on the Earth by asteroids and comets: Assessing the hazard. Nature, 367: 33-40. Chapman, D.J. and Schopf, J.W., 1983. Biological and biochemical effects of the development of an aerobic environment. In: J.W. Schopf (Ed.), Earth's Earliest Biosphere: Its Origin and Evolution. Princeton University Press, Princeton, New Jersey, pp. 302-320. Chardon, D., Choukroune, P. and Jayananda, M., 1996. Strain patterns, d6collement and incipient sagducted greenstone terrains in the Archaean Dharwar craton (south India). J. Struct. Geol., 18: 991-1004. Chardon, D., Choukroune, P. and Jayananda, M., 1998. Sinking of the Dharwar Basin (South India): implications for Archaean tectonics. Precambrian Res., 91: 15-39. Chardon, D., Peucat, J.-J., Jayananda, M., Choukroune, P. and Fanning, C.M., 2002. Archean granite-greenstone tectonics at Kolar (South India): interplay of diapirism and bulk inhomogeneous contraction during juvenile magmatic accretion. Tectonics, 21 (3): 10.1029/2001TC901032. Chartrand, E and Cattalani, S., 1990. Massive sulfide deposits in northwestern Quebec. In: M. Rive, P. Verpaelst, Y. Gagnon, J.M. Lulin, G. Riverin and A. Simard (Eds.), La Ceinture polym6tallique du Nord-Ouest qu6b6cois. Can. Inst. Min. Metall. Spec. Vol., 43, pp. 77-91. Chauvel, C., Dupr6, B. and Jenner, G.A., 1985. The Sm-Nd age of Kambalda volcanics is 500 Ma too old! Earth Planet. Sci. Lett., 74:315-324. Chen, G.Q. and Ahrens, T.J., 1997. Erosion of terrestrial planet atmosphere by surface motion after a large impact. Phys. Earth. Planet. Interior, 100:21-26. Chen, Z., Liu, Y., Hodges, K.V., Burchfiel, B.C., Royden, L.H. and Deng, C., 1990. The Kangmar dome: a metamorphic core complex in southern Xizang (Tibet). Science, 250: 1552-1556. Cheney, E.S., 1996. Sequence stratigraphy and plate tectonic significance of the Transvaal succession of southern Africa and its equivalent in Western Australia. Precambrian Res., 79: 3-24. Cheney, E.S., Barton, J.M., Jr. and Brandl, G., 1990. Extent and age of the Soutpansberg sequences of southern Africa. S. Afr. J. Geol., 93: 664-675. Chetty, D. and Frimmel, H.E., 2000. The role of evaporites in the genesis of base metal sulphide mineralisation in the Northern Platform of the Pan-African Damara Belt, Namibia: Geochemical and fluid inclusion evidence from carbonate wall rock alteration. Mineral. Deposita, 35: 364-376.

794

References

Choukroune, E, Bouhallier, H. and Arndt, N.T., 1995. Soft lithosphere during periods of Archean crustal growth or crustal reworking. In: M.E Coward and A.C. Ries (Eds.), Early Precambrian Processes. Geol. Soc. London Spec. Publ., 95, pp. 67-86. Choukroune, P., Luddon, J., Calvert, A.J. and Bouhallier, H., 1997. Archean crustal growth and tectonic processes: a comparison of the Superior Province, Canada and the Dharwar Craton, India. In: J. Burg and M. Ford (Eds.), Orogeny through Time. Geol. Soc. London Spec. Publ., 121, pp. 63-98. Chown, E.H., Daigneault, R., Mueller, W. and Mortensen, J., 1992. Tectonic evolution of the northern volcanic zone, Abitibi belt, Quebec. Can. J. Earth Sci., 29:2211-2225. Chown, E.H., N'Dah, E. and Mueller, W.U., 2000. The relation between iron-formation and low temperature hydrothermal alteration in an Archean volcanic environment. Precambrian Res., 101: 263-275. Chown, E.H., Harrap, R. and Mouhksil, A., 2002. The role of granitic intrusions in the evolution of the Abitibi Belt, Canada. Precambrian Res., 115: 291-310. Christensen, U.R. and Yuen, D.A., 1985. Layered convection induced by phase transitions. J. Geophys. Res., 90:291-300. Christie-Blick, N., 1991. Onlap, offlap, and the origin of unconformity-bounded depositional sequences. Marine Geol., 97: 35-56. Christie-Blick, N., Grotzinger, J.P. and Von der Borch, C.C., 1988. Sequence stratigraphy in Proterozoic successions. Geology, 16:100-104. Chung, S.-L., Wang, K.-L., Crawford, A.J., Kamenetsky, V.S., Chen, C.-H., Lan, C.-Y. and Chen, C.-H., 2001. High-Mg potassic rocks from Taiwan: implications for the genesis of orogenic potassic lavas. Lithos, 59: 153-170. Cisne, J.L., 1984. A basin model for massive banded iron-formation and its geophysical applications. J. Geol., 5: 471-489. Claesson, S., Bogdanova, S.V., Bibikova, E.V. and Gorbatschev, R., 2001. Isotopic evidence for Palaeoproterozoic accretion in the basement of the East European Craton. Tectonophysics, 339: 1-18. Clague, D.A., Davis, A.S., Bischoff, J.L., Dixon, J.E. and Geyer, R., 2000. Lava bubble-wall fragments formed by submarine hydrovolcanic explosions on Loihi Seamount and Kilauea Volcano. Bull. Volcanol., 61: 437--449. Clague, D.A., Davis, A.S. and Dixon, J.E., 2003. Submarine strombolian eruptions on the Gorda mid-ocean ridge. In: J. White, D. Clague and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph (in press). Clark, S., 2000. Life on Other Worlds and How to Find It. Springer-Verlag, Berlin, 255 pp. Clarke, G.L., Sun, S.-S. and White, R.W., 1995. Grenville-age belts and associated older terranes in Australia and Antarctica. Australian Geological Survey Organisation Journal of Australian Geology and Geophysics, 16: 25-39. Clauer, N., Ey, E and Gauthier-Lafaye, E, 1985. K-Ar dating of different rock types from the Cluff Lake uranium ore deposits (Saskatchewan, Canada). In: R. Lain6, D. Alonso and M. Svab (Eds.), The Carswell structure uranium deposits, Saskatchewan. Geol. Assoc. Can. Spec. Paper, 29, pp. 87-103. Claypool, G.E., Holser, W.T., Kaplan, I.R., Sakai, H. and Zak, I., 1980. The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretations. Chem. Geol., 28: 199-260. Clayton, D.D., 1982. Cosmic chemical memory: a new astronomy. Quart. J. Roy. Astron. Soc., 23: 174-212.

References

795

Clayton, D.D., 1986. Interstellar fossil 26Mg and other extinct nuclides by low-energy heavy cosmic rays in molecular clouds. Nature, 368: 222-224. Clemens, J.D., 1998. Observations on the origins and ascent mechanisms of granite magmas. J. Geol. Soc., 155: 843-851. Clemens, J.D. and Wall, V.J., 1981. Origin and crystallization of some peraluminous (S-type) granitic magmas. Can. Mineral., 19:111-131. Clemmensen, L.B. and Abrahamsen, K., 1983. Aeolian stratification and facies associations in desert sediments, Arran Basin (Permian), Scotland. Sedimentology, 30:311-339. Clemmenson, L.B. and Dam, G., 1993. Aeolian sand sheet deposits in the lower Cambrian Neksr Sandstone Formation, Bornholm, Denmark: sedimentary architecture and genesis. Sediment. Geol., 83: 71-85. Clemmey, H. and Badham, J.P.N., 1982. Oxygen in the Precambrian atmosphere: an evaluation of the geological evidence. Geology, 10:141-146. Clendenin, C.W., Charlesworth, E.G. and Maske, S., 1988. Tectonic style and mechanism of early Proterozoic successor basin development, Southern Africa. Tectonophysics, 156: 275-291. Clifford, S.M. and Parker, T.J., 2001. The evolution of the Martian hydrosphere: implications for the fate of a primordial ocean and the current state of the Northern Plains. Icarus, 154: 40-79. Clifton, H.E., 1984. Sedimentation units in stratified resedimented conglomerate: Paleocene submarine canyon fill, Point Lobos, California. In: E.H. Koster and R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Can. Soc. Petrol. Geol., Calgary, Memoir, 10, pp. 429--441. Cloos, M., 1993. Lithospheric buoyancy and collisional orogenesis, subduction of oceanic plateaus, continental margins, island arcs, spreading ridges, and seamounts. Geol. Soc. Amer. Bull., 105: 717-737. Cloud, P.E., 1965. Significance of the Gunflint (Precambrian) microflora. Science, 148: 27-45. Cloud, P.E., 1968. Atmospheric and hydrospheric evolution on the primitive Earth. Science, 160: 729-736. Cloud, P.E., 1972. A working model of the primitive Earth. Amer. J. Sci., 272: 537-548. Cloud, P.E., 1973. Paleoecological significance of banded iron-formation. Econ. Geol., 68: 1135-1143. Cloud, P.E., 1983. Banded iron-formation a gradualist's dilemma. In: A.F. Trendall and R.C. Morris (Eds.), Iron-formations: Facts and Problems. Elsevier, Amsterdam, pp. 401-416. Cloud, P.E., 1988. Oasis in space. Norton and Company, New York, 372 pp. Cloud, P.E., Jr. and Semikhatov, M.A., 1969. Proterozoic stromatolite zonation. Amer. J. Sci., 267: 1017-1061. Coffin, M.E and Eldholm, O., 1994. Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics, 32: 1-36. Coffin, M.F. and Eldholm, O., 2001. Large igneous provinces: progenitors of some ophiolites? In: R.E. Ernst and K.L. Buchan (Eds.), Mantle Plumes: Their Identification Through Time. Geol. Soc. Amer. Spec. Paper, 352, pp. 59-70. Cogley, J.G., 1984. Continental margins and the extent and number of continents. Rev. Geophys. Space Phys., 22: 101-122. Coleman, A.E, 1908. The Lower Huronian Ice Age. J. Geol., 16: 149-158. Coleman, R.G., 1977. Ophiolites. Springer-Verlag, New York, 229 pp. Collins, C.D.N., 1983. Crustal structure of the southern McArthur Basin, northern Australia, from deep seismic sounding. BMR Journal of Australian Geology and Geophysics, 8: 19-34. Collins, W.H., 1925. North shore of lake Huron. Geological Survey of Canada, Memoir 143, 160 pp.

796

References

Collins, W.J., 1989. Polydiapirism of the Archean Mount Edgar Batholith, Pilbara Block, Western Australia. Precambrian Res., 43: 47-62. Collins, W.J., 1993. Melting of Archaean sialic crust under high H20 conditions: genesis of 3300 Ma Na-rich granitoids in the Mount Edgar Batholith, Pilbara Block, Western Australia. Precambrian Res., 60: 151-174. Collins, W.J. and Teyssier, C., 1990. Mount Edgar Batholith and Wanawoona Syncline. In: S.E. Ho, J.E. Glover, J.S. Myers and J.R. Muhling (Eds.), Excursion Guidebook No. 5, Pilbara and Hamersley Basin. Third Intern. Archaean Symp., Perth, 1990, Univ. Western Australia, Geology Dept. and Univ. Extension, Publ. 21, pp. 36--45. Collins, W.J. and Van Kranendonk, M.J., 1999. Model for the development of kyanite during partial convective overturn of Archaean granite-greenstone terranes: the Pilbara Craton, Australia. J. Metamorphic Petrol., 17: 145-156. Collins, W.J., Van Kranendonk, M.J. and Teyssier, C., 1998. Partial convective overturn of Archaean crust in the east Pilbara craton, Western Australia: driving mechanisms and tectonic implications. J. Struct. Geol., 20: 1405-1424. Collinson, J.D., 1983. Sedimentology of unconformities within a fluvio-lacustrine sequence: Middle Proterozoic of eastern North Greenland. Sediment. Geol., 34: 145-166. Collinson, J.D. and Thompson, 1989. Sedimentary Structures, 2nd edn. Unwin Hyman, London, 207 pp. Colliston, W.P., Reimold, W.U. and Crozier, J., 1999. Pseudotachylitic breccia occurrences and geometry in the Vredefort Dome and Witwatersrand Basin, South Africa. Meteoritics Planet. Sci., 34: 28. Coltice, N., Albar~de, E and Gillet, P., 2000. 4~176 constraints on recycling continental crust into the mantle. Science, 288: 845-847. Compston, W. and Pigeon, R.T., 1986. Jack Hills, evidence of more very old detrital zircons in Western Australia. Nature, 321: 766-769. Condie, K.C., 1981. Archean Greenstone Belts. Elsevier, Amsterdam, 434 pp. Condie, K.C., 1989. Geochemical changes in basalts and andesites across the Archean-Proterozoic boundary: identification and significance. Lithos, 23:1-18. Condie, K.C., 1990. Growth and accretion of continental crust: inferences based on Laurentia. Chem. Geol., 83: 183-194. Condie, K.C. (Ed.), 1992a. Proterozoic Crustal Evolution. Elsevier, Amsterdam, 510 pp. Condie, K.C., 1992b. Proterozoic terranes and continental accretion in southwestern North America. In: K.C. Condie (Ed.), Proterozoic Crustal Evolution. Elsevier, Amsterdam, pp. 447-479. Condie, K.C., 1994a. Archean Crustal Evolution. Elsevier, Amsterdam, 420 pp. Condie, K.C., 1994b. Greenstones through time. In: K.C. Condie (Ed.), Archean Crustal Evolution, Chapter 3. Elsevier, Amsterdam, pp. 85-120. Condie, K.C., 1995. Episodic ages of greenstones: a key to mantle dynamics. Geophys. Res. Lett., 22:2215-2218. Condie, K.C., 1997a. Plate Tectonics and Crustal Evolution. Butterworth-Heinemann, Oxford, 282 pp. Condie, K.C., 1997b. Contrasting sources from upper and lower continental crust: the greenstone connection. J. Geol., 105: 729-736. Condie, K.C., 1998. Episodic continental growth and supercontinents: a mantle avalanche connection? Earth Planet. Sci. Lett., 163: 97-108. Condie, K.C., 2000. Episodic continental growth models: afterthoughts and extensions. Tectonophys., 322: 153-162.

References

797

Condie, K.C., 2001 a. Mantle Plumes and Their Record in Earth History. Cambridge University Press, Cambridge, 305 pp. Condie, K.C., 200lb. Continental growth during formation of Rodinia at 1.35-0.9 Ga. Gondwana Res., 4: 5-16. Condie, K.C., 2002a. Breakup of a Paleoproterozoic supercontinent. Gondwana Res., 5(1): 41-43. Condie, K.C., 2002b. Incompatible element ratios in basalts from mantle plumes: Tracking deep mantle sources and continental growth rates with time (submitted). Condie, K.C., 2002c. Continental growth during a 1.9-Ga superplume event. J. Geodynam., 34: 249-264. Condie, K.C., Des Marais, D.J. and Abbott, D., 2000. Geologic evidence for a mantle superplume event at 1.9 Ga. Geochem. Geophys. Geosystems, 1: Paper No. 2000GC000095. Condie, K.C., Des Marais, D.J. and Abbott, D., 2001. Precambrian superplumes and supercontinents: a record in black shales, carbon isotopes, and paleoclimates? Precambrian Res., 106/3-4: 239-260. Condon, D.J., Prave, A.R. and Benn, D.I., 2002. Neoproterozoic glacial-rainout intervals: observations and implications. Geology, 30: 35-38. Coney, P.J., 1980. Cordilleran metamorphic core complexes: An overview. In: M.D. Crittenden Jr., P.J. Coney and G.H. Davis (Eds.), Cordilleran Metamorphic Core Complexes. Geol. Soc. Amer. Memoir, 153, pp. 7-31. Constantinou, M.A., 2001. A petrographic and microstructural study of the best preserved, oldest known, metasedimentary rocks in the 3.8-3.7 Ga Isua greenstone belt, southwest Greenland. B.Sc. Honours Thesis (unpubl.), Carleton University, Ottawa, 57 pp. Cook, D.G. and MacLean, B.C., 1995. The intracratonic Paleoproterozoic Forward orogeny, and implications for regional correlations, Northwest Territories, Canada. Can. J. Earth Sci., 32: 1991-2008. Cooke, D.L. and Moorhouse, W.W., 1969. Timiskaming volcanism in the Kirkland Lake area, Ontario, Canada. Can. J. Earth Sci., 6:117-132. Coombs, D.S. and Landis, C.A., 1966. Pumice from the South Sandwich eruption of March 1962 reaches New Zealand. Nature, 209: 289-290. Corcoran, P.L., 2000. Recognizing distinct portions of seamounts using volcanic facies analysis: examples from the Archean Slave Province, NWT, Canada. Precambrian Res., 101:277-261. Corcoran, P.L. and Dostal, J., 2001. Development of an ancient back-arc basin overlying continental crust: the Archean Peltier Formation, Northwest Territories, Canada. J. Geol., 109: 329-348. Corcoran, P.L. and Mueller, W.U., 2002. The effects of weathering, sorting and source composition in Archean high-relief basins: examples from the Slave Province, Northwest Territories, Canada. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: Modem Approach to Ancient Depositional Systems. Blackwell, Oxford, pp. 183-212. Corcoran, P.L., Mueller, W.U. and Chown, E.H., 1998. Climatic and tectonic influences on fan deltas and wave- to tide-controlled shoreface deposits: evidence from the Archean Keskarrah Formation, Slave Province, Canada. Sediment. Geol., 120: 125-152. Corcoran, P.L., Mueller, W.U. and Padgham, W.A., 1999. Influence of tectonism and climate on lithofacies distribution and sandstone and conglomerate composition in the Archean Beaulieu Rapids formation, Northwest Territories, Canada. Precambrian Res., 94: 175-204. Corfu, E and Easton, R.M., 2000. U-Pb evidence for polymetamorphic history of Huronian rocks within the Grenville front tectonic zone east of Sudbury, Ontario, Canada. Chem. Geol., 172: 149-171.

798

References

Corkeron, M.L. and George, A.D., 2001. Glacial incursion on a Neoproterozoic carbonate platform in the Kimberley region, Australia. Geol. Soc. Amer. Bull., 113:1121-1132. Cornell, D.H., Krrner, A., Humphreys, H. and Griffin, G., 1990. Age and origin of the polymetamorphosed Copperton Formation, Namaqua-Natal Province, determined by single grain zircon Pb-Pb dating. S. Afr. J. Geol., 93: 709-716. Cornell, D.H., SchiJtte, S.S. and Eglington, B.L., 1996. The Ongeluk Basaltic Andesite Formation in Griqualand West, South Africa: submarine alteration in a 2222 Ma Proterozoic sea. Precambrian Res., 79: 101-124. Corner, B., 1994. Geological evolution of western Dronning Maud Land within a Gondwana framework. Geophysics sub-programme, Pretoria, Department of Environmental Affairs and Tourism Report, 21 pp. Corner, B., 2000. Crustal framework of Namibia derived from magnetic and gravity data. Communications of the Geological Survey of Namibia, 12: 13-19. Corrigan, D., MacHattie, T.G. and Chakungal, J., 2000. The nature of the Wathaman Batholith and its relationship to the Archean Peter Lake domain along the Reindeer Lake Transect, Saskatchewan. Current Research, Geological Survey of Canada, Paper, 2000-C13, p. 10 (CD-ROM). Cotter, E., 1978. The evolution of fluvial style, with special reference to the central Appalachian Paleozoic. In: A.D. Miall (Ed.), Fluvial Sedimentology. Can. Soc. Petrol. Geol., Calgary, Alberta, Memoir, 5, pp. 361-383. Courtillot, V., Jaupart, C., Manighetti, I., Tapponnier, P. and Besse, J., 1999. On causal links between flood basalts and continental breakup. Earth Planet. Sci. Lett., 166: 177-195. Cousens, B.L., 1999. Geochemistry and geologic setting of the Baker Lake Group, Angikuni Lake to Yathkyed Lake. Indian and Northern Affairs Canada, DIAND Geology Division, EGS, 1999-05, p. 59. Cousens, B.L., 2000. Geochemistry of the Archean Kam Group, Yellowknife greenstone belt, Slave Province. Can. J. Geol., 108:181-197. Cousens, B.L., Aspler, L.B., Chiarenzelli, J.R., Donaldson, J.A., Sandeman, H., Peterson, T.D. and le Cheminant, A.N., 2001. Enriched Archean lithospheric mantle beneath western Churchill province tapped during Paleoproterozoic orogenesis. Geology, 29: 827-830. Cousens, B.L., Spera, E and Dobson, P., 1993. Post-eruptive alteration of silicic ignimbrites and lavas, Gran Canaria, Canary Islands: Sr, Nd, Pb and O isotopic evidence. Geochim. Cosmochim. Acta, 57: 631-640. Cousineau, P.A., 1994. Subaqueous pyroclastic deposits in an Ordovician fore-arc basin: an example from the Saint-Victor Formation, Quebec Appalachians, Canada. J. Sediment. Res., A64: 867-880. Cowan, E.A., Cai, J., Powell, R.D., Seramur, K.C. and Spurgeon, V.L., 1998. Modem tidal rhythmites deposited in a deep-water estuary. Geo-Marine Lett., 18: 40--48. Craig, D.C. and Loughnan, EC., 1964. Chemical and mineralogical transformations accompanying the weathering of basic volcanic rocks of New South Wales. Australian J. Soil Sci., 2: 218-234. Crawford, A.J., Fallon, T.J. and Green, D.H., 1989. Classification, petrogenesis and tectonic setting of boninites. In: A.J. Crawford (Ed.), Boninites and Related Rocks. Unwin Hyman, London, pp. 1-49. Creer, K.M., 1965. An expanding Earth? Nature, 205: 539-544. Crossley, D.J. and Stevens, R.K., 1976. Expansion of the Earth due to a secular decrease in G - evidence from Mercury. Can. J. Earth Sci., 13: 1723-1725. Crowell, J.C., 1957. Origin of pebbly mudstones. Geol. Soc. Amer. Bull., 68: 993-1010.

References

799

Crowell, J.C., 1974. Sedimentation along the San Andreas Fault, California. In: R.H. Dott Jr. and R.H. Shaver (Eds.), Modern and Ancient Geosynclinal Sedimentation. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 19, pp. 292-303. Crowell, J.C., 1999. Pre-Mesozoic Ice Ages: Their bearing on understanding the climate system. Geol. Soc. Amer. Memoir, 192, 106 pp. Crowley, T.J., Hyde, W.T. and Short, D.A., 1989. Seasonal cycle variations on the supercontinent of Pangaea. Geology, 17: 457-460. Crowley, T.J., Hyde, W.T. and Peltier, W.R., 2001. CO2 levels required for deglaciation of a "NearSnowball" Earth. Geophys. Res. Lett., 28: 283-286. Crown, D.A. and Baloga, S.M., 1998. Pahoehoe toe dimensions, morphology, and branching relationships at Mauna Ulu, Kilauea volcano, Hawai'i. Bull. Volcanol., 61: 288-305. Culler, T.S., Becker, T.A., Muller, R.A. and Renne, P.R., 2000. Lunar impact history from 40Ar/39Ar dating of glass spherules. Science, 287: 1785-1788. Cullers, R.L., DiMarco, M., Lowe, D.R. and Stone, J., 1993. Geochemistry of a silicified, felsic volcaniclastic suite from the early Archaean Panorama Formation, Pilbara Block, Western Australia: an evaluation of depositional and post-depositional processes with special emphasis on the rare-earth elements. Precambrian Res., 60:99-116. Cumming, G.L. and Krstic, D., 1992. The age of unconformity-related uranium mineralization in the Athabasca Basin, northern Saskatchewan. Can. J. Earth Sci., 29: 1623-1639. Cumming, G.L., Krstic, D. and Wilson, J.A., 1987. Age of the Athabasca Group, northern Alberta. Program with Abstracts, Geol. Assoc. Can., Mineral. Assoc. Can., Can. Geophys. Union, Joint Annual Meeting, Vol. 12, p. 35. Cummings, G., Eckstrand, O. and Peredery, W., 1982. Geochronological interpretations of Pb isotope ratios in nickel sulfides of the Thompson Belt, Manitoba. Can. J. Earth Sci., 19: 2306-2324. Curray, J.R., 1964. Transgressions and regressions. In: R.L. Miller (Ed.), Papers in Marine Geology. MacMillan, New York, pp. 175-203. Dahle, K., Flesja, K., Talbot, M.R. and Dreyer, T., 1997. Correlation of fluvial deposits by the use of Sm-Nd isotope analysis and mapping of sedimentary architecture in the Escanilla Formation (Ainsa Basin, Spain) and the Statfjord Formation (Norwegian North Sea). Abstracts, Sixth International Conference on Fluvial Sedimentology, Cape Town, 1997, South Africa, p. 46. Daigneault, R. and Archambault, G., 1990. Les grands couloirs de d6formation de la sous-province de l'Abitibi. In: M. Rive, G. Riverin, A. Simard, J.M. Lulin and Y. Gagnon (Eds.), The Northwestern Quebec Polymetallic Belt: A summary of 60 years of mining exploration. Can. Inst. Mines Metal. Spec. Vol., 43, pp. 43-64. Daigneault, R., St.-Julien, P. and Allard, G.O., 1990. Tectonic evolution of the northeastern portion of the Archean Abitibi greenstone belt, Chibougamau area, Quebec. Can. J. Earth Sci., 27: 1714-1736. Daigneault, R., Mueller, W. and Chown, E.H., 2002. Oblique Archean subduction: accretion and exhumation of an oceanic arc during dextral transpression, Southern Volcanic Zone, Abitibi Subprovince Canada. Precambrian Res., 115:261-290. Dalgarno, C.R. and Johnson, J.E., 1965. The Holowilena ironstone, a Sturtian glacigene unit. Quarterly Notes, Geol. Surv. S. Aust., 13: 2-4. Dalrymple, R.W. and Makino, Y., 1989. Description and genesis of tidal bedding in the Cobequid Bay-Salmon River estuary, Bay of Fundy, Canada. In: A. Taira and E Masuda (Eds.), Sedimentary Facies in the Active Plate Margin. Terra, Tokyo, pp. 151-177.

800

References

Dalrymple, R.W., Knight, R.J. and Zaitlain, B.A., 1990. Dynamics and facies model of a macrotidal sand-bar complex, Cobequid Bay-Salmon River estuary (Bay of Fundy). Sedimentology, 37: 577-612. Dalton, R., 2002. Squaring up over ancient life. Nature, 417: 782-784. Daly, J.S., Balagansky, V.V., Timmerman, M.J., Whitehouse, M.J., de Jong, K., Guise, P., Bogdanova, S., Gorbatschev, R. and Bridgwater, D., 2001. Ion microprobe U-Pb zircon geochronology and isotopic evidence for a trans-crustal suture in the Lapland-Kola orogen, northern Fennoscandian Shield. Precambrian Res., 105:289-314. Dalziel, I.W.D., 1991. Pacific margins of Laurentia and east Antarctica-Australia as a conjugate rift pair: evidence and implications for an Eocambrian supercontinent. Geology, 19: 598-601. Dalziel, I.W.D., 1992. Antarctica: a tale of two supercontinents. Annu. Rev. Earth Planet. Sci., 20: 501-526. Dalziel, I.W.D., 1997. Neoproterozoic-Paleozoic geography and tectonics: review, hypothesis and environmental speculation. Geol. Soc. Amer. Bull., 109:16-42. Dalziel, I.W.D., Mosher, S. and Gahagan, L.M., 2000. Laurentia-Kalahari collision and the assembly of Rodinia. J. Geol., 108: 499-513. Danielson, A., M611er, P. and Dulski, P., 1992. The europium anomalies in banded iron formations and the thermal history of the oceanic crust. Chem. Geol., 97: 89-100. Dann, J.C., 2000. The 3.5 Ga Komati Formation, Barberton greenstone belt, South Africa, Part I: new maps and magmatic architecture. S. Afr. J. Geol., 103: 47-68. Dann, J.C., 2001. Vesicular komatiites, 3.5 Ga Komati Formation, Barberton greenstone belt, South Africa: inflation of submarine lavas and origin of spinifex zones. Bull. Volcanol., 63: 462-481. Davies, G.E, 1992a. On the emergence of plate tectonics. Geology, 20: 963-966. Davies, G.E, 1992b. Plates and plumes: dynamos of the Earth's mantle. Science, 257: 493-494. Davies, G.E, 1993. Conjectures on the thermal and tectonic evolution of the Earth. Lithos, 30: 281-289. Davies, G.E, 1995a. Penetration of plates and plumes through the mantle transition zone. Earth Planet. Sci. Lett., 133: 507-516. Davies, G.F., 1995b. Punctuated tectonic evolution of the earth. Earth Planet. Sci. Lett., 136: 363-379. Davis, D.W., 1992. Ages constraints on deposition and provenance of Archean sediments in the southern Abitibi and Pontiac Subprovinces from U-Pb analysis of detrital zircons. Lithoprobe, Univ. British Columbia, Canada, Rep. 25, pp. 147-150. Davis, D.W., 2002. U-Pb geochronology of Archean metasedimentary rocks in the Pontiac and Abitibi subprovinces, Qu6bec. Precambrian Res., 115:97-117. Davis, G.H., 1980. Structural characteristics of metamorphic core complexes, southern Arizona. In: M.D. Crittenden, Jr., P.J. Coney and G.H. Davis (Eds.), Cordilleran Metamorphic Core Complexes. Geol. Soc. Amer. Mem. 153, pp. 35-77. Davis, R.A., 1968. Algal stromatolites composed of quartz sandstone. J. Sediment. Petrol., 38: 953-955. Davis, W.J., Lacroix, S., Gariepy, C. and Machado, N., 2000. Geochronology and radiogenic isotope geochemistry of plutonic rocks from the central Abitibi subprovince; significance to the internal subdivision and plutono-tectonic evolution of the Abitibi belt. Can. J. Earth Sci., 37:117-133. Davy, R., 1983. A contribution on the chemical composition of the Precambrian iron-formations. In: A.E Trendall and R.C. Morris (Eds.), Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 325-343.

References

801

Dawes, ER., Smithies, R.H., Centofanti, J. and Podmore, D.C., 1995. Sunrise Hill unconformity: a newly discovered regional hiatus between Archaean granites and greenstones in the northeastern Pilbara Craton. Aust. J. Earth Sci., 42: 635-639. de Almeida, EEM., de Brito Neves, B.B. and Carneiro, C.D.M., 2000. The origin and evolution of the South American Platform. Earth-Science Rev., 50:77-111. de Beer, J.H. and Stettler, E.H., 1992. The deep structure of the Limpopo Belt from geophysical studies. Precambrian Res., 55:173-186. de Beer, C.H., Gresse, EG., Theron, J.N. and Almond, J.E., 2002. The Geology of the Calvinia Area. Explanation, Sheet 3118 Calvinia 1 : 250000 scale, Council for Geoscience, Pretoria, 92 pp. de Boer, EL., Oost, A.P. and Visser, M.J., 1989. The diurnal inequality of the tide as a parameter for recognizing tidal influences. J. Sediment. Petrol., 59:912-921. de Raaf, J.EM., Boersma, J.R. and van Gelder, A., 1977. Wave-generated structures and sequences from a shallow marine succession, Lower Carboniferous, County Cork, Ireland. Sedimentology, 24:451-483. de Ronde, C.E.J. and de Wit, M.J., 1994. Tectonic history of the Barberton Greenstone Belt, South Africa: 490 million years of Archaean evolution. Tectonics, 13: 983-1005. de Ronde, C.E.J., De Wit, M.J. and Spooner, E.T.C., 1994. Early Archaean (> 3.2 Ga) Fe-oxide-rich, hydrothermal discharge vents in the Barberton greenstone belt, South Africa. Geol. Soc. Amer. Bull., 106: 86-104. de Rosen-Spence, A.E, 1976. Stratigraphy, development and petrogenesis of the central Noranda volcanic pile, Noranda, Quebec. Ph.D. Thesis (unpubl.), Univ. Toronto, Canada, 166 pp. de Rosen-Spence, A.E, Provost, G., Dimroth, E., Gochnauer, K. and Owen, V., 1980. Archean subaqueous felsic flows, Rouyn-Noranda, Quebec, Canada and their Quaternary equivalents. Precambrian Res., 12: 43-78. de Souza Filho, C.R. and Drury, S.A., 1998. A Neoproterozoic supra-subduction terrane in northern Eritrea, NE Africa. J. Geol. Soc. London, 155:551-566. de Wit, M.J., 1982. Gliding and overthrust nappe tectonics in the Barberton greenstone belt. J. Struct. Geol., 4:117-136. de Wit, M.J., 1983. Notes on a preliminary 1:25000 geological map of the southern part of the Barberton Mountain Land. Geol. Soc. S. Afr., Johannesburg, Spec. Publ., 9, pp. 185-187. de Wit, M.J., 1991. Archaean greenstone belt tectonism and basin development: some insights from the Barberton and Pietersburg greenstone belts, Kaapvaal Craton, South Africa. J. Afr. Earth Sci., 13: 45-63. de Wit, M.J., 1998. On Archean granites, greenstones, cratons and tectonics: does the evidence demand a verdict? Precambrian Res., 91 : 181-226. de Wit, M.J. and Ashwal, L.D., 1997a. Preface: Convergence towards divergent models of greenstone belts. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Science Publications, Oxford, pp. ix-xvii. de Wit, M.J. and Ashwal, L.D. (Eds.), 1997b. Greenstone Belts. Oxford Science Publications, Oxford, 809 pp. de Wit, M.J. and Hart, R.A., 1993. Earth's earliest continental lithosphere, hydrothermal flux and crustal recycling. Lithos, 30: 309-335. de Wit, M.J. and Hynes, A., 1995. The onset of interaction between the hydrosphere and oceanic crust, and the origin of the first continental lithosphere. In: M.E Coward and A.C. Ries (Eds.), Early Precambrian Processes. Geol. Soc. London Spec. Publ., 95, pp. 1-9.

802

References

de Wit, M.J., Armstrong, R.A., Hart, R.J. and Wilson, A.H., 1987a. Felsic igneous rocks within the 3.3- to 3.5-Ga Barberton greenstone belt: high crustal level equivalents of the surrounding tonalite--trondhjemite terrain, emplaced during thrusting. Tectonics, 6: 529-549. de Wit, M.J., Hart, R.A. and Hart, R.J., 1987b. The Jamestown ophiolite complex, Barberton Mountain Land: a section through 3.5 Ga oceanic crust. J. Afr. Earth Sci., 5:681-730. de Wit, M.J., Roering, C., Hart, R.J., Armstrong, R.A., de Ronde, C.E.J., Green, R.W.E., Tredoux, M., Peberdy, E. and Hart, R.A., 1992. Formation of an Archaean continent, Nature, 357: 553-562. De, S.K., Chacko, T., Creaser, R.A. and Muehlenbachs, K., 2000. Geochemical and Nd-Pb-O isotope systematics of granites from the Taltson Magmatic Zone, NE Alberta: implications for Early Proterozoic tectonics in western Laurentia. Precambrian Res., 102:221-249. Deamer, D.W., 1993. Prebiotic conditions and the first living cells. In: J.H. Lipps (Ed.), Fossil Prokaryotes and Protists. Blackwell, Oxford, pp. 11-18. Decho, A., 1990. Microbial exopolymer secretions in ocean environments: their role(s) in foodwebs and marine processes. Oceanogr. Mar. Biol. Annu. Rev., 28: 73-154. Decker, EL., 1990. Style and mechanics of liquefaction-related deformation, lower Absaroka Volcanic Supergroup (Eocene), Wyoming. Geol. Soc. Amer. Spec. Paper, 240, 78 pp. Dehler, C.M., Elrick, M., Karlstrom, K.E., Smith, G.A., Crossey, L.J. and Timmons, J.M., 2001. Neoproterozoic Chuar Group (~ 800-742 Ma), Grand Canyon: a record of cyclic marine deposition during global cooling and supercontinent rifting. Sediment. Geol., 141/142: 465-499. Delsemme, A., 1998. Cosmic origin of the biosphere. In: A. Brack (Ed.), The Molecular Origins of Life. Cambridge Univ. Press, Cambridge, pp. 100-118. Delsemme, A.H., 1995. Cometary origin of the biosphere: a progress report. Adv. Space Res., 15(3): 49-57. Demarest, J.M. and Kraft, J.C., 1987. Stratigraphic record of Quaternary sea levels: implications for more ancient strata. In: D. Nummedal, O.H. Pilkey and J.D. Howard (Eds.), Sea Level Fluctuation and Coastal Evolution. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 41, pp. 223-239. Dempster, T.J., Rogers, G., Tanner, EW.G., Bluck, B.J., Muir, R.J., Redwood, S.D., Ireland, T.R. and Paterson, B.A., 2002. Timing of deposition, orogenesis and glaciation within the Dalradian rocks of Scotland: constraints from U-Pb zircon ages. J. Geol. Soc., London, 159: 83-94. DePaolo, D.J., 1980. Crustal growth and mantle evolution: inferences from models of element transport and Nd and Sr isotopes. Geochim. Cosmochim. Acta, 44:1185-1196. DePaolo, D.J., 1983. The mean life of continents: estimates of continent recycling rates from Nd and Hf isotopic data and implications for mantle structure. Geophys. Res. Lett., 10: 705-708. Derry, L.A. and Jacobsen, S.B., 1990. The chemical evolution of Precambrian seawater: evidence from REEs in banded iron formations. Geochim. Cosmochim. Acta, 54: 2965-2977. Derry, L.A., Brasier, M.D., Corfield, R.M., Rozanov, A.Yu. and Zhuravlev, A.Yu., 1994. Sr and C isotopes in Lower Cambrian carbonates from the Siberian craton: a paleoenvironmental record during the "Cambrian explosion". Earth Planet. Sci. Lett., 128: 671-681. Derry, L.A., Kaufman, A.J. and Jacobsen, S.B., 1992. Sedimentary cycling and environmental change in the late Proterozoic; evidence from stable and radiogenic isotopes. Geochim. Cosmochim. Acta, 56: 1317-1329. Des Marais, D.J., 1994a. Tectonic control of the crustal organic carbon reservoir during the Precambrian. Chem. Geol., 114:303-314. Des Marais, D.J., 1994b. The Archean atmosphere: its composition and fate. In: K.C. Condie (Ed.), Archean Crustal Evolution. Developments in Precambrian Geology, 11. Elsevier, Amsterdam, pp. 505-519.

References

803

Des Marais, D.J., 1997. Isotopic evolution of the biogeochemical cycle during the Proterozoic Eon. Organic Geochem., 27: 185-193. Des Marais, D.J., Strauss, H., Summons, R.E. and Hayes, J.M., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature, 359: 605-609. Desrochers, J.-P. and Hubert, C., 1996. Structural evolution and early accretion of the Archean Malartic Composition Block, southern Abitibi Greenstone Belt, Quebec, Canada. Can. J. Earth Sci., 33: 1556-1569. Desrochers, J.-P., Hubert, C., Ludden, J.N. and Pilote, P., 1993. Accretion of Archean oceanic plateau fragments in the Abitibi greenstone belt, Canada. Geology, 21:451-454. Detmers, J., Brtichert, V., Habicht, K.S. and Kuever, J., 2001. Diversity of sulfur isotope fractionations by sulfate-reducing prokaryotes. App. Environ. Microbio., 67: 888-894. Deubner, E-L., 1990. Discussion on Late Precambrian tidal rhythmites in South Australia and the history of the Earth's rotation. J. Geol. Soc. London, 147: 1083-1084. Devaney, J.R., 1987. Sedimentology and stratigraphy of the northern and central metasedimentary belts in the Beardmore-Geraldton area of northern Ontario. M.Sc. Thesis (unpubl.), Lakehead University, Ontario, 227 pp. Devaney, J.R., 1991. Archean tidalites and shallow marine sandstones: examples from the Beardmore-Geraldton area, Wabigoon-Quetico subprovincial transition in northern Ontario. Geol. Assoc. Can.-Mineral. Assoc. Can.-Soc. Econ. Geol., Joint Annual Meeting, Program with Abstracts, Vol. 16, p. A30. Devaney, J.R., 1999a. Stages of volcanism, sedimentation, tectonism, and mineralization in the evolution of the Sioux Lookout orogenic belt, western Wabigoon Subprovince. Ontario Geol. Surv., Summary of Field Work and Other Activities 1998, Misc. Paper, 169: 156-167. Devaney, J.R., 1999b. Precambrian geology of the Colpitts-Bury lakes area (eastern Melchett Lake greenstone belt). Ontario Geol. Surv., Open File Rep. 5983, 96 pp. Devaney, J.R., 1999c. Late orogenic history of the Birch-Uchi Archean greenstone belt: interpretation of clastic pull-apart basins marginal to an hypothesized indentor/promontory block. In: R.M. Harrap and H.H. Helmstaedt (Eds.), 1999 Western Superior Transect, Fifth Annual Workshop. Lithoprobe Report #70, Lithoprobe Secretariat, Univ. British Columbia, pp. 55-65. Devaney, J.R., 1999d. Some major problems in Archean sedimentary basin analysis. Abstracts, Geol. Assoc. Canada-Mineral. Assoc. Canada, Joint Annual Meeting, Vol. 24, pp. 30. Devaney, J.R., 2000a. Regional geology of the Sioux Lookout orogenic belt, western Wabigoon Subprovince: stages of volcanism, sedimentation, tectonism and mineralization. Ontario Geol. Surv., Open File Rep. 6017, 158 pp. Devaney, J.R., 2000b. A summary of late orogenic wrench tectonics: strike-slip faulting, transpression, transtension and pull-apart basins, and application to greenstone belts and lode gold deposits of Superior Province. In: R.M. Harrap and H.H. Helmstaedt (Eds.), 2000 Western Superior Transect, Sixth Annual Workshop. Lithoprobe Report #77, Lithoprobe Secretariat, Univ. British Columbia, pp. 45-50. Devaney, J.R., 2001a. Stratigraphy of epiclastic and volcaniclastic facies units, northern Birch-Uchi greenstone belt, Uchi Subprovince. Ontario Geol. Surv., Open File Rep. 6030, 49 pp. Devaney, J.R., 200lb. Sedimentology and volcanology of selected tectonized stratigraphic units, southern Birch-Uchi greenstone belt, Uchi Subprovince. Ontario Geol. Surv., Open File Rep. 6031, 47 pp. Devaney, J.R. and Williams, H.R., 1989. Evolution of an Archean subprovince boundary: a sedimentological and structural study of part of the Wabigoon-Quetico boundary in northern Ontario. Can. J. Earth Sci., 26: 1013-1026.

804

References

Dewey, J.E, 1988. Extensional collapse of orogens. Tectonics, 7:1123-1139. Dewey, J.E and Windley, B.E, 1981. Growth and differentiation of continental crust. Philos. Trans. Roy. Soc. London Ser. A, 301:189-206. Deynoux, M., 1982. Periglacial polygonal structures and sand wedges in the Late Precambrian glacial formations of the Taoudeni Basin in Adrar of Mauretania (West Africa). Palaeogeogr., Palaeoclimatol., Palaeoecol., 39: 55-70. Deynoux, M., Duringer, P., Khatib, R. and Villeneuve, M., 1993. Laterally and vertically accreted tidal deposits in the Upper Proterozoic Madina-Kouta Basin, Southeastern Senegal, West Africa. Sediment. Geol., 84:179-188. Dickey, J.O., Bender, P.L., Failer, J.E., Newhall, X.X., Ricklefs, R.L., Ries, J.G., Shelus, P.J., Veillet, C., Whipple, A.L., Wiant, J.R., Williams, J.G. and Yoder, C.E, 1994. Lunar laser ranging: a continuing legacy of the Apollo program. Science, 265: 482-490. Dickinson, W.R., Klute, M.A., Hayes, M.J., Janecke, S.U., Lundin, E.R., McKittrick, M.A. and Olivares, M.D., 1988. Paleogeographic and paleotectonic setting of Laramide sedimentary basins in the Rocky Mountain region. Geol. Soc. Amer. Bull., 100: 1023-1039. Dilek, Y., Thy, P. and Moores, E., 1991. Episodic dike intrusions in the northwestern Sierra Nevada, California--Implications for multistage evolution of a Jurassic arc terrane. Geology, 19:180-184. Dillard, K.A., Simpson, E.L., Noto, R.C. and Wizevich, M., 1999. Characterization of fluvial deposits interbedded with flood basalts, Neoproterozoic Catochin Formation, Central Appalachians, USA. Precambrian Res., 97:115-134. DiMarco, M.J. and Lowe, D.R., 1989. Stratigraphy and sedimentology of an Early Archean felsic volcanic sequence, Eastern Pilbara Block, Western Australia, with special reference to the Duffer Formation and implications for crustal evolution. Precambrian Res., 44: 147-169. Dimroth, E., 1975. Paleo-environment of iron-rich sedimentary rocks. Geol. Rundschau, 64: 751-767. Dimroth, E., 1976. Aspects of the sedimentary petrology of cherty iron-formation. In: K.H. Wolf (Ed.), Handbook of Strata-bound and Stratiform Ore Deposits, 7. Elsevier, Amsterdam, pp. 203-254. Dimroth, E., 1982. The oldest rocks on Earth: stratigraphy and sedimentology of the 3.8 billion years old Isua supracrustal sequence. In: A.V. Sidorenko (Ed.), Sedimentary Geology of Highly Metamorphosed Precambrian Complexes. Nauka Publishing House, Moscow, pp. 16-27. Dimroth, E., 1985. A mass balance between Archean and Phanerozoic rates of magma emplacement, crustal growth and erosion: implications for recycling of the continental crust. Chem. Geol., 53: 17-24. Dimroth, E. and Kimberley, M.M., 1976. Precambrian atmospheric oxygen: evidence in the sedimentary distributions of carbon, sulfur, uranium, and iron. Canadian J. Earth Sci., 13:1161-1185. Dimroth, E. and Lichtblau, A.P., 1978. Oxygen in the Archean ocean; comparison of ferric oxide crusts on Archean and Cainozoic pillow basalts. Neues. Jahrbuch Min. Abhndl., 133: 1-22. Dimroth, E. and Rocheleau, M., 1979. Volcanology and sedimentology of the Rouyn-Noranda area, Quebec. Geol. Assoc. Canada, Field Trip A-1, Guide Book, Universit6 Laval, Quebec, 193 pp. Dimroth, E., Cousineau, P., Leduc, M. and Sanschagrin, Y., 1978. Structure and organization of Archean subaqueous basalt flows, Rouyn-Noranda area, Quebec, Canada. Can. J. Earth Sci., 15: 902-918. Dimroth, E., Imreh, L., Rocheleau, M. and Goulet, N., 1982. Evolution of the south-central part of the Archean Abitibi Belt, Quebec. Part I: Stratigraphy and paleogeographic model. Can. J. Earth Sci., 19:1729-1758.

References

805

Dimroth, E., Imreh, L., Goulet, N. and Rocheleau, M., 1983. Evolution of the south-central segment of the Archean Abitibi belt, Quebec. Part II: Tectonic evolution and geomechanical model. Can. J. Earth Sci., 20: 1355-1373. Dimroth, E., lmreh, L., Cousineau, P., Leduc, M. and Sanschagrin, Y., 1985. Paleogeographic analysis of mafic submarine flows and its use in the exploration for massive sulfide deposits. In: L.D. Ayres, P.C. Thurston, K.D. Card and W. Weber (Eds.), Evolution of Archean Supracrustal Sequences. Geol. Assoc. Can. Spec. Paper, 28, pp. 203-222. Dirks, P.H.G.M. and Jelsma, H.A., 1998. Horizontal accretion leading to stabilization on the Archean Zimbabwe Craton. Geology, 26:11-14. Dirks, P.H.G.M. and Jelsma, H.A., 2002. Crust-mantle decoupling and the growth of the Archaean Zimbabwe craton. J. Afr. Earth Sci., 34: 157-166. Dirks, P.H.G.M. and Wilson, C.J.L., 1995. Crustal evolution of the East Antarctica mobile belt in Prydz Bay: continental collision at 500 Ma? Precambrian Res., 75:189-207. Dixon, E.T., Honda, M., McDougall, I., Campbell, I.H. and Sigurdsson, I., 2000. Preservation of near-solar neon isotopic ratios in Icelandic basalts. Earth Planet. Sci. Lett., 180: 309-324. Dixon, J.M. and Summers, J.M., 1983. Patterns of total and incremental strain in subsiding troughs: experimental centrifuged models of inter-diapir synclines. Can. J. Earth Sci., 20:1843-1861. Doin, M.P. and Henry, P., 2001. Subduction initiation and continental crust recycling: the roles of rheology and eclogitization. Tectonophysics, 342:163-191. Dolozi, M.B., 1988. Physical volcanology of the Early Proterozoic Bear Lake mafic metavolcanic succession, Flin Flon, Manitoba. Ph.D. Thesis (unpubl.), Univ. Manitoba, 441 pp. Dolozi, M.B. and Ayres, L.D., 1991. Early Proterozoic, basaltic andesite tuff-breccia: downslope, subaqueous mass transport of phreatomagmatically-generated tephra. Bull. Volcanol., 53: 477-495. Dominguez, J.M.L. and Wanless, H.R., 1991. Facies architecture of a falling sea-level strandplain, Doce River coast, Brazil. In: D.J.P. Swift, G.E Oertel, R.W. Tillman and J.A. Thorne (Eds.), Shelf Sand and Sandstone Bodies: Geometry, Facies and Sequence Stratigraphy. IAS Spec. Publ., 14. Blackwell, Oxford, pp. 259-281. Donaldson, C.H., 1974. Olivine crystal types in harrisitic rocks of the Rhum pluton and Archean spinifex rocks. Geol. Soc. Amer. Bull., 85: 1721-1726. Donaldson, C.H., 1976. An experimental study of olivine morphology. Contrib. Mineral. Petrol., 57: 187-213. Donaldson, C.H., 1982. Spinifex-textured komatiites: A review of textures, mineral compositions, and layering. In: N.T. Arndt and E.G. Nisbet (Eds.), Komatiites. George Allen and Unwin, London, pp. 211-244. Donaldson, J.A. and de Kemp, E.A., 1998. Archaean quartz arenites in the Canadian Shield: examples from the Superior and Churchill Provinces. Sediment. Geol., 120: 153-176. Donaldson, J.A., Eriksson, P.G. and Altermann, W., 2002. Actualistic versus non-actualistic conditions in the Precambrian: a reappraisal of an enduring discussion. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modem Approach to Ancient Depositional Systems. IAS Spec. Publ. 33. Blackwell, Oxford, pp. 3-13. Dones, L. and Tremaine, S., 1993. Why does the Earth spin forward? Science, 259: 350-354. Donnelly, T.H. and Jackson, M.J., 1988. Sedimentology and geochemistry of a mid-Proterozoic lacustrine unit from northern Australia. Sediment. Geol., 58: 145-169. Donovan, A.D., 2001. Free market theory and sequence stratigraphy. Abstract, Amer. Assoc. Petrol. Geol. Hedberg Research Conference on "Sequence Stratigraphic and Allostratigraphic Principles and Concepts", Dallas, August 26-29, 2001, p. 22.

806

References

Dorr, J.V., 1973. Iron-formation in South America. Econ. Geol., 68: 1005-1022. Dorr, J.V.N., II, 1969. Physiographic, stratigraphic and structural development of the Quadril~itero Ferrffero, Minas Gerais, Brazil. U.S. Geol. Surv. Prof. Paper, 64 I-A, 110 pp. Dostal, J. and Mueller, W., 1992. Archean shoshonites from the Abitibi greenstone belt, Chibougamau (Qu6bec, Canada): geochemistry and tectonic setting. J. Volcanol. Geotherm. Res., 53: 145-165. Dostal, J. and Mueller, W., 1996. An Archean oceanic felsic dyke swarm in a nascent arc: the Hunter Mine Group, Abitibi greenstone belt. J. Volcanol. Geotherm. Res., 72: 37-57. Dostal, J. and Mueller, W., 1997. Komatiite flooding of a rifted Archean rhyolitic arc complex: geochemical signature and tectonic significance of the Stoughton-Roquemaure Group, Abitibi greenstone belt, Canada. J. Geol., 105: 545-563. Dostal, J. and Mueller, W.U., 2002. Archean hyaloclastites: fragmentation process and composition of the 2.72 Ga Stoughton-Roquemaure komatiites and komatiitic basalts. Abstract, AGUChapman conference, Explosive Subaqueous Volcanism, 21-25th January 2002, Dunedin, New Zealand, p 35. Dott, R.H., Jr., Byers, C.W., Fielder, G.W., Stenzel, S.R. and Winfree, K.E., 1986. Aeolian to marine transition in Cambro-Ordovician cratonic sheet sandstones of the northern Mississippi valley, U.S.A. Sedimentology, 33: 345-367. Doucet, P., Mueller, W. and Chartrand, E, 1994. Archean, deep-marine, volcanic eruptive products associated with the Coniagas massive sulfide deposit, Quebec, Canada. Can. J. Earth Sci., 31: 1569-1584. Doumbia, S., Pouclet, A., Kouamelan, A., Peucat, J.J., Vidal, M. and Delor, C., 1998. Petrogenesis of juvenile-type Birimian (Paleoproterozoic) granitoids in central C6te d'Ivoire, West Africa: geochemistry and geochronology. Precambrian Res., 87: 33-63. Draganic, I.G., Bjergbakke, E., Draganic, Z.D. and Sehested, K., 1991. Decomposition of ocean waters by potassium-40 radiation 3800 Ma ago as a source of oxygen and oxidizing species. Precambrian Res., 52: 321-336. Drake, M.J. and Righter, K., 2002. Determining the composition of the Earth. Nature, 416: 39-44. Dreibus, G. and W~inke, H., 2002. Comment on "Comparative geochemistry of Basalts from the Moon, Earth, HED asteroid and Mars: Implications for the origin of the Moon" by A. Ruzicka, G.A. Snyder and L.A. Taylor (2001), Geochim. Cosmochim. Acta, 65: 979-997. Geochim. Cosmochim. Acta, 66:2631-2632. Dressier, B.O. and Reimold, W.U., 2001. Terrestrial impact melt rocks and glasses. Earth-Sci. Rev., 56: 205-284. Drever, J.I., 1997. The Geochemistry of Natural Waters. Prentice-Hall, Englewood Cliffs, New Jersey, 436 pp. Drexel, J.E, Preiss, W.V. and Parker, A.J. (Eds.), 1993. Geol. Surv. S. Aust. Bull., 54(1), 242 pp. Dreybrodt, W., 1999. Chemical kinetics, speleothem growth and climate. Boreas, 28: 347-356. Driese, S.G., Mora, C.I., Cotter, E. and Foreman, J.L., 1992. Paleopedology and stable isotope chemistry of later Silurian vertic paleosols, Bloomsburg Formation, central Pennsylvania. J. Sediment. Petrol., 62: 825-841. Driese, S.G., Simpson, E.L. and Eriksson, K.A., 1995. Redoximorphic paleosols in alluvial and lacustrine deposits, 1.8 Ga Lochness Formation, Mount Isa, Australia: pedogenic processes and implications for paleoclimate. J. Sediment. Res., A65: 675-689. Drummond, C.N. and Wilkinson, B.H., 1996. Stratal thickness frequencies and the prevalence of orderedness in stratigraphic sequences. J. Geol., 104: 1-18.

References

807

Drummond, M.S. and Defant, M.J., 1990. A model for trondjhemite-tonalite-dacite genesis and crustal growth via slab melting: Archean to Modern comparisons. J. Geophys. Res., 95: 21,503-21,522. Drury, S.A., 1977. Structures induced by granite diapirs in the Archaean greenstone belt at Yellowknife, Canada: implications for Archaean geotectonics. J. Geol., 85: 345-358. Duchac, K.C. and Hanor, J.S., 1987. Origin and timing of the metasomatic silicification of an early Archaean komatiite sequence, Barberton Mountain Land, South Africa. Precambrian Res., 37: 125-146. Ducharme, Y., Stevenson, R.K. and Machado, N., 1997. Sm-Nd geochemistry and U-Pb geochronology of the Preissac and Lamotte leucogranites, Abitibi Subprovince. Can. J. Earth Sci., 34: 1059-1071. Duchesne, J.C., Vander Auwera, J. and Clifford, T., 2001. Geochemical constraints on the petrogenesis of the O'Kiep intrusive rocks (Namaqualand, South Africa): a lower crustal mafic source. Abstracts, EUG XI Conference, Strasbourg, 8-12 April 2001, p. 773. Dunbar, G.J. and McCall, G.J.H., 1971. Archaean turbidites and banded ironstones of the Mt. Belches area (Western Australia). Sediment. Geol., 5: 93-133. Duncan, C.C. and Turcotte, D.L., 1994. On the breakup and coalescence of continents. Geology, 22: 103-106. Dunlop, J.S.R., 1978. Shallow water sedimentation at North Pole, Pilbara, Western Australia. In: J.E. Glover and D.I. Groves (Eds.), Archean Cherty Metasediments--Their sedimentology, Micropalaeontology, Biochemistry and significance to mineralization. Geol. Dept. and Extension Service, Univ. Western Australia, Publ. 2, pp. 30-38. Dunlop, J.S.R., Muir, M.D., Milne, V.A. and Groves, D.I., 1978. A new microfossil assemblage from the Archaean of Western Australia. Nature, 274: 676-678. Durrheim, R.J. and Mooney, W.D., 1994. Evolution of the Precambrian Lithosphere: seismological and geochemical constraints. J. Geophys. Res., 99: 15,359-15,374. Duve de, C., 1991. Blueprint for a Cell: The Nature and Origin of Life. Neil Patterson Publ., Burlington, North Carolina, 296 pp. Dziggel, A., Stevens, G., Poujol, M., Anhaeusser, C.R. and Armstrong, R.A., 2002. Metamorphism of the granite-greenstone terrane south of the Barberton greenstone belt, South Africa: an insight into the tectono-thermal evolution of the "lower" portions of the Onverwacht Group. Precambrian Res., 114: 221-247. Edwards, M.B., 1979. Late Precambrian loessites from North Norway and Svalbard. J. Sediment. Petrol., 49: 85-92. Eggins, S.M., 1993. Origin and differentiation of picritic arc magmas, Ambae (Aoba), Vanuatu. Contrib. Mineral. Petrol., 114: 79-100. Eggleton, R.A., 1986. The relation between crystal structure and silicate weathering. In: S.M. Colman and D.E Dethier (Eds.), Rates of Chemical Weathering of Rocks and Minerals. Academic Press, Orlando, Florida, pp. 21-40. Eglington, B.M., Harmer, R.E. and Kerr, A., 1989. Isotope and geochemical constraints on Proterozoic crustal evolution in south-eastern Africa. Precambrian Res., 45: 159-174. Egyed, L., 1969. The slow expansion hypothesis. In: S.K. Runcorn (Ed.), The Application of Modern Physics to the Earth and Planetary Interiors. Wiley Interscience, London, pp. 65-75. Ehrenfreund, E and Menten, K.M., 2002. From molecular clouds to the origin of life. In: G. Horneck and C. Baumstark-Khan (Eds.), Astrobiology: The Quest for the Conditions of Life. Springer, Berlin, pp. 7-24.

808

References

Eide, E.A. and Torsvik, T.H., 1996. Paleozoic supercontinental assembly, mantle flushing and genesis of the Kiaman superchron. Earth Planet. Sci. Lett., 144: 389-402. Einsele, G., 1992. Sedimentary Basins: Evolution, Facies and Sediment Budget. Springer, Berlin, 628 pp. Eisbacher, G.H., 1985. Late Proterozoic rifting, glacial sedimentation, and sedimentary cycles in the light of Windermere deposition, Western Canada. Palaeogeog. Palaeoclim. Palaeoecol., 51: 231-254. Elliott, T., 1986. Siliciclastic shorelines. In: H.G. Reading (Ed.), Sedimentary Environments and Facies. Blackwell, Oxford, pp. 155-188. Els, B.G., 1990. Determination of some palaeohydraulic parameters for a fluvial Witwatersrand succession. S. Afr., J. Geol., 93:531-537. Els, B.G., 1998. The auriferous Late Archaean sedimentation systems of South Africa: unique palaeoenvironmental conditions? Sediment. Geol., 120: 205-224. Embleton, B.J.J. and Williams, G.E., 1986. Low palaeolatitudes of deposition for late Precambrian periglacial varvites in South Australia: implications for palaeoclimatology. Earth Planet. Sci. Lett., 79: 419-430. Embry, A.F., 1993. Transgressive-regressive (T-R) sequence analysis of the Jurassic succession of the Sverdrup Basin, Canadian Arctic Archipelago. Can. J. Earth Sci., 30: 301-320. Embry, A.F., 1995. Sequence boundaries and sequence hierarchies: problems and proposals. In: R.J. Steel, V.L. Felt, E.P. Johannessen and C. Mathieu (Eds.), Sequence stratigraphy on the northwest European Margin. Norwegian Petroleum Society (NPF) Spec. Publ., 5. Elsevier, Amsterdam, pp. 1-11. Embry, A.F. and Catuneanu, O., 2002. Practical Sequence Stratigraphy: Concepts and Applications. Can. Soc. Petrol. Geol., Calgary, short course notes, 167 pp. Embry, A.F. and Johannessen, E.P., 1992. T-R sequence stratigraphy, facies analysis and reservoir distribution in the uppermost Triassic-Lower Jurassic succession, western Sverdrup Basin, Arctic Canada. In: T.O. Vorren, E. Bergsager, O.A. Dahl-Stamnes, E. Holter, B. Johansen, E. Lie and T.B. Lund (Eds.), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF) Spec. Publ., 2. Elsevier, Amsterdam, pp. 121-146. Emery, D. and Myers, K., 1996. Sequence Stratigraphy. Blackwell, Oxford, 297 pp. Emery, K.O. and Uchupi, E., 1972. Western North Atlantic Ocean: topography, rocks, structure, water, life and sediments. AAPG Mem., 17, 532 pp. Endress, M., Zinner, E. and Bischoff, A., 1996. Early aqueous activity on primitive meteorite parent bodies. Nature, 379:701-703. Engel, A.E.J., Dixon, T.H. and Stern, R.J., 1980. Late Precambrian evolution of Afro-Arabian crust from ocean arc to craton. Geol. Soc. Amer. Bull., 91: 699-706. Engelbrecht, C.J., Baumbach, G.W.S., Matthysen, J.L. and Fletcher, P., 1986. The West Wits line. In: C.R. Anhaeusser and S. Maske (Eds.), Mineral Deposits of Southern Africa. Geol. Soc. S. Afr., Johannesburg, pp. 599-649. England, G., Rasmussen, B., Krapez, B. and Groves, D.I., 2001. The origin of uraninite, bitumen nodules, and carbon seams in Witwatersrand gold-uranium-pyrite ore deposits, based on a PermoTriassic analogue. Econ. Geol., 96: 1907-1920. England, G., Rasmussen, B., Krapez, B. and Groves, D.I., 2002. Palaeoenvironmental significance of rounded pyrite in siliciclastic sequences of the Late Archean Witwatersrand basin: oxygendeficient atmosphere or hydrothermal alteration? Sedimentology, 49:1133-1156.

References

809

England, P.C. and Thompson, A.B., 1984. Pressure-temperature-time paths of regional metamorphism. 1. Heat transfer during the evolution of regions of thickened continental crust. J. Petrol., 25: 894-928. Eriksson, K.A., 1977. Tidal deposits from the Archaean Moodies Group, Barberton Mountain Land, South Africa. Sediment. Geol., 18:257-281. Eriksson, K.A., 1978. Alluvial and destructive beach facies from the Archaean Moodies Group, Barberton Mountain Land, South Africa and Swaziland. In: A.D. Miall (Ed.), Fluvial Sedimentology. Can. Soc. Petrol. Geol., Calgary, Mem., 5, pp. 287-311. Eriksson, K.A., 1979. Marginal marine depositional processes from the Archaean Moodies Group, Barberton Mountain Land, South Africa: Evidence and significance. Precambrian Res., 8: 153-182. Eriksson, K.A., 1980. Transitional sedimentation styles in the Fig Tree and Moodies Group, Barberton Mountain Land, South Africa: evidence favouring an Atlantic or Japan sea-type Archean continental margin. Precambrian Res., 12:141-160. Eriksson, K.A., 1982. Geometry and internal characteristics of Archean submarine channel deposits, Pilbara Block, Western Australia. J. Sediment. Petrol., 52: 383-393. Eriksson, K.A., 1983. Siliciclastic-hosted iron-formations in the early Archean Barberton and Pilbara sequences. J. Geol. Soc. Aust., 30: 473-482. Eriksson, K.A., 1995. Crustal growth, surface processes, and atmospheric evolution on the early Earth. In: M.P. Coward and A.C. Ries (Eds.), Early Precambrian Processes. Geol. Soc. London Spec. Publ., 95, pp. 11-25. Eriksson, K.A. and Fedo, C.M., 1994. Archean Cratonic Sedimentary Successions. In: K.C. Condie (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 171-204. Eriksson, K.A. and Simpson, E.L., 1990. Recognition of high-frequency sea-level changes in Proterozoic siliciclastic tidal deposits, Mount Isa, Australia. Geology, 18: 474--477. Eriksson, K.A. and Simpson, E.L., 1993. Siliciclastic braided-alluvial sediments intercalated within continental flood basalts in the Early to Middle Proterozoic Mount Isa Inlier, Australia. In: M. Marzo and C. Puigdeffibreagas (Eds.), Alluvial Sedimentation. IAS Spec. Publ., 17. Blackwell, Oxford, pp. 167-184. Eriksson, K.A. and Simpson, E.L., 1998. Controls on spatial and temporal distribution of Precambrian eolianites. Sediment Geol., 120: 275-294. Eriksson, K.A. and Simpson, E.L., 2000. Quantifying the oldest tidal record: the 3.2 Ga Moodies Group, Barberton greenstone Belt, South Africa. Geology, 28:831-834. Eriksson, K.A. and Simpson, E.L., 2002. Precambrian tidalites: recognition and significance. Abstract, 16th Intern. Sedimentology Congress, Johannesburg, July 2002, p. 92. Eriksson, K.A., Krapez, B. and Fralick, P.W., 1994. Sedimentology of Archean greenstone belts: signatures of tectonic evolution. Earth Sci. Rev., 37: 1-88. Eriksson, K.A., Krapez, B. and Fralick, P.W., 1997. Sedimentological aspects. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 33-54. Eriksson, K.A., Turner, B.R. and Vos, R.G., 1981. Evidence of tidal processes from the lower part of the Witwatersrand Supergroup, South Africa. Sediment. Geol., 29: 309-325. Eriksson, P.G., 1999. Sea level changes and the continental freeboard concept: general principles and application to the Precambrian. Precambrian Res., 97: 143-154. Eriksson, P.G. and Altermann, W., 1998. An overview of the geology of the Transvaal Supergroup dolomites (South Africa). Env. Geol., 36(1/2): 179-188.

810

References

Eriksson, EG. and Cheney, E.S., 1992. Evidence for the transition to an oxygen-rich atmosphere during the evolution of red beds in the Lower Proterozoic sequences of southern Africa. Precambrian Res., 54: 257-269. Eriksson, EG. and Reczko, B.EE, 1995. The sedimentary and tectonic setting of the Transvaal Supergroup floor rocks to the Bushveld complex. J. Afr. Earth Sci., 21: 487-504. Eriksson, EG. and Reczko, B.EE, 1998. Contourites associated with pelagic mudrocks and distal delta-fed turbidites in the Lower Proterozoic Timeball Hill Formation epeiric basin (Transvaal Supergroup), South Africa. Sediment. Geol., 120: 319-335. Eriksson, EG., Altermann, W., Catuneanu, O., van der Merwe, R. and Bumby, A.J., 2001c. Major influences on the evolution of the 2.67-2.1 Ga Transvaal basin, Kaapvaal craton. Sediment. Geol., 141-142: 205-231. Eriksson, EG., Altermann, W., Eberhardt, L., Ahrend-Heidbrinck, S. and Bumby, A.J., 2002a. Palaeoproterozoic epeiric sea palaeoenvironments: the Silverton Formation (Pretoria Group, Transvaal Supergroup), South Africa. In: W. Altermann and EL. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. IAS Spec. Publ., 33. Blackwell, Oxford, pp. 351-368. Eriksson, EG., Schreiber, U.M. and Van der Neut, M., 1991. A review of the sedimentology of the Early Proterozoic Pretoria Group, Transvaal Sequence, South Africa: implications for tectonic setting. J. Afr. Earth Sci., 13:107-119. Eriksson, E G., Schreiber, U.M., Van der Neut, M., Labuschagne, H., Van der Schyff, W. and Potgieter, G., 1993. Alternative marine and fluvial models for the non-fossiliferous quartzitic sandstones of the Early Proterozoic Daspoort Formation, Transvaal Sequence of southern Africa. J. Afr. Earth Sci., 16: 355-366. Eriksson, EG., Bose, EK. and Altermann, W., 1999a. Preface, special issue: variation in sea level and continental freeboard: evidence from the Precambrian volcano-sedimentary record. Precambrian Res., 97: 137-141. Eriksson, EG., Catuneanu, O., Aspler, L.B., Chiarenzelli, J.R. and Martins-Neto, M.A., 2001a. Preface, special issue: the influence of magmatism, tectonics, sea level change and palaeo-climate on Precambrian basin evolution: change over time. Sediment. Geol., 141/142: vii-xi. Eriksson, EG., Condie, K.C., Tirsgaard, H., Mueller, W.U., Altermann, W., Miall, A.D., Aspler, L.B., Catuneanu, O. and Chiarenzelli, J.R., 1998b. Precambrian clastic sedimentation systems. Sediment. Geol., 120: 5-53. Eriksson, EG., Condie, K.C., Van der Westhuizen, W., Van der Merwe, R., de Bruiyn, H., Nelson, D.R., Altermann, W., Catuneanu, O., Bumby, A.J., Lindsay, J. and Cunningham, M.J., 2002b. Late Archaean superplume events: a Kaapvaal-Pilbara perspective. J. Geodynam., 34: 207-247. Eriksson, E G., Martins-Neto, M.A., Nelson, D.R., Aspler, L.B., Chiarenzelli, J.R., Catuneanu, O., Sarkar, S., Altermann, W. and Rautenbach, C.J. de W., 200lb. An introduction to Precambrian basins: their characteristics and genesis. Sediment. Geol., 141-142: 1-35. Eriksson, EG., Mazumder, R., Sarkar, S., Bose, EK., Altermann, W. and Van der Merwe, R., 1999b. The 2.7-2.0 Ga volcano-sedimentary record of Africa, India and Australia: evidence for global and local changes in sea level and continental freeboard. Precambrian Res., 97: 269-302. Eriksson, EG., Mueller, W.U. and Tirsgaard, H., 1998a. Preface, special issue: Precambrian clastic sedimentation systems. Sediment. Geol., 120: 1-4. Eriksson, EG., Reczko, B.EF., Boshoff, A.J., Schreiber, U.M., Van der Neut, M. and Snyman, C.E, 1995. Architectural elements from Lower Proterozoic braid-delta and high-energy tidal flat de-

References

811

posits in the Magaliesberg Formation, Transvaal Supergroup, South Africa. Sediment. Geol., 97: 99-117. Eriksson, P.G., Reczko, B.EE and Piper, D.P., 1996. An interpretation of boron contents within a Palaeoproterozoic volcano-sedimentary succe,;sion: Pretoria Group, Transvaal Supergroup, South Africa. Precambrian Res., 78: 273-287. Eriksson, P.G., Simpson, E.L., Eriksson, K.A., Bumby, A.J., Steyn, G.L. and Sarkar, S., 2000. Muddy roll-up structures in siliciclastic interdune beds of the c. 1.8 Ga Waterberg Group, South Africa. Palaios, 15: 177-183. Ernst, R. and Buchan, K., 1997. Giant radiating dyke swarms: their use in identifying pre-Mesozoic large igneous provinces and mantle plumes. In: J. Mahoney and M. Coffin (Eds.), Large Igneous Provinces. American Geophysical Union Monograph Series, 100, pp. 297-333. Ernst, R.E. and Buchan, K.L., 2001 a. The use of mafic dike swarms in identifying and locating mantle plumes. In: R.E. Ernst and K.L. Buchan (Eds.), Mantle Plumes: their Identification through Time. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 352, pp. 247-265. Ernst, R.E. and Buchan, K.L., 2001 b. Large mafic magmatic events through time and links to mantleplume heads. In: R.E. Ernst and K.L. Buchan (Eds.), Mantle Plumes: Their Identification through Time. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 352, pp. 483-575. Ernst, R.E. and Buchan, K.L., 2002a. Maximum size and distribution in time and space of mantle plumes: evidence from large igneous provinces. J. Geodynam. 34: 309-342. Ernst, R.E. and Buchan, K.L., 2002b. Maximum size and distribution in time and space of mantle plumes: evidence from large igneous provinces. In: K.C. Condie, D. Abbott, D. Des Marais (Eds.), Superplumes in Time and Space, special issue, J. Geodynamics, 34: 309-342. (Erratum, 2002, 34: 711-714.) Ernst, R.E. and Buchan, K.L., 2003. Recognizing Mantle Plumes in the Geological Record. Annu. Rev. Earth Planet. Sci. (submitted). Esperanqa, S., Carlson, R.W., Shirey, S.B. and Smith, D., 1997. Dating crust-mantle separation: ReOs isotopic study of mafic xenoliths from central Arizona. Geology, 25:651-654. Etheridge, M.A., Rutland, R.W.R. and Wyborn, L.A., 1987. Orogenic and tectonic processes in the Early to Middle Proterozoic of northern Australia. In: A. Kroner, (Ed.), Precambrian Lithospheric Evolution. American Geophysical Union Geodynamics Series, 17, pp. 131-147. Ethridge, E G., Germanoski, D., Schumm, S.A. and Wood, L.J., 2001. The morphologic and stratigraphic effects of base-level change: a review of experimental studies. Abstracts, Seventh International Conference on Fluvial Sedimentology, Lincoln, August 6-10, 2001, p. 95. Ethridge, F,G., Tyler, N. and Burns, L.K., 1984. Sedimentology of a Precambrian quartz pebble conglomerate, southwest Colorado. In: E.H.Koster and R.J.Steel (Eds.), Sedimentology of Gravels and Conglomerate. Can. Soc. Petrol. Geol., Mem. 10, pp. 165-174. Evans, D.A., 1998. True polar wander, a supercontinental legacy. Earth Planet. Sci. Lett., 157: 1-8. Evans, D.A., Beukes, N.J. and Kirschvink, J.L., 1997. Low latitude glaciation in the Palaeoproterozoic era. Nature, 386: 262-266. Evans, D.A., Ripperdan, R.L. and Kirschvink, J.L., 1998. Polar wander and the Cambrian-reply. Science, 279, 9a. Evans, D.A.D., 2000. Stratigraphic, geochronological and paleomagnetic constraints upon the Neoproterozoic climatic paradox. Amer. J. Sci., 300: 347-433. Evans, D.M., Windrim, D.P. and Armstrong, R.A., 2002. Northern margin of the Namaqua-Natal mobile belt in the Karas Mountains, Namibia: Mapping and radiometric age constraints. Abstract, 1l th Quadrennial IAGOD Symposium and Geocongress 2002, Windhoek, 22-26 July 2002 (CD-ROM).

812

References

Evans, N.J., Shahinpoor, M. and Ahrens, T.J., 1994. Hypervelocity impact: ejecta velocity, angle and composition. In: B.O. Dressier, R.A.E Grieve and V.L. Sharpton (Eds.), Large Meteorite Impacts and Planetary Evolution. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 293, pp. 93-101. Ewers, W.E. and Morris, R.C., 1981. Studies on the Dales Gorge Member of the Brockman Iron Formation. Econ. Geol., 76: 1929-1953. Eyles, N., 1993. Earth's glacial record and its tectonic setting. Earth-Sci. Rev., 35: 1-248. Eyles, N. and Young, G.M., 1994. Geodynamic controls on glaciation in Earth history. In: M. Deynoux, J.M.G. Miller, E.W. Domack, N. Eyles, I.J. Fairchild and G.M. Young (Eds.), Earth's Glacial Record. Cambridge University Press, Cambridge, pp. 1-28. Eyles, C.H., Eyles, N. and Lagoe, M.B., 1991. The Yakataga Formation: a late Miocene to Pleistocene record of temperate glacial marine sedimentation in the Gulf of Alaska. In: J.B. Anderson and G.M. Ashley (Eds.), Glacial Marine Sedimentation: Paleoclimatic Signifcance. GSA Spec. Paper, 261, pp. 159-180. Fagerstrom, J.A., 1967. Development, flotation, and transportation of mud crusts--neglected factors in sedimentology. J. Sediment. Petrol., 37: 73-79. Fahrig, W.E, 1961. The geology of the Athabasca Formation. Geol. Surv. Can. Bull., 68: 1-41. Fahrig, W.E, 1987. The tectonic settings of continental mafic dyke swarms: failed arm and early passive margin. In: H.C. Halls and W.E Fahrig (Eds.), Mafic Dyke Swarms. Geol. Assoc. Can. Spec. Paper, 34, pp. 331-348. Fairchild, I.J., 1991. Origins of carbonate in Neoproterozoic stromatolites and the identification of modern analogues. Precambrian Res., 53: 281-299. Fairchild, I.J. and Hambrey, M.J., 1995. Vendian basin evolution in East Greenland and Northeast Svalbard. Precambrian Res., 73: 217-233. Fairchild, I.J. and Herrington, P.M., 1989. A tempestite-stromatolite-evaporite association (Late Vendian, East Greenland): a shoreface-lagoon model. Precambrian Res., 43: 101-127. Farmer, J.D., 2000. Hydrothermal systems: doorways to early biosphere evolution. GSA Today, 10(7): 1-9. Farquhar, J. and Wing, B.A., 2003. Multiple sulfur isotopes and the evolution of the atmosphere. Earth Planet. Sci. Lett., 213: 1-13. Farquhar, J., Bao, H. and Thiemens, M., 2000a. Atmospheric influence of Earth's earliest sulfur cycle. Science, 289: 756-758. Farquhar, J., Chacko, T. and Ellis, D.J., 1996. Preservation of oxygen isotope compositions in granulites from Northwestern Canada and Enderby Land, Antarctica: implications for hightemperature isotopic thermometry. Contrib. Miner. Petrol., 125:213-224. Farquhar, J., Savarino, J., Jackson, T.L. and Thiemens, M.H., 2000b. Evidence of atmospheric sulphur in the martian regolith from sulphur isotopes in meteorites. Nature, 404: 50-52. Farquhar, J., Savarino, J., Airieau, S. and Thiemens, M.H., 2001. Observation of wavelengthsensitive mass-independent sulfur isotope effects during SO2 photolysis: application to the early atmosphere. J. Geophys. Res. Planets, 106: 32,829-32,839. Farquhar, J., Wing, B.A., McKeegan, K.D., Harris, J.W., Cartigny, P. and Thiemens, M.H., 2002. Mass-independent sulfur of inclusions in diamond and sulfur recycling on early Earth. Science, 298: 2369-2372. Faure, E, 2001. Les textures de croissance rapide dans les roches magmatiques basiques et ultrabasiques: 6tude exp6rimentale et nanoscopique. Ph.D. Thesis (unpubl.), Universit6 Blaise Pascal, Clermont Ferrand, 206 pp. Faure, G., 2001. Origin of Igneous Rocks: The Isotopic Evidence. Springer-Verlag, New York, 496 pp.

References

813

Fedo, C.M., 2000. Setting and origin for problematic rocks from the > 3.7 Ga Isua Greenstone Belt, southern west Greenland: Earth's oldest coarse clastic sediments. Precambrian Res., 101: 69-78. Fedo, C.M. and Cooper, J.D., 2001. Sedimentology and sequence stratigraphy of Neoproterozoic and Cambrian units across a craton-margin hinge zone, southeastern California, and implications for the early evolution of the Cordilleran margin. Sediment. Geol., 141/142:501-522. Fedo, C.M. and Eriksson, K.A., 1996. Stratigraphic framework of the ~ 3.0 Ga Buhwa greenstone belt: a unique stable shelf succession in the Zimbabwe Archean craton. Precambrian Res., 77: 161-178. Fedo, C.M. and Whitehouse, M.J., 2002. Metasomatic origin of quartz-pyroxene rock, Akilia, Greenland, and implications for Earth's earliest life. Science, 296: 1448-1452. Fedo, C.M., Eriksson, K.A. and Krogstad, E.J., 1996. Geochemistry of shales from the Archean (,~ 3.0 Ga) Buhwa Greenstone Belt, Zimbabwe: implications for provenance and source-area weathering. Geochim. Cosmochim. Acta, 60: 1751-1763. Fedo, C.M., Myers, J.S. and Appel, P.W.U., 2001. Depositional Setting and paleogeographic implications of Earth's oldest supracrustal rocks, the > 3.7 Ga Isua greenstone belt, West Greenland. Sediment. Geol., 141/142: 61-77. Fedo, C.M., Nesbitt, H.W. and Young, G.M., 1995. Unraveling the effects of potassium metasomatism in sedimentary rocks and paleosols, with implications for paleoweathering conditions and provenance. Geology, 23:921-924. Fedo, C.M., Young, G.M. and Nesbitt, H.W., 1997. Paleoclimatic control on the composition of the Paleoproterozoic Serpent Formation, Huronian Supergroup, Canada: a greenhouse to icehouse transition: Precambrian Res., 86: 201-223. Fedonkin, M.A., 1996. Geobiological trends and events in the Precambrian Biosphere. In: O.H. Walliser (Ed.), Global Events and Event Stratigraphy. Springer-Verlag, Berlin, pp. 89-112. Feldstein, S.N. and Lange, R.A., 1999. Pliocene potassic magmas from the Kings River region, Sierra Nevada, California: evidence for melting of a subduction-modified mantle. J. Petrol., 40: 1301-1320. Feng, R. and Kerrich, R., 1991. Single zircon age constraints on the tectonic juxtaposition of the Archean Abitibi greenstone belt and Pontiac subprovince, Quebec, Canada. Geochim. Cosmochim. Acta, 55: 3437-3441. Feng, R. and Kerrich, R., 1992. Geochemical evolution of granitoids from the Archaean Abitibi Southern Volcanic Zone and Pontiac subprovince, Superior Province, Canada. Chem. Geol., 98: 23-70. Ferreira, E.P., 1988. The sedimentology and stratigraphy of the Ahlmannryggen Group, Antarctica. M.Sc. Thesis (unpubl.), Univ. Stellenbosch, 193 pp. Ferreira, W.S., 1984. A physical comparison of subaerial and subaqueous eruptive environments in the Proterozoic Amisk Group, Saskatchewan, Canada. M.Sc. Thesis (unpubl.), Univ. Manitoba, 371 pp. Feybesse, J.-L. and Mil6si, J.-P., 1994. The Archaean/Proterozoic contact zone in West Africa: a mountain belt of d6collement thrusting and folding on a continental margin related to a 2.1 Ga convergence of Archaean cratons? Precambrian Res., 69: 199-227. Fialko, Y.A. and Rubin, A.M., 1999. Thermal and mechanical aspects of magma emplacement in giant dike swarms. J. Geophys. Res., B, Solid Earth and Planets, 104(10): 23,033-23,049. Fisher, R.V., 1961. Proposed classification of volcaniclastic sediments and rocks. Geol. Soc. Amer. Bull., 72: 1409-1414. Fisher, R.V., 1966. Rocks composed of volcanic fragments and their classification. Earth Sci. Rev., 1: 287-298.

814

References

Fisher, R.V. and Schmincke, H.-U., 1984. Pyroclastic Rocks. Springer-Verlag, New York, 472 pp. Fisher, R.V. and Smith, G.A., 1991. Volcanism, tectonics and sedimentation. In: R.V. Fisher and G.A. Smith (Eds.), Sedimentation in Volcanic Settings. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Vol., 45, pp. 1-5. Fisk, M.R. and Giovannoni, S.J., 1999. Sources of nutrients and energy for a deep biosphere on Mars. J. Geophys. Res., 104:11,805-11,815. Fiske, R.S., 1963. Subaqueous pyroclastic flows in the Ohanapecosh Formation, Washington. Geol. Soc. Amer. Bull., 74: 391-406. Fiske, R.S. and Matsuda, T., 1964. Submarine equivalents of ash flows in the Tokiwa Formation, Japan. Amer. J. Sci., 262: 76-106. Fiske, R.S. and Tobisch, O.T., 1994. Middle Cretaceous ash-flow tuff and caldera-collapse deposits in the Minarets Caldera, east-central Sierra Nevada, CA. Geol. Soc. Amer. Bull., 106: 582-593. Fiske, R.S., Naka, J., lizasa, K., Yuasa, M. and Klaus, A., 2001. Submarine silicic caldera at the front of the Izu-Bonin Arc, Japan; voluminous seafloor eruptions of rhyolite pumice. Geol. Soc. Amer. Bull., 113:813-824. Fitches, W.R., Graham, R.H., Hussein, I.M., Ries, A.C., Shackleton, R.M. and Price, R.C., 1983. The late Proterozoic ophiolite of Sol Hamed, NE Sudan. Precambrian Res., 19:385-4 11. Fitzsimons, I.C.W., 2000. Grenville-age basement provinces in East Antarctica: evidence for three separate collisional orogens. Geology, 28: 879-882. Fitzsimons, I.C.W., Kinny, P.D. and Harley, S.L., 1997. Two stages of zircon and monazite growth in anatectic leucogneiss: SHRIMP constraints on the duration and intensity of Pan-African metamorphism in Prydz Bay, East Antarctica. Terra Nova, 9:47-51. Fleet, M.E. and MacRae, N.D., 1975. A spinifex rock from Munro Township, Ontario. Can. J. Earth Sci., 12: 928-939. Flemings, P.B. and Jordan, T.E., 1990. Stratigraphic modelling of foreland basins: interpreting thrust deformation and lithospheric rheology. Geology, 18: 430435. Foden, J., Suck, H.S., Turner, S., Elburg, M., Smith, P.B., Van der Steldt, B. and Van Penglis, D., 2002. Geochemical evolution of lithospheric mantle beneath SE South Australia. Chem. Geol., 182: 663-695. Foerster, J.W., 1964. The use of calcium and magnesium ions to stimulate sheaths formation in Oscillatoria limosa (Roth) C.A. Agardh. Trans. American Microscop. Soc., 83: 420--427. Foley, S.E, 1992. Vein-plus-wall-rock melting mechanisms in the lithosphere and the origin of potassic alkaline magmas. Lithos, 28: 435-453. Foley, S., Tiepolo, M. and Vannucci, R., 2002. Growth of early continental crust controlled by melting of amphibolite in subduction zones. Nature, 417: 837-840. Foley, S.E, Venturelli, G., Green, D.H. and Toscani, L., 1987. The ultrapotassic rocks: characteristics, classification and constraints for petrogenetic models. Earth Sci. Rev., 24:81-134. FSlling, P.G. and Frimmel, H.E., 2002. Chemostratigraphic correlation of carbonate successions in the Gariep and Saldania Belts, Namibia/South Africa. Basin Res., 14: 69-88. Frlling, P.G., Zartman, R.E. and Frimmel, H.E., 2000. A novel approach to double-spike Pb-Pb dating of carbonate rocks: examples from Neoproterozoic sequences in southern Africa. Chem. Geol., 171: 97-122. Forsyth, D.W. and Uyeda, S., 1975. On the relative importance of the driving forces of plate motion. Geophys. J. Roy. Astr. Soc., 43: 163-200. Forterre, P., Confalonieri, E, Charbonnier, E and Duguet, M., 1995. Speculations on the origins of life and thermophily: review of available information reverse gyrase suggests that hyperthermophilic procaryotes are not so primitive. Origins Life Evol. Biosphere, 25: 235-249.

References

815

Foster, EN. and Boss, A.E, 1996. Triggering star formation with stellar ejecta. J. Astrophys., 468: 784-796. Fouquet, Y., Eissen, J.-E, Ondr6as, H., Barriga, E, Batiza, R. and Danyushevshy, L., 1998. Extensive volcaniclastic deposits at the Mid-Atlantic Ridge axis: results of deep-water basaltic explosive volcanic activity? Terra Nova, 10: 280-286. Fowler, A.D. and Roach, D.E., 1996. A model and simulation of branching mineral growth from cooling contacts and glasses. Mineral. Mag., 60:595-601. Fowler, A.D., Jensen, L.S. and Peloquin, S.A., 1986. Varioles in Archean basalts: products of spherulitic crystallization. Can. Mineral., 25: 275-289. Fowler, A.D., Stanley, H.E. and Daccord, G., 1989. Disequilibrium silicate minerals: fractal and nonfractal features. Nature, 341: 134-138. Fowler, A.D., Berger, B., Shore, M., Jones, M.I. and Ropchan, J., 2002. Supercooled rocks: development and significance of varioles, spherulites, dendrites and spinifex in Archean volcanic rocks, Abitibi Greenstone Belt, Canada. Precambrian Res., 115:311-328. Fox, J., 1976. Some comments on the volcanic stratigraphy and economic potential of the west Amisk Lake, Saskatchewan. Saskatchewan Research Council, Circ. 9, 31 pp. Fox, J.S. and Johnston, W.G.Q., 1981. Komatiites, "boninites" and tholeiitic picrites in the central La Ronge metavolcanic belt, Saskatchewan and Manitoba, and their possible economic significance. Can. Inst. Min. Metal. Bull., 74, 83 l: 73-82. Fralick, EW. and Miall, A.D., 1989. Sedimentology of the Lower Huronian Supergroup (Early Proterozoic), Elliot Lake area, Ontario, Canada. Sediment. Geol., 63: 127-153. Francis, D. and Ludden, J.N., 1990. The mantle source for olivine nephelinite, basanite, and alkali olivine basalt at Fort Selkirk, Yukon, Canada. J. Petrol., 3 l" 371-400. Francois, L.M., 1986. Extensive deposition of banded iron formations was possible without photosynthesis. Nature, 320: 352-354. Frank, T.D. and Lyons, T.W., 1998. "Molar-tooth" structures: a geochemical perspective on a Proterozoic enigma. Geology, 26: 683-686. Frank, T.D., Kah, L.C. and Lyons, T.W., 2003. Changes in organic matter production and accumulation as a mechanism for isotopic evolution in the Mesoproterozoic ocean. Geol. Mag., 140: 397-420. Frarey, M.J. and Krogh, T.E., 1986. U-Pb zircon ages of late internal plutons of the Abitibi and eastern Wawa subprovinces, Ontario and Quebec. Current Research, Part A., Geol. Surv. Can. Paper, 86-1 A, pp. 43-48. Frazier, D.E., 1974. Depositional episodes: their relationship to the Quaternary stratigraphic framework in the northwestern portion of the Gulf Basin. Bureau of Economic Geology, Univ. Texas at Austin, Geological Circular, 4(1), 28 pp. Frei, R. and Rosing, M.T., 2001. The least radiogenic terrestrial leads; implications for the early Archean crustal evolution and hydrothermal-metasomatic processes in the Isua supracrustal belt (West Greenland). Chem. Geol., 181: 47-66. Frei, R., Rosing, M., Waight, T.E., Krogstad, E.J., Storey, M., Ulfbeck, D.G. and Albar~de, E, 2002. Hydrothermal-metasomatic and tectono-metamorphic processes in the Isua greenstone belt (West Greenland): a multi-isotopic investigation of their effects on the Earth's oldest oceanic crustal sequence. Geochim. Cosmochim. Acta, 66: 467-486. French, B.M., 1998. Traces of Catastrophe: A Handbook of Shock-Metamorphic Effects in Terrestrial Meteorite Impact Structures. Lunar and Planetary Inst., Houston, Texas, 120 pp. Freund, E, Dickinson, J.T. and Cash, M., 2002. Hydrogen in rocks: an energy source for deep microbial communities. Astrobiology, 2: 83-92.

816

References

Friedmann, E.I., 1980. Endolithic microbial life in hot and cold deserts. Origins of Life, 10: 223-235. Friedmann, E.I. and Koriem, A.M., 1989. Life on Mars: how it disappeared (if it was ever there). Adv. Space Res., 9(6): 167-172. Friedman, G.M., Sanders, J.E. and Kopaska-Merkel, D.C., 1992. Principles of Sedimentary Deposits. Macmillan, New York, 717 pp. Friend, C.R.L., Bennett, V.C. and Nutman, A.P., 2002. Abyssal peridotites > 3,800 Ma from southern West Greenland: field relationships, petrography, geochronology, whole-rock and mineral chemistry of dunite and harzburgite inclusions in the Itsaq Gneiss Complex. Contrib. Mineral. Petrol., 143:71-92. Friend, C.R.L., Nutman, A.P., Baadsgaard, B., Kinny, P.D. and McGregor, V.R., 1996. Timing of late Archean terrane assembly, crustal thickening and granite emplacement in the Nuuk region, southern West Greenland. Earth Planet. Sci. Lett., 142: 353-365. Friend, P.E, 1978. Distinctive features of some ancient river systems. In: A.D. Miall (Ed.), Fluvial Sedimentology. Can. Soc. Petrol. Geol. Mem., 5, pp. 531-542. Frimmel, H.E., 2000a. The Pan-African Gariep Belt in southwestern Namibia and western South Africa. Communications of the Geological Survey of Namibia, 12: 197-209. Frimmel, H.E., 2000b. The stratigraphy of the Chameis Sub-terrane in the Gariep Belt in southwestern Namibia. Communications of the Geological Survey of Namibia, 12:179-186. Frimmel, H.E. and Frank, W., 1998. Neoproterozoic tectono-thermal evolution of the Gariep Belt and its basement, Namibia/South Africa. Precambrian Res., 90: 1-28. Frimmel, H.E. and F611ing, P.G., in prep. Late Vendian closure of the Adamastor Ocean: Timing of tectonic inversion and syn-orogenic sedimentation in the Gariep Basin. Gondwana Res. Frimmel, H.E. and Hartnady, C.J.H., 1992. Blue amphiboles and their significance for the metamorphic history of the Pan-African Gariep belt, Namibia. J. Metamorphic Geol., 10:651-669. Frimmel, H.E. and Jiang, S.-Y., 2001. Marine evaporites from an oceanic island in the Neoproterozoic Adamastor ocean. Precambrian Res., 105: 57-71. Frimmel, H.E., Deane, J.G. and Chadwick, P.J., 1996a. Pan-African tectonism and the genesis of base metal sulfide deposits in the northern foreland of the Damara Orogen, Namibia. In: D.E Sangster (Ed.), Carbonate-Hosted Lead-Zinc Deposits. Soc. Econ. Geol. Spec. Publ., 4, pp. 204-217. Frimmel, H.E., F611ing, P.G. and Eriksson, P.G., 2002. Neoproterozoic tectonic and climatic evolution recorded in the Gariep Belt, Namibia/South Africa. Basin Res., 14: 55-67. Frimmel, H.E., Hartnady, C.J.H. and Koller, E, 1996b. Geochemistry and tectonic setting of magmatic units in the Pan-African Gariep Belt, Namibia. Chem. Geol., 130: 101-121. Frimmel, H.E., K16tzli, U. and Siegfried, P., 1996c. New Pb-Pb single zircon age constraints on the timing of Neoproterozoic glaciation and continental break-up in Namibia. J. Geol., 104: 459-469. Frimmel, H.E., Zartman, R.E. and Sp~ith, A., 2001. Dating Neoproterozoic continental break-up in the Richtersveld Igneous Complex, South Africa. J. Geol., 109: 493-508. Frisch, T., Jackson, G.D., Glebovitsky, V.A., Yefimov, M.M., Bogdanova, M.N. and Parrish, R.R., 1995. U-Pb ages in zircon from the Kolvitsa gabbro--anorthosite complex, southern Kola Peninsula, Russia. Petrology, 3(3): 219-225 (official English translation). Fritz, W.J. and Moore, J.N., 1988. Basics of Physical Stratigraphy and Sedimentology. John Wiley and Sons, New York, 371 pp. Froude, D.O., Ireland, T.R., Kinny, P.K., Williams, I.S., Compston, W., Williams, I.R. and Myers, J.S., 1983. Ion Microprobe identification of 4,100--4,200 Myr-old terrestrial zircons. Nature, 304:616-618.

References

817

Fryberger, S.G., Hesp, E and Hastings, K., 1992. Aeolian granule ripple deposits, Namibia. Sediment. Geol., 39:319-333. Fryberger, S.G., Schenk, C.J. and Hesp, E, 1988. Pin stripe lamination; a distinctive feature of modern and ancient eolian sediments. Sediment. Geol., 55:1-15. Fryer, B.J., 1983. Rare earth elements in iron-formation. In: A.F. Trendall and R.C. Morris (Eds.), Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 345-358. Fujioka, K., Kobayashi, K., Kato, K., Aoki, M., Mitsuzawa, K., Kinoshita, M. and Nishizawa, A., 1997. Tide-related variability of TAG hydrothermal activity observed by deep-sea monitoring system and OBSH. Earth Planet. Sci. Lett., 153: 239-250. Fuller, A.O., 1985. A contribution to the conceptual modelling of pre-Devonian fluvial systems. Trans. Geol. Soc. S. Afr., 88: 189-194. Fumes, H. and Sturt, B.A., 1976. Beach/shallow marine hyaloclastite deposits and their geological significance--an example from Gran Canaria. J. Geol., 84: 439-453. Furnes, H., Fridleifsson, I.B. and Atkins, EB., 1980. Subglacial volcanics---on the formation of acid hyaloclastites. J. Volcanol. Geotherm. Res., 54:95-110. Furnes, H., Shimron, A.E. and Roberts, D., 1985. Geochemistry of Pan-African volcanic arc sequences in southeastern Sinai Peninsula and plate tectonic implications. Precambrian Res., 29: 359-382. Furniss, G., Rittel, J.E and Winston, D., 1998. Gas bubble and expansion crack origin of "molartooth" calcite structures in the Middle Proterozoic Belt Supergroup, western Montana. J. Sediment. Res., 68:104-114. Fyfe, W.S., 1978. The evolution of the Earth's crust: modern plate tectonics to ancient hotspot tectonics? Chem. Geol., 23:89-114. Fysson, W.K., Herd, R.K. and Ermanovics, I.E, 1978. Diapiric structures and regional compression in an Archean greenstone belt, Island Lake, Manitoba. Can. J. Earth Sci., 15: 1817-1825. Ga~il, G., 1992. Global Proterozoic tectonic cycles and Early Proterozoic metallogeny. S. Afr. J. Geol., 95(3): 79-87. Ga~il, G. and Gorbatschev, R., 1987. An outline of the Precambrian evolution of the Baltic Shield. Precambrian Res., 35: 15-52. Gala, M.G., Symons, D.T.A. and Palmer, H.C., 1998. Geotectonics of the Hanson Lake block, Trans-Hudson orogen, central Canada: a preliminary paleomagnetic report. Precambrian Res., 90: 85-101. Galer, S.J.G. and Metzger, K., 1988. Metamorphism, denudation and sea level in the Archean and cooling of the Earth. Precambrian Res., 92: 389-412. Gall, Q., Peterson, T.D. and Donaldson, J.A., 1992. A proposed revision of early Proterozoic stratigraphy of the Thelon and Baker Lake basins, Northwest Territories. Current Res., Part C, Geol. Surv. Can. Paper, 92-1 C: 129-137. Gallardo, V.A., 1977. Large benthic microbial communities in sulfide biota under Peru-Chile subsurface counter current. Nature, 268:331-332. Galley, A.G., 1993. Characteristics of semi-conformable alteration zones associated with volcanogenic massive sulphide districts. J. Geochem. Explor., 48: 175-200. Galley, A.G., 1994. Geology of the Ansil Cu-Zn massive sulfide deposit, Rouyn-Noranda, Quebec, Canada. Ph.D. Thesis (unpubl.), Carleton University, Ottawa, 386 pp. Galloway, W.E., 1989. Genetic stratigraphic sequences in basin analysis, I. Architecture and genesis of flooding-surface bounded depositional units. Amer. Assoc. Petrol. Geol. Bull., 73: 125-142.

818

References

Galloway, W.E., 2001. The many faces of submarine erosion: theory meets reality in selection of sequence boundaries. Abstracts, A.A.EG. Hedberg Research Conference on "Sequence Stratigraphic and Allostratigraphic Principles and Concepts", Dallas, August 26-29, 2001, pp. 28-29. Galloway, W.E. and Hobday, D.K., 1983. Terrigenous Clastic Depositional Systems. Springer-Verlag, New York, 423 pp. Galloway, W.E. and Hobday, D.K., 1996. Terrigenous Clastic Depositional Systems, 2nd edn. Springer-Verlag, Berlin, 489 pp. Galy, A., Young, E.D., Ash, R.D. and O'Nions, R.K., 2000. The formation of chondrules at high pressures in the solar nebula. Science, 290:1751-1753. Garlick, W.G., 1981. Sabkhas, slumping, and compaction at Mufulira, Zambia. Econ. Geol., 76: 1817-1847. Garlick, W.G., 1988. Algal mats, load structures, and synsedimentary sulfides in Revett Quartzites of Montana and Idaho. Econ. Geol., 83: 1259-1278. Garrels, R.M., 1967. Genesis of some ground waters from Igneous rocks. In: P.H. Abelson (Ed.), Researches in Geochemistry 2. John Wiley and Sons, Chichester, pp. 405-422. Garrels, R.M., 1987. A model for the deposition of the microbanded Precambrian iron formations. Amer. J. Sci., 287: 81-106. Garrels, R.M. and Mackenzie, E T., 1967. Origin of the chemical composition of some springs and lakes. In: Equilibrium Concepts of Natural Water System. Amer. Chem. Soc., Adv. Chem., 67, pp. 222-242. Gauthier-Lafaye, E, Holliger, E and Blanc, EL., 1996. Natural fission reactors in the Franceville Basin, Gabon: a review of the conditions and results of a "critical event" in a geologic system. Geochim. Cosmochim. Acta, 60:4831-4852. Gebelein, C.D. and Hoffman, P., 1973. Algal origin of dolomite laminations in stromatolitic limestone. J. Sediment. Petrol., 43:603--613. Gee, R.D., 1979. Structure and tectonic style of the Western Australian Shield. Tectonophysics, 58: 327-362. Gee, R.D. and Grey, K., 1993. Proterozoic rocks on the Glengarry 1:250000 sheet--stratigraphy, structure and stromatolite biostratigraphy. Geological Survey of Western Australia, Report 41, 33 pp. Gehling, J.G., 1999. Microbial mats in terminal Proterozoic siliciclastics: Ediacaran death masks. Palaios, 14: 40-57. Gehling, J.G., 2000. Environmental interpretation and a sequence stratigraphic framework for the terminal Proterozoic Ediacara Member within the Rawnsley Quartzite, South Australia. Precambrian Res., 100: 65-95. G61inas, L. and Brooks, C., 1974. Archean quench-textured tholeiites. Can. J. Earth Sci., 11" 324-340. G61inas, L., Brooks, C. and Trzcienski, W.E., 1977b. Archean variolites quenched immiscible liquids. Can. J. Earth Sci., 13:21 0-230. G61inas, L., Lajoie, J. and Brooks, C., 1977a. The origin and significance of Archaean ultramafic volcaniclastics from Spinifex Ridge, Lamotte Township, Qu6bec. In: W.R.A. Baragar, L.C. Coleman and J.M. Hall (Eds.), Volcanic Regimes in Canada. Geol. Assoc. Can. Spec. Paper, 16, pp. 297-309. Gellatly, A.M. and Lyons, T.W., submitted. Trace sulfate in Mesoproterozoic carbonates and the implications for sea water sulfate and oxygen availability. Geochim. Cosmochim. Acta. Gerdes, G. and Krumbein, W.E., 1987. Biolaminated Deposits. Lecture Notes in Earth Sciences, 9. Springer-Verlag, New York, 183 pp.

References

819

Gerdes, G., Claes, M., Dunajtschik-Piewak, K., Riege, H., Krumbein, W.E. and Reineck, H.-E., 1993. Contribution of microbial mats to sedimentary surface structures. Facies, 29:61-74. Gerdes, G., Krumbein, W.E. and Reineck, H.-E., 1985. The depositional record of sandy, versicolored tidal flats (Mellum Island, southern North Sea). J. Sediment. Petrol., 55: 265-278. Gerdes, G., Noffke, N., Klenke, Th. and Krumbein, W.E., 2000. Microbial signatures in peritidal sediments--a catalogue. Sedimentology, 47: 279-308. Gerlach, D.C., Shirey, S.B. and Carlson, R.W., 1988. Nd isotopes in Proterozoic iron formations: evidence for mixed-age provenance and depositional variability. Eos, 69(44): 1615. Germs, G.J.B., 1983. Implications of a sedimentary facies and depositional environmental analysis of the Nama Group in South West Africa/Namibia. In: R.M. Miller (Ed.), Evolution of the Damara Orogen of South West Africa/Namibia. Geol. Soc. S. Afr., Johannesburg, Spec. Publ., 11, pp. 89-114. Gersonde, R., Kyte, E T., B leil, U., Diekmann, B., Flores, J.A., Gohl, K., Grahl, G., Hagen, R., Kuhn, G., Sierro, EJ., V61ker, D., Abelmann, A. and Bostwick, J.A., 1997. Geological record and reconstruction of the late Pliocene impact of the Eltanin asteroid in the Southern Ocean. Nature, 390: 357-363. Gerya, T.V., Percuk, L.L., Van Reenen, D.D. and Smit, C.A., 2000. Two-dimensional numerical modeling of pressure-temperature-time paths for the exhumation of some granulite facies terrains in the Precambrian. J. Geodynamics, 30:17-35. Gessner, E, 1957. Van G61tiIZUr Limnologie des grol3en Soda Sees in Ostanatolien (Ttirkei). Arch. Hydrobiol., 53: 1-22. Ghosh, S.K., 1991. Palaeoenvironmental analysis of the Late Proterozoic Nagthat Formation, NV Kumaun Lesser Himalaya, India. Sediment. Geol., 71: 33-45. Gibb, R.A., 1978. Slave-Churchill collision tectonics. Nature, 271: 50-52. Gibb, R.A., 1983. Model for suturing of Superior and Churchill plates: an example of double indentation tectonics. Geology, 11: 413-417. Gibbs, A.D., 1987. Development of extension and mixed-mode sedimentary basins. In: M.P. Coward, J.E Dewey and P.L. Hancock (Eds.), Continental extensional tectonics. Geol. Soc. London Spec. Publ., 28, pp. 19-33. Gibbs, M.T., Rees, P.McA., Kutzbach, J.E., Ziegler, A.M., Behling, P.J. and Rowley, D.B., 2002. Simulations of Permian climate and comparisons with climate-sensitive sediments. J. Geol., 110: 33-55. Gibson, G.M. and Ireland, T.R., 1995. Granulite formation during continental extension on Fiordland, New Zealand. Nature, 375: 479-482. Gibson, G.M. and Ireland, T.R., 1999. Black Giants Anorthosite, New Zealand: a Paleozoic analogue of Archean stratiform anorthosite and implications for the formation of Archean high-grade terranes. Geology, 27(2): 131-134. Gibson, H.L., 1989. The Mine Sequence of the Central Noranda volcanic complex: Geology, alteration, massive sulfide deposits and volcanological reconstruction. Ph.D. Thesis (unpubl.), Carleton University, Ottawa, 715 pp. Gibson, H.L. and Watkinson, D.H., 1990. Volcanogenic massive sulphide deposits of the Noranda cauldron and shield volcano, Quebec. In: M. Rive, P. Verpaelst, Y. Gagnon, J.-M. Lulin, G. Riverin and A. Simard (Eds.), La Ceinture Polym6tallique du Nord-Ouest Qu6b6cois. Can. Inst. Min. Metall., Spec. Vol. 43, pp. 119-132. Gibson, H.L., Morton, R. and Hudak, G., 1997. Subaqueous volcanism: environments and controls on VMS mineralization. In: T.C. Barrie and M.D. Hannington (Eds.), Volcanic-Associated Massive

820

References

Sulfide Deposits: Processes and Examples in Modern and Ancient Settings. GAC-MDD-SEG Short Course. Geol. Assoc. Can.-Mineral. Assoc. Can., pp. 13-51. Gilbert, H.P., Syme, E.C. and Zwanzig, H.V., 1980. Geology of the metavolcanic and volcaniclastic metasedimentary rocks in the Lynn Lake area. Manitoba Energy and Mines, Geol. Paper, GP80-1, 118 pp. Gilichinsky, D.A., 2002. Permafrost model of extraterrestrial habitat. In: G. Horneck and C. Baumstark-Khan (Eds.), Astrobiology: The Quest for the Conditions of Life. Springer-Verlag, Berlin, pp. 125-142. Gill, J., 1981. Orogenic Andesites and Plate Tectonics. Springer-Verlag, Heidelberg, 390 pp. Gill, J., Torssandu, P., Lapierre, H., Taylor, R., Kaiho, K., Koyama, M., Kusakabe, M., Aitchison, J., Cisonski, C., Dadey, K., Fujioka, K., Klaus, A., Lovell, M., Marsaglia, K., Pesard, P., Taylor, B. and Tazaki, K., 1990. Explosive deep water basalt in the Sumisu backarc rift. Science, 248: 1214-1217. Gilmour, J.D. and Saxton, J.M., 2001. A time-scale of formation of the first solids. Philos. Trans. Roy. Soc. London Ser. A, 359: 2037-2048. Glaser, K.S. and Droxler, A.W., 1991. High production and highstand shedding from deeply submerged carbonate banks, northern Nicaragua Rise. J. Sediment. Petrol., 61: 128-142. Glass, B.E, 1984. Solution of naturally-occurring glasses in the geological environment. J. NonCrystalline Solids, 67: 265-286. Glass, B.E, 1989. North American tektite debris and impact ejecta from DSDP Site 612. Meteoritics, 24:209-218. Glass, B.E, 1990. Tektites and microtektites: key facts and inferences. Tectonophysics, 171: 393-404. Glass, B.E, Swincki, M.B. and Zwart, EA., 1979. Australasian, Ivory Coast and North American tektite strewnfields; size, mass and correlation with geomagnetic reversals and other Earth events. Proc. Lunar Planet. Sci. Conf., 3: 2535-2545. Glebovitsky, V.A., Alexeyev, N.L. and Dolivo-Dobrovolcky, D.V., 1997. Reaction structures and T-P regimes of cooling of the deep crustal formations in the Kandalaksha-Kolvitsa structural zone, Kola Peninsula. Proc. Russ. Mineral. Soc., 2:1-22 (in Russian). Glikson, A.Y., 1972. Early Precambrian evidence of primitive ocean crust and island nucleii of sodic granite. Geol. Soc. Arner. Bull., 83: 3323-3344. Glikson, A.Y., 1976. Trace element geochemistry and origin of early Precambrian acid igneous rock series, Barberton Mountain Land, Transvaal. Geochim. Cosmochim. Acta, 40: 1261-1280. Glikson, A.Y., 1979. Early Precambrian tonalite-trondhjemite sialic nuclei. Earth Sci. Rev., 15: 1-73. Glikson, A.Y., 1999. Oceanic mega-impacts and crustal evolution. Geology, 27: 387-390. Goff, S.E, Godfrey, J.D. and Holland, J.G., 1986. Geochemistry of the Canadien Shield of Northeastern Alberta. Bull. Alberta Res. Council, Edmonton, 51, 60 pp. Goldhaber, M.B. and Kaplan, I.R., 1974. The sulfur cycle. In: E.D. Goldberg (Ed.), The Sea, Vol. 5. John Wiley and Sons, New York, pp. 569-655. Goldich, S.S., 1973. Ages of Precambrian iron-formations. Econ. Geol., 68:1126-1135. Goldreich, E, 1966. History of the lunar orbit. Rev. Geophys., 4:411-439. Goldstein, S.L., Arndt, N.T. and Stallard, R.F., 1997. The history of a continent from Sm-Nd isotopes in Orinoco basin river sediments and U-Pb ages of zircons from Orinoco river sand. Chem. Geol., 139:271-298. Gole, M., 1981. Archean banded iron-formations, Yilgarn Block, Western Australia. Econ. Geol., 76: 1954-1974.

References

82i

Gole, M. and Klein, C., 1981. Banded iron-formations through much of Precambrian time. J. Geol., 89: 169-183. Goodfellow, W.D. and Jonasson, I.R., 1984. Ocean stagnation and ventilation defined by ~34S secular trends in pyrite and barite, Selwyn Basin, Yukon. Geology, 12: 583-586. Goodwin, A.M., 1982. Archean volcanoes in the southwestern Abitibi Belt, Ontario and Quebec: form, composition, and development. Can. J. Earth Sci., 19:1140-1155. Goodwin, A.M., 1996. Princples of Precambrian Geology. Academic Press, London, 327 pp. Gorbatschev, R. and Bogdanova, S., 1993. Frontiers in the Baltic Shield. Precambrian Res., 64: 3-21. Gordon, R.G., McWilliams, M.O. and Cox, A., 1979. Pre-Tertiary velocities of the continents: a lower bound from palaeomagnetic data. J. Geophys. Res., 84: 5480-5486. Gorjan, E, Veevers, J.J. and Walter, M.R., 2000. Neoproterozoic sulfur-isotope variation in Australia and global implications. Precambrian Res., 100:151-179. Gorman, B.E., Pearce, T.H. and Birkett, T.C., 1978. On the structure of Archaean greenstone belts. Precambrian Res., 6: 23-41. Gose, W.A., Helper, M.A., Connelly, J.N., Hutson, EE. and Dalziel, I.W.D., 1997. Paleomagnetic data and U-Pb isotopic age determinations from Coats Land, Antarctica: implications for Neoproterozoic plate reconstructions. J. Geophys. Res., 102: 7887-7902. Goswami, J.N., 2000. Chronology of early solar system events: dating with short-lived nuclides. In: O. Manuel (Ed.), Origin of Elements in the Solar System: Implications of Post 1957 Observations. Kluwer Academic/Plenum Publishers, New York, pp. 407-430. Goth, K., 1990. Der Messeler Oelschiefer; ein Algenlaminit. CFS Courier Forschungsinstitut Senckenberg, 131, 141 pp. Goth, K. and Schiller, W., 1994. Miozaene Algenlaminite von Hausen in der Rhoen. Palaeontologische Zeitschrift, 68: 287-297. Goto, Y. and McPhie, J., 1998. Endogenous growth of a Miocene submarine cryptodome, Rebun Island, Hokkaido, Japan. J. Volcanol. Geotherm. Res., 84: 273-286. Goulet, N., 1978. Stratigraphy and structural relationships across the Cadillac-Larder Lake fault, Rouyn-Beauchatel area, Qu6bec. Ph.D. Thesis (unpubl.), Queen's University, Kingston, Ontario, Canada, 141 pp. Goutier, J., Melanqon, M., Dion, C., Verpaelst, E and Rive, M., 1994. Compilation des datations isotopiques des sous-Provinces de l'Abitibi et quelques unit6s adjacentes (Qu6bec et Ontario). Minist6re des Ressources naturelles du Qu6bec, MB 94-55, 42 pp. Gradstein, F.M., Sandvik, K.O. and Milton, N.J., 1998. Sequence StratigraphymConcepts and Applications. Norwegian Petroleum Society (NPF) Spec. Publ., 8. Elsevier, Amsterdam, 437 pp. Grand, S.E, van der Hilst, R.D and Widiyantoro, S., 1997. Global seismic tomography: a snapshot of convection in the Earth. GSA Today, 7(4): 1-7. Grantham, G.H., Jackson, C., Moyes, A.B., Groenewald, EB., Harris, ED., Ferrar, G. and Krynauw, J.R., 1995. The tectonothermal evolution of the Kirwanveggen-H.U. Sverdrupfjella areas, Dronning Maud Land, Antarctica. Precambrian Res., 75: 209-229. Grantham, J.H. and Velbel, M.A., 1988. The influence of climate and topography on rock-fragment abundance in modern fluvial sands of the southern Blue Ridge Mountains, North Carolina. J. Sediment. Petrol., 58:219-227. Grassineau, N.V., Nisbet, E.G., Bickle, M.J., Fowler, C.M.R., Lowry, D., Mattey, D.E, Abell, E and Martin, A., 2001. Antiquity of the biological sulphur cycle: evidence from sulphur and carbon isotopes in 2700 million-year-old rocks of the Belingwe Belt, Zimbabwe. Proc. Roy. Soc. London Bio. Sci., 268:113-119. Gray, C.M. and Compston, W., 1974. Excess 26Mg in the Allende meteorite. Nature, 251: 495-497.

822

References

Greeley, R., 1982. The Snake River Plain, Idaho: representative of a new category of volcanism. J. Geophys. Res., 87: 2705-2712. Green, A.G., Hajnal, Z. and Weber, W., 1985. An evolutionary model of the western Churchill Province and western margin of the Superior Province in Canada and the north-central United States. Tectonophysics, 116:281-322. Green, M.G., Sylvester, J. and Buick, R., 2000. Growth and recycling of early Archaean continental crust: geochemical evidence from the Coonterunah and Warrawoona Groups, Pilbara Craton, Australia. Tectonophysics, 322: 69-88. Greff-Lefftz, M. and Legros, H., 1999a. Core rotational dynamics and geological events. Science, 286:1707-1709. Greff-Lefftz, M. and Legros, H., 1999b. Correlation between some major geological events and resonances between the free core nutation and luni-solar tidal waves. Geophys. J. Internat., 139: 131-151. Gresse, P.G., 1992. The tectono-sedimentary history of the Vanrhynsdorp Group. Mem. Geol. Surv. S. Afr., 79: 1-163. Grey, K., 1984. Biostratigraphic studies of stromatolites from the Proterozoic Earaheedy Group, Nabbera Basin, Western Australia. Bull. Geol. Surv. Western Australia, 128:221-241. Grey, K., 2001. Surviving the snowball earth: The acritarch record. Geol. Soc. Aust. Abs., 65: 45-47. Grey, K., 2002. Ediacarian palynology of Australia. Mem. Aust. Ass. Palaeontol. (in press). Grieve, R.A.E, 1998. Extraterrestrial impacts on Earth: the evidence and the consequences. In: M.M. Grady, R. Hutchison, G.J.H. McCall and D.A. Rothery (Eds.), Meteorites: Flux with Time and Impact Effects. Geol. Soc. London Spec. Publ., 140, pp. 105-131. Griffin, W.L., O'Reilly, S.Y., Ryan, C.G., Gaul, O. and lonov, D.A., 1998. Secular variation in the composition of subcontinental lithospheric mantle: geophysical and geodynamic implications. Geol. Soc. America, Amer. Geophys. Union, Geodynamic Ser., 26: 1-25. Groenewald, P.B., Moyes, A.B., Grantham, G.H. and Krynauw, J.R., 1995. East Antarctic crustal evolution: geological constraints and modelling in western Dronning Maud Land. Precambrian Res., 75:231-250. Gross, G.A., 1965. Geology of Iron Deposits in Canada; Vol. I, General Geology and Evaluation of Iron Deposits. Geol. Surv. Can., Econom. Geol. Rep. 22, 181 pp. Gross, G.A., 1967. Geology and Iron Deposits of Canada, Vol. II, Iron Deposits in the Appalachian and Grenville Regions of Canada. Geol. Surv. Can., Econom. Geol. Rep. 22, 179 pp. Gross, G.A., 1968. Geology of Iron Deposits of Canada, Vol. III, Iron Ranges in the Labrador Geosyncline. Geol. Surv. Can., Econom. Geol. Rep. 22, 111 pp. Gross, G.A., 1973. The depositional environment of the principal types of Precambrian iron formations. Proceedings Kiev Symposium: Genesis of Precambrian Iron and Manganese Deposits, 1970. UNESCO, Paris, Earth Sciences 9, pp. 15-21. Gross, G.A., 1980. A classification of iron formations based on depositional environments. Can. Mineral., 18: 215-222. Gross, G.A., 1991. Genetic concepts for iron-formation and metalliferous sediments. In: R.W. Hutchinson and R.I. Grauch (Eds.), Historical Perspectives of Genetic Concepts and Case Histories of Famous Discoveries. Economic Geology Monograph 8, pp. 51-81. Gross, M.G., 1990. Oceanography, 5th edn. Prentice Hall, Englewood Cliffs, New Jersey, 441 pp. Grotzinger, J.P., 1989. Facies and evolution of Precambrian carbonate depositional systems: emergence of the modem platform archetype. In: P.D. Crevello, J.L. Wilson, J.E Sarg and J.F. Read (Eds.), Controls on Carbonate Platform and Basin Development. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 44, pp. 79-101.

References

823

Grotzinger, J.E, 1990. Geochemical model for Proterozoic stromatolite decline. Amer. J. Sci., 290-A: 80-103. Grotzinger, J.P., 1994. Trends in Precambrian carbonate sediments and their implications for understanding evolution. In: S. Bengtson (Ed.), Early Life on Earth. Nobel Symposium, No. 84. Columbia Univ. Press, New York, pp. 245-258. Grotzinger, J.P. and James, N.P., 2000. Carbonate Sedimentation and Diagenesis in the Evolving Precambrian World. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 67, 364 pp. Grotzinger, J.E and Kasting, J.E, 1993. New constraints of Precambrian ocean composition. J. Geol., 101: 235-243. Grotzinger, J.P. and Knoll, A.H., 1995. Anomalous carbonate precipitates: is the Precambrian the key to the Permian? Palaios, 10: 578-596. Grotzinger, J.E and Knoll, A.H., 1999. Stromatolites in Precambrian carbonates: Evolutionary mileposts or environmental dipsticks? Ann. Rev. Earth Planet. Sci., 27: 313-358. Grotzinger, J.E and Rothmans, D.H., 1996. An abiotic model for stromatolite morphogenesis. Nature, 383: 423-425. Grotzinger, J.E, Bowring, S.A., Saylor, B.Z. and Kaufman, A.J., 1995. Biostratigraphic and geochronologic constraints on early animal evolution. Science, 270: 598-604. Grove, T.L., Gaetani, G.A. and de Wit, M.J., 1994. Spinifex textures in the 3.49 Ga Barberton mountain belt komatiites: evidence for crystalization of water-bearing cool magmas in the Archean. Eos. Trans. AGU, 75: 354. Grove, T.L., Gaetani, G.A., Parman, S., Dann, J. and de Wit, M.J., 1996. Origin of spinifex textures in 3.49 Ga komatiite magmas from the Barberton Mountainland South Africa. Eos, 77:281. Grove, T.L., Parman, S.W. and Dann, J.C., 1999. Conditions of magma generation for Archean komatiites from the Barberton Mountainland, South Africa. In: Y. Fei, C.M. Bertka and B.O. Mysen (Eds.), Mantle Petrology: Field Observations and High-Pressure Experimentation. Geochemic. Soc., Houston, pp. 155-167. Groves, D.A., Morton, R.L. and Franklin, J.M., 1988. Physical volcanology of the footwall rocks near the Mattabi massive sulfide deposit, Sturgeon Lake, northwestern Ontario. Can. J. Earth Sci., 25: 280-291. Groves, D.I., Dunlop, J.S.R. and Buick, R., 1981. An early habitat of life. Sci. Amer., 245: 64-73. Gruau, G., Chauvel, C. and Jahn, B.M., 1990. Anomalous Sm-Nd ages for the Early Archean Onverwacht Group volcanics. Contrib. Mineral. Petrol., 104: 27-34. Gruau, G., Jahn, B.-M., Glikson, A.Y., Davy, R., Hickman, A.H. and Chauvel, C., 1987. Age of the Archaean Talga-Talga subgroup, Pilbara Block, Western Australia, and early evolution of the mantle: new Sm-Nd isotopic evidence. Earth Planet. Sci. Lett., 85:105-116. Gruau, G., Rosing, M., Bridgwater, D. and Gill, R.C.O., 1996. Resetting of Sm-Nd systematics during metamorphism of > 3.7 Ga rocks: implications for isotopic models of early Earth differentiation. Chem. Geol., 133: 225-240. Gruau, G., Toupin S., Fourcade, S. and Blais, S., 1992. Loss of isotopic (Nd, O) and chemical LREE memory during metamorphism of komatiites; new evidence from eastern Finland. Contrib. Mineral. Petrol., 117: 66-82. Gruner, J.W., 1922. The origins of sedimentary iron formations: the Biwabik formation of the Mesabi Range. Econ. Geol., 17: 407-460. Gruner, J.W., 1924. Contributions to the geology of the Mesabi Range. Minnesota Geol. Surv. Bull., 19: 1-67. Gruner, J.W., 1946. The Mineralogy and Geology of the Taconites and Iron Ores of the Mesabi Range, Minnesota. Iron Range Resources and Rehabilitation Commission, St. Paul, 127 pp.

824

References

Guilbert, J.M. and Park, C.E, 1986. The Geology of Ore Deposits, Freeman, San Francisco, CA, 985 pp. Gtirich, G., 1906. Les Spongiostromides du Visen de la province de Namur. Mus6e Royal d'Histoire Naturelle de Belgique, Memoir 3, 55 pp. Gurnis, M., 1988. Large-scale mantle convection and the aggregation and dispersal of supercontinents. Nature, 332: 695-699. Gurnis, M., 1990. Ridge spreading, subduction and sea level fluctuations. Science, 250: 970-972. Gurnis, M. and Torsvik, T.H., 1994. Rapid drift of large continents during the late Precambrian and Palaeozoic: palaeomagnetic constraints and dynamic models. Geology, 22: 1023-1026. Gurov, E.P., Gurova, E.P. and Rakitskaya, R.B., 1995. Impact diamonds in the craters of the Ukrainian Shield. Meteoritics Planet. Sci., 30:515-516. Gurov, E.P., Gurova, E.P. and Rakitskaya, R.B., 1996. Impact diamonds of the Zapadnaya Crater; phase composition and some properties. Meteoritics Planet. Sci., 31: 56. Gusev, G.S. and Peskov, A.I., 1992. The Akitkan ridge Proterozoic trachyandesites and mafites (northern Baikal region): geochemistry and tectonic nature. Geotectonics, 4:75-86 (in Russian). Gutzmer, J. and Beukes, N.J., 1998. High-grade manganese ores in the Kalahari manganese field: Characterisation and dating of ore-forming events. Unpubl. rep., Johannesburg, 221 pp. Gutzmer, J., Beukes, N.J., Pickard, A. and Barley, M.E., 2000. 1170 SHRIMP age for Koras Group bimodal volcanism, Northern Cape Province. S. Afr. J. Geol., 103: 32-37. Habicht, K.S. and Canfield, D.E., 1996. Sulphur isotope fractionation in modern microbial mats and the evolution of the sulphur cycle. Nature, 382: 342-343. Habicht, K.S. and Canfield, D.E., 2001. Isotope fractionation by sulfate-reducing natural populations and the isotopic composition of sulfide in marine sediments. Geology, 29: 555-558. Habicht, K.S., Gade, M., Thamdrup, B., Berg, P. and Canfield, D., 2002. Calibration of sulfate levels in the Archean Ocean. Science, 298: 2372-2374. Haddad, D., Watts, A.B. and Lindsay, J.E, 2001. Evolution of the intracratonic Officer Basin, central Australia: implications from subsidence analysis and gravity modelling. Basin Res., 13: 217-238. Hadlari, T. and Rainbird, R.H., 2000. Sequence stratigraphy and sedimentology of the Paleoproterozoic Baker Lake Group in the Baker Lake Basin, Thirty Mile Lake, Nunavut Territory. Geol. Surv. Can., Current Res. 2000-C9, 10 pp. Hagadorn, J.W. and Bottjer, D.J., 1997. Wrinkle structures: microbially mediated sedimentary structures common in subtidal siliciclastic settings at the Proterozoic-Phanerozoic transition. Geology, 25: 1047-1050. Hagadorn, J.W. and Bottjer, D.J., 1999. Restriction of a late Neoproterozoic biotope: suspectmicrobial structures and trace fossils at the Vendian-Cambrian transition. Palaios, 14: 73-85. Hagadorn, J.W., Pfltiger, E and Bottjer, D.J., 1999. Unexplored microbial worlds. Palaios, 14: 1-2. Hahn, T.-M. and Runnegar, B., 1992. Megascopic eukaryotic algae from the 2.1 billion-year-old Negaunee Iron-Formation, Michigan. Science, 257: 232-235. Hajnal, Z., Lucas, S., White, D., Lewry, J., Bezdan, S., Stauffer, M.R. and Thomas, M.D., 1996. Seismic reflection images of high-angle faults and linked detachments in the Trans-Hudson Orogen. Tectonics, 15(2): 427-439. Hall, J.D., 1883. Cryptozo6n (poliferum) n.g. and s.p. Rep. New York State Mus., 36: 1-6. Hallam, A., 1981. Facies Interpretation and the Stratigraphic Record. W.H. Freeman, Oxford, 291 pp. Hallberg, J.A., 1986. Archaean basin development and crustal extension in the northeastern Yilgarn Block, Western Australia. Precambrian Res., 31: 133-156. Halliday, A.N., 2000. Terrestrial accretion rates and the origin of the Moon. Earth Planet. Sci. Lett., 176: 17-30.

References

825

Halliday, A.N. and Porcelli, D., 2001. In search of lost planets--the palaeocosmochemistry of the inner solar system. Earth Planet. Sci. Lett., 192: 545-559. Halliday, A.N., Lee, D.-C. and Jacobsen, S.B., 2000. Tungsten isotopes, the timing of metal-silicate fractionation, and the origin of the Earth and .Moon. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 45-62. Hambrey, M. J., 1983. Correlation of Late Proterozoic tillites in the North Atlantic region and Europe. Geol. Mag., 120: 209-220. Hambrey, M.J. and Harland, W.B. (Eds.), 1981. Earth's Pre-Pleistocene Glacial Record. Cambridge Univ. Press, Cambridge, 1004 pp. Hambrey, M.J. and Harland, W.B., 1985. The late Proterozoic glacial era. Palaeogeog. Palaeoclimatol. Palaeoecol., 51: 255-272. Hamilton, D.S. and Tadros, N.Z., 1994. Utility of coal seams as genetic stratigraphic sequence boundaries in non-marine basins: an example from the Gunnedah basin, Australia. Amer. Assoc. Petrol. Geol. Bull., 78: 267-286. Hamilton, W.B., 1998. Archean magmatism and deformation were not products of plate tectonics. Precambrian Res., 91: 143-179. Han, T.-M. and Runnegar, B., 1992. Megascopic eukaryotic algae from the 2.1-billion-year-old Negaunee Iron-Formation, Michigan. Science, 257: 232-235. Hanmer, S. and Greene, D.C., 2002. A modern structural regime in the Paleoarchean (~ 3.64 Ga); Isua Greenstone Belt, southern West Greenland. Tectonophysics, 346: 201-222. Hanmer, S., Hamilton, M.A. and Crowley, J.L., 2002. Geochronological constraints on Paleoarchean thrust-nappe and Neoarchean accretionary tectonics in southern West Greenland. Tectonophysics, 350:255-271. Hanmer, S., Williams, M. and Kopf, C., 1995. Striding-Athabasca mylonite zone: implications for the Archean and Early Proterozoic tectonics of the western Canadian Shield. Can. J. Earth Sci., 32: 178-196. Hansen, K.S., 1982. Secular effects of oceanic tidal dissipation on the Moon's orbit and the Earth's rotation. Rev. Geophys. Space Phys., 20: 457-480. Hanski, E.J., Huhma, H., Lehtonen, M.I. and Rastas, E, 1998. 2.0 Ga old oceanic crust in northern Finland. In: E. Hanski and J. Vuollo (Eds.), International Ophiolite Symposium and Field Excursion, Abstracts. Geol. Surv. Finland, Espoo, Spec. Paper, 26, p. 24. Hanson, R.E., Martin, M.W., Bowring, S.A. and Munyanyiwa, H., 1998. U-Pb zircon age for the Umkondo dolerites, eastern Zimbabwe: 1.1 Ga large igneous province in southern Africa-East Antarctica and possible Rodinia correlations. Geology, 26:1143-1146. Haq, B.U., 1991. Sequence stratigraphy, sea-level change, and significance for the deep sea. In: D.I.M. Macdonald (Ed.), Sedimentation, Tectonics and Eustasy. IAS Spec. Publ., 12. Blackwell, Oxford, pp. 3-39. Haq, B.U., Hardenbol, J. and Vail, ER., 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235:1156-1166. Hardebeck, J. and Anderson, D.L., 1996. Eustasy as a test of a Cretaceous superplume hypothesis. Earth Planet. Sci. Lett., 137: 101-108. Hargraves, R.B., 1986. Faster spreading or greater ridge length in the Archaean? Geology, 14: 750-752. Harland, W.B., 1964. Critical evidence for a great Infra-Cambrian glaciation. Geol. Rund., 54: 45-61. Harland, W.B. and Bidgood, D.E.T., 1959. Paleomagnetism in some Norwegian sparagmites and the late-Precambrian ice age. Nature, 184:1860-1862. Harley, S.L., 1989. The origin of granulites: a metamorphic perspective. Geol. Mag., 126: 215-247.

826

References

Harley, S.L., 1992. Proterozoic granulite terranes. In: K.C. Condie (Ed.), Proterozoic Crustal Evolution. Elsevier, Amsterdam, pp. 301-359. Harley, S.L., 1998. An appraisal of peak temperatures and thermal histories in ultrahigh-temperature (UHT) crustal metamorphism: the significance of aluminous orthopyroxene. In: Y. Motoyoshi and K. Shiraishi (Eds.), Origin and Evolution of Continents. Tokyo. Mem. Nat. Inst. of Polar Research. Spec. Issue, 53, pp. 49-73. Harmer, R.E. and von Gruenewaldt, G., 1991. A review of magmatism associated with the Transvaal Basin--implications for its tectonic setting. S. Afr. J. Geol., 94: 104-122. Harnik, P.G. and Simonson, B.M., 2000. Contrasts in spherule textures suggest change in nature of impact strewn fields through geologic time. Abstracts with Programs, Annu. Meeting, Geol. Soc. Amer., Denver, Colorado, 2000, 31, p. 174. Harper, C.L. and Jacobsen, S.B., 1996a. Evidence for ~82Hfin the early solar system and constraints on the timescale of terrestrial accretion and core formation. Geochim. Cosmochim. Acta, 60: 1131-1153. Harper, C.L. and Jacobsen, S.B., 1996b. Noble gases and Earth's accretion. Science, 273:1814-1818. Harper, C.T., 1996. Geology of the Maurice Bay area (NTS 74N-5 and part of 74N-12). Saskatchewan Energy and Mines, Rep. 219, pp. 1-27. Harris, C.W. and Eriksson, K.A., 1990. Allogenic controls on the evolution of storm to tidal shelf sequences in the early Proterozoic Uncompahgre Group, southwest Colorado, U.S.A. Sedimentology, 37: 189-213. Harris, N.B.W., Hawkesworth, C.J. and Tindle, A.G., 1993. The growth of continental crust during the Late Proterozoic: geochemical evidence from the Arabian Shield. In: H.M. Pritchard, T. Alabaster, N.B.W. Harris and C.R. Neary (Eds.), Magmatic Processes and Plate Tectonics. Geol. Soc. London, Spec. Publ., 76, pp. 363-371. Harris, P.D., Moyes, A.B., Fanning, C.M. and Armstrong, R.A., 1995. Zircon ion microprobe resuits from the Maudheim high-grade gneiss terrane, western Dronning Maud Land, Antarctica. Abstract, Centennial Geocongress 1995, Johannesburg, pp. 240--243. Harrison, A.G. and Thode, H.G., 1958. Mechanism of the bacterial reduction of sulphate from isotope fractionation studies. Trans. Faraday Soc., 54: 84-92. Harry, D.L. and Leeman, W.P., 1995. Partial melting of melt-metasomatized subcontinental mantle and the magma source potential of the lower lithosphere. J. Geophys. Res., 100: 10,255-10,270. Hart, M., 1979. Habitable zones around main sequence stars. Icarus, 37: 351-357. Hart, S.R. and Brooks, C., 1977. The geochemistry and evolution of the early Precambrian mantle. Contrib. Mineral. Petrol., 61: 109-128. Hartlaub, R.P., 1999. New insights into the geology of the Murmac Bay Group, Rae Province, northwestern Saskatchewan. Summary of Investigations, Saskatchewan Geological Survey, 1999, 2(99-4.2), pp. 17-26. Hartman, H. and McKay, C.P., 1995. Oxygenic photosynthesis and the oxidation state of Mars. Planet. Space Sci., 43: 123-128. Hartmann, W.K. and Vail, S.M., 1986. Giant impactors: plausible sizes and populations. In: W.K. Hartmann, R.J. Phillips and G.J. Taylor (Eds.), Origin of the Moon. Lunar Planet. Inst., Houston, pp. 551-566. Hartmann, W.K., Ryder, G., Dones, L. and Grinspoon, D., 2001. The time-dependent intense bombardment of the primordial Earth/Moon system. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 493-512. Hartzer, EJ., 1989. Stratigraphy, structure and tectonic evolution of the Crocodile River Fragment. S. Afr. J. Geol., 92:110-124.

References

827

Hassler, S.W. and Simonson, B.M., 2001. The sedimentary record of extraterrestrial impacts in deep shelf environments---evidence from the early Precambrian. J. Geol., 109: 1-22. Hassler, S.W., Robey, H.E and Simonson, B.M., 2000. Bedforms produced by impact-generated tsunami, ~ 2.6 Ga Hamersley basin, Western Australia. Sediment. Geol., 135: 283-294. Hattori, K., Krouse, H.R. and Campbell, EA., 1983. The start of sulfur oxidation in continental environments about 2.2 • 109 years ago. Science, 221: 549-551. Hawkesworth, C.J., Hergt, J.M., McDermott, E and Ellam, R.M., 1991. Destructive margin magmatism and the contributions from the mantle wedge and subducted crust. Aust. J. Earth Sci., 38: 577-594. Hay, R.L., Hildreth, W. and Lambe, R.N., 1979. Globule ignimbrite of Mount Suswa, Kenya. In: C.E. Chapin and W.E. Elston (Eds.), Ash-flow Tufts. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 180, pp. 167-175. Hayes, J.M., 1983. Geochemical evidence bearing on the origin of aerobiosis, a speculative hypothesis. In: J.W. Schopf (Ed.), Earth's Earliest Biosphere: Its Origin and Evolution. Princeton Univ. Press, Princeton, pp. 291-301. Hayes, J.M., 1994. Global methanotrophy at the Archean-Proterozoic transition. In: S. Bengston (Ed.), Early Life on Earth. Nobel Symposium No. 84. Columbia Univ. Press, New York, pp. 220-236. Hayes, J.M., Kaplan, I.R. and Wedeking, K.W., 1983. Precambrian organic geochemistry, preservation of the record. In: J.W. Schopf (Ed.), Earth's Earliest Biosphere, its Origin and Evolution. Princeton Univ. Press, Princeton, New Jersey, pp. 93-134. Head, J.W. and Wilson, L., 2003. Deep submarine pyroclastic eruptions: theory and predicted landforms and deposits. J. Volcanol. Geotherm. Res., 151: 155-193. Heaman, L.M., 1997. Global mafic magmatism at 2.45 Ga: Remnants of an ancient large igneous province? Geology, 25: 299-302. Heiken, G. and Lofgren, G., 1971. Terrestrial glass spheres. Geol. Soc. Amer. Bull., 82: 1045-1050. Hein, EJ., 1984. Deep-sea and fluvial braided channel conglomerates: a comparison of two case studies. In: E.H. Koster and R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Can. Soc. Petrol. Geol., Mem., 10: 33-49. Helland-Hansen, W. and Gjelberg, J.B., 1994. Conceptual basis and variability in sequence stratigraphy: a different perspective. Sediment. Geol., 92:31-52. Helland-Hansen, W. and Martinsen, O.J., 1996. Shoreline trajectories and sequences: description of variable depositional-dip scenarios. J. Sediment. Res., 66(4): 670-688. Hellings, R.W., Adams, P.J., Anderson, J.D., Keesey, M.S., Lau, E.L., Standish, E.M., Canuto, V.M. and Goldman, I., 1983. Experimental test of the variability of G using Viking Lander ranging data. Phys. Rev. Lett., 51: 1609-1612. Helz, G.R., Miller, C.V., Charnock, J.M., Mosselmans, J.EW., Pattrick, R.A.D., Garner, C.D. and Vaughan, D.J., 1996. Mechanisms of molybdenum removal from the sea and its concentration in black shales: EXAFS evidence. Geochim. Cosmochim. Acta, 60:3631-3642. Helz, R.T., 1976. Phase relations of basalts in their melting ranges at P(H20) = 5 kb: II, Melt compositions. J. Petrol., 17: 139-193. Henderson, J., McGrath, P., Theriault, R. and van Breemen, O., 1990. Intracratonic indentation of the Archean Slave Province into the Early Proterozoic Thelon Tectonic Zone of the Churchill Province, northwestern Canadian Shield. Can. J. Earth Sci., 27: 1699-1713. Henderson-Sellers, A., 1979. Clouds and the long-term stability of the Earth's atmosphere and climate. Nature, 279: 786-788. Henderson-Sellers, A., 1988. Equable climate in the early Archaean. Nature, 336:117-118.

828

References

Hendry, H.E. and Wheatley, K.L., 1985. The Carswell Formation, northern Saskatchewan: stratigraphy, sedimentology, and structure. In: R. Lain& D. Alonso and M. Svab (Eds.), The Carswell Structure Uranium Deposits, Saskatchewan. Geol. Assoc. Can. Spec. Paper, 29, pp. 87-103. Herzberg, C., 1999. Phase equilibrium constraints on the formation of cratonic mantle. In Y. Fei, C.M. Bertka and M.O. Mysen (Eds.), Mantle Petrology: Field Observations and High Pressure Experimentation: a Tribute to Francis R. (Joe) Boyd. Geochem. Soc., Spec. Publ. No. 6, pp. 241-257. Herzberg, C.T., 1995. Generation of plume magmas through time: an experimental perspective. Chem. Geol., 126: 1-16. Herzberg, C.T. and O'Hara, M.J., 1998. Phase equilibrium constraints of the origin of basalts, picrites and komatiites. Earth Sci. Rev., 44: 39-79. Herzberg, C.T. and Ohtani, E., 1988. Origin of komatiite at high pressures, Earth Planet. Sci. Lett., 88:321-329. Hess, H., 1962. History of ocean basins. In: A. Engel and B. Leonard (Eds.), Petrologic studies: A Volume in Honor of A. F. Buddington. Geol. Soc. Amer., Denver, pp. 599-620. Heubeck, C. and Lowe, D.R., 1994. Depositional and tectonic setting of the Archean Moodies Group, Barberton Greenstone Belt, South Africa. Precambrian Res., 68: 257-290. Heymann, D., Dressier, B.O. and Thiemens, M.H., 1997a. Origin of native sulfur in rocks from the Sudbury Structure, Ontario, Canada. LPI Contribution, 922: 22-23. Heymann, D., Yancey, T.E. and Thiemens, M.H., 1997b. Native sulfur in sediments from KT boundary sites of the Brazos River, Texas. Abstracts of Papers Submitted to the Lunar and Planetary Science Conference, Houston, Texas, 1997, 28, pp. 565-566. Hickman, A.H., 1975. Precambrian structural geology of part of the Pilbara region. West. Aust. Geol. Surv., Annu. Rep., 1974, pp. 68-73. Hickman, A.H., 1983. Geology of the Pilbara Block and its environs. Geol. Surv. W. Aust. Bull., 127, 268 pp. Hickman, A.H., 1984. Archaean diapirism in the Pilbara Block, Western Australia. In: A. Kr6ner and R. Greiling (Eds.), Precambrian Tectonics Illustrated. E. Schweizerbarts'che Verlagsbuchhandlung, Stuttgart, pp. 113-127. Hickman, A.H., 1990. Geology of the Pilbara Craton. In: S.E. Ho, J.E. Glover, J.S. Myers and J.R. Muhling (Eds.), Excursion Guidebook No. 5, Pilbara and Hamersley Basin. Third Intern. Archaean Syrup., Perth, 1990, Univ. Western Aust. Geol. Dept. and Univ. Extension, Publ. 21, pp. 1-13. Hickman, A.H., 2001. East Pilbara diapirism: new evidence from mapping. West. Aust. Geol. Surv., Record 2001/5, pp. 23-25. Hickman, A.H., in press. Two contrasting granite-greenstone terranes in the Pilbara Craton, Australia: evidence for vertical and horizontal tectonic regimes prior to 2900 Ma. Precambrian Res. Higgins, M.D. and van Breeman, O., 1998. The age of the Sept lies layered mafic intrusion, Canada: Implications for the late Neoproterozoic/Cambrian history of southeastern Canada. J. Geol., 106: 421-433. Hildebrand, R.S., 1981. Early Proterozoic Labine Group of Wopmay Orogen: Remnant of a continental volcanic arc developed during oblique convergence. In: EH.A. Campbell (Ed.), Proterozoic Basins of Canada. Geol. Surv. Can. Paper, 81 (10), pp. 133-156. Hill, A.C. and Walter, M.R., 2000. Mid-Neoproterozoic ('~ 830-750 Ma) isotope stratigraphy of Australia and global correlation. Precambrian Res., 100:181-211. Hill, E.J. and Baldwin, S.L., 1993. Exhumation of high-pressure metamorphic rocks during crustal extension in the D'Entrecasteaux region, Papua New Guinea. J. Metamorphic Geol., 11: 261-277.

References

829

Hill, E.J., Baldwin, S.L. and Lister, G.S., 1992. Unroofing of active metamorphic core complexes in the D'Entrecasteaux Islands, Papua New Guinea. Geology, 20:907-910. Hill, R.E.T., Barnes, S.-J., Gole, M.J. and Dowling, S.E., 1995. The volcanology of komatiites as deduced from field relationships in the Norseman-Wiluna greenstone belt, Western Australia. Lithos, 34: 159-188. Hill, R.I., Campbell, I.H., Davies, G.E and Griffiths, R.W., 1992. Mantle plumes and continental tectonics. Science, 256: 186-193. Hinderer, J. and Legros, H., 1988. Tidal flow in the Earth's core: a search for the epoch and amplitude of exact resonance in the past. Geophys. Monogr., 46: 79-82. Hinrichs, K.-U., 2002. Microbial fixation of methane carbon at 2.7 Ga: Was an anaerobic mechanism possible? Geochem. Geophys. Geosyst., 3, Paper No. 2001GC000286. Hinrichs, K.-U., Hayes, J.M., Sylva, S.P., Brewer, P.G. and DeLong, E.E, 1999. Methane-consuming archaebacteria in marine sediments. Nature, 398: 802-805. Hippert, J. and Davis, B.: 2000. Dome emplacement and formation of kilometre-scale synclines in a granite-greenstone terrain (Quadrilatero Ferrifero, southeastern Brazil). Precambrian Res., 102: 99-121. Hirdes, W. and Davis, D.W., 2002. U-Pb geochronology of Paleoproterozoic rocks in the southern part of the Kedougou-K6ni6ba inlier, Senegal, West Africa: evidence for diachronous accretionary development of the Eburnean province. Precambrian Res., 118: 83-99. Hirdes, H., Davis, D.W. and Eisenlohr, B.N., 1992. Reassessment of Proterozoic granitoid ages in Ghana on the basis of U-Pb zircon and monazite dating. Precambrian Res., 56: 89-96. Hirdes, W., Davis, D.W., LiJdtke, G. and Konan, G., 1996. Two generations of Birimian (Paleoproterozoic) volcanic belts in northeastern C6te d'Ivoire (West Africa): consequences for the "Birimian controversy". Precambrian Res., 80:173-191. Hjellbakk, A., 1993. A "flash-flood" dominated braid delta in the upper Proterozoic Na~ringselva Member, Varanger Peninsula, northern Norway. Norsk Geol. Tidsskr., 73: 63-80. Hjellbakk, A., 1997. Facies and fluvial architecture of a high-energy braided river: the upper Proterozoic Selodden Member, Varanger Peninsula, northern Norway. Sediment. Geol., 114:131-161. Hoal, B.G. and Heaman, L.M., 1995. The Sinclear Sequence: U-Pb age constraints from the Awasib Mountain area. Communications of the Geological Survey of Namibia, 10:83-91. Hobson, G.D. and MacAuley, H.A., 1969. A seismic reconnaissance survey of the Athabasca Formation, Alberta and Saskatchewan. Geol. Surv. Can. Paper, 69(18), pp. 1-23. Hodgson, C.J. and Hamilton, J.V., 1989. Gold mineralization in the Abitibi greenstone belt: end-stage result of Archean collisional tectonics? Econ. Geol. Mono., 6: 86-100. Hoering, T.C., 1967. The organic geochemistry of Precambrian rocks. In: P.H. Abelson (Ed.), Researches in Geochemistry, Vol. 2. John Wiley and Sons, New York, pp. 87-111. Hoeve, J., Cumming, G.L., Baadsgaard, H. and Morton, R.D., 1981. Geochronology of uranium metallogenesis in Saskatchewan. Canadian Institute of Mining Uranium Symposium, September 8-13, 1981, Saskatoon, Technical Program, pp. 11-12. Hoffman, P.E, 1976. Stromatolite morphogenesis in Shark Bay, Western Australia. In: M.R. Walter (Ed.), Stromatolites. Elsevier, New York, Developments in Sedimentology, 20, pp. 261-271. Hoffman, P.E, 1987. Proterozoic foredeeps, foredeep magmatism, and superior-type iron-formations of the Canadian Shield. In: A. Kroner (Ed.), Proterozoic Lithospheric Evolution. Amer. Geophys. Union Geodynamics Ser., 17, pp. 85-98. Hoffman, P.E, 1988. United plates of America, the birth of a craton: Early Proterozoic assembly and growth of Laurentia. Annu. Rev. Earth Planet. Sci., 16: 543-603.

830

References

Hoffman, RE, 1989a. Speculations on Laurentia's first gigayear (2.0 to 1.0 Ga). Geology, 17: 135-138. Hoffman, EE, 1989b. Do contemporaneous 1.7 Ga basins signify that Laurentia aggregated above a mantle downwelling? Abstracts, Geol. Assoc. Can., 14: 74. Hoffman, EE, 1989c. Precambrian geology and tectonic history of North America. In: A.W. Bally and A.R. Palmer (Eds.), The Geology of North America, Vol. A. Geol. Soc. Amer., Boulder, Colorado, pp. 447-512. Hoffman, EE, 1990. Subdivision of the Churchill Province and extent of the Trans-Hudson Orogen, In: J.E Lewry and M.R. Stauffer (Eds.), The Early Proterozoic Trans-Hudson Orogen of North America. Geol. Assoc. Can. Spec. Paper, 37, pp. 15-39. Hoffman, EE, 1991. Did the breakout of Laurentia turn Gondwanaland inside-out? Science, 252: 1409-1412. Hoffman, EE, 1992. Supercontinents. In: Encyclopedia of Earth System Science. Academic Press, London, 4, pp. 323-328. Hoffman, EE, 1999a. "Snowball Earth" theory still stands. Nature, 400: 708. Hoffman, EE, 1999b. The break-up of Rodinia, birth of Gondwana, true polar wander and the snowball Earth. J. Afr. Earth Sci., 28:17-33. Hoffman, EE, 2000. Discussion: Vreeland diamictites--Neoproterozoic glaciogenic slope deposits, Rocky Mountains, southeast British Columbia. Bull. Can. Petrol. Geol., 48: 360-363. Hoffman, EE, 2001. Snowball Earth: promise and problems. Earth System Processes, Geol. Soc. Amer. Geol. Soc. London Meeting, Edinburgh, Programme with Abstracts, p. 109. Hoffman, EE and Ranalli, G., 1988. Archaean oceanic flake tectonics. Geophysic. Res. Lett., 15: 1077-1080. Hoffman, EE and Schrag, D.E, 2002. The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14: 129-155. Hoffman, EE, Hawkins, D.P., Isachsen, C.E. and Bowring, S.A., 1996. Precise U-Pb zircon ages for early Damaran magmatism in the Summas Mountains and Welwitschia inlier, northern Damara belt, Namibia. Communications of the Geological Survey of Namibia, 11: 47-52. Hoffman, EE, Kaufman, A.J. and Halverson, G.E, 1998a. Comings and goings of global glaciations on a Neoproterozoic tropical platform in Namibia. GSA Today, 8(5): 1-9. Hoffman, EE, Kaufman, A.J., Halverson, G.E and Schrag, D.E, 1998b. A Neoproterozoic Snowball Earth. Sci., 281: 1342-1346. Hoffmann, K.H., 1989. New aspects of lithostratigraphic subdivision and correlation of late Proterozoic to early Cambrian rocks of the southern Damara Belt and their correlation with the central and northern Damara Belt and the Gariep Belt. Communications of the Geological Survey of Namibia, 5: 59-67. Hoffmann, K.H. and Prave, A.R., 1996. A preliminary note on a revised subdivision and regional correlation of the Otavi Group based on glaciogenic diamictites and associated cap dolostones. Communications of the Geological Survey of Namibia, 11: 77-82. Hofmann, A., Dirks, E and Jelsma, H., 2001. Late Archaean foreland basin evolution, Belingwe greenstone belt, Zimbabwe. Sediment. Geol., 141/142:131-168. Hofmann, A.W., 1988. Chemical differentiation of the Earth: the relationships between mantle, continental crust, and oceanic crust. Earth Planet. Sci. Lett., 90:297-314. Hofmann, H.J., 1973. Stromatolites: characteristics and utility. Earth. Sci. Rev., 9: 339-373. Hofmann, H.J., 1976. Precambrian microflora, Belcher Islands, Canada: significance and systematics. J. Paleont., 50:1040-1073.

References

831

Hofmann, H.J., 1998. Synopsis of Precambrian fossil occurrences in North America. Chapter 4, Geol. Soc. America, The Geology of North America, Vol. C-1, pp. 273-293. Hofmann, H.J., 2000. Archean stromatolites as microbial archives. In: R.E. Riding and S.M. Awramik (Eds.), Microbial Sediments. Springer-Verlag, Berlin, pp. 315-327. Hofmann, H.J. and Masson, M., 1994. Archean stromatolites from Abitibi greenstone belt, Quebec, Canada. Geol. Soc. Amer. Bull., 106: 424-429. Hofmann, H.J., Grey, K.A.H. and Thorpe, R.I., 1999. Origin of 3.45 Ga coniform stromatolites in Warrawoona Group, Western Australia. Geol. Soc. Amer. Bull., 111:1256-1262. Holland, H.D., 1964. On the chemical evolution of the terrestrial and Cytherean atmosphere. In: P.J. Brancazio and A.G.W. Cameron (Eds.), The Origin and Evolution of Atmospheres and Oceans, John Wiley and Sons, Chichester, pp. 86-97. Holland, H.D., 1966. Origin and evolution of the atmosphere. McGraw-Hill 1966 Yearbook, Science and Technology, pp. 53-59. Holland, H.D., 1973. The oceans: a possible source of iron in iron-formations. Econ. Geol., 68: 1169-1172. Holland, H.D., 1978. The Chemistry of the Atmosphere and Oceans. Princeton Univ. Press, Princeton, New Jersey, 351 pp. Holland, H.D., 1984. The Chemical Evolution of the Atmosphere and the Oceans. Princeton Univ. Press, Princeton, New Jersey, 582 pp. Holland, H.D., 1994. Early Proterozoic atmospheric change. In: S. Bengston (Ed.), Early Life on Earth. Nobel Symposium No. 84. Columbia Univ. Press, New York, pp. 237-244. Holland, H.D., 1999. When did the atmosphere become oxic? A reply. Geochemical News, 100: 20-22. Holland, H.D., 2002. Volcanic gases, black smokers, and the great oxidation event. Geochim. Cosmochim. Acta, 66:3811-3826. Holland, H.D. and Beukes, N.J., 1990. A paleoweathering profile from Griqualand West, South Africa: evidence for a dramatic rise in atmospheric oxygen between 2.2 and 1.9 by bp. Amer. J. Sci., 290A: 1-34. Hollings, P. and Ansdell, K., 2002. Paleoproterozoic arc magmatism imposed on an older backarc basin: implications for the tectonic evolution of the Trans-Hudson orogen, Canada. Geol. Soc. Amer. Bull., 114: 153-168. Hollings, P., Wyman, D. and Kerrich, R., 1999. Komatiite-basalt-rhyolite volcanic associations in northern Superior Province greenstone belts: significance of plume-arc interaction in the generation of the proto-continental Superior Province. Lithos, 46: 137-161. Holser, W.T. and Margaritz, M., 1987. Events near the Permian-Triassic boundary. Mod. Geol., 11: 155-180. Hrltt~i, P., 1988. Metamorphic zones and evolution of granulite grade metamorphism in the early Proterozoic Pielavesi area, central Finland. Geol. Surv. Finland, Bull., 344, 50 pp. Holzer, L., Frei, R., Barton, J.M., Jr. and Kramers, J.D., 1998. Unravelling the record of successive high grade events in the Central Zone of the Limpopo belt using Pb single phase dating of metamorphic minerals. Precambrian Res., 87:87-115. Hon, K., Heliker, C. and Kjargaard, J.I., 1998. Limu o Pele: a new kind of hydroclastic tephra from Kilauea Volcano, Hawaii. Geol. Soc. Amer., Abstracts with Programs, 20:112-113. Hon, K., Kauahikaua, J., Denlinger, R. and McKay, K., 1994. Emplacement and inflation of pahoehoe sheet flows: observations and measurement of active lava flows on Kilauea Volcano, Hawaii. Geol. Soc. Amer. Bull., 106: 351-370.

832

References

Honda, M., McDougall, I., Patterson. D., Doulgeris, A. and Clague, D.A., 1991. Possible solar noblegas component in Hawaiian basalts. Nature, 349: 149. Honda, S., Yoshida, M., Ootorii, S. and lwase, Y., 2000. The timescales of plume generation caused by continental aggregation. Earth Planet. Sci. Lett., 176:31-43. Honjo, S., Francois, R., Krishfield, R. and Manganini, S., 2002. Global model of export flux and roles of CaCO3 particles in removing atmospheric COe to the oceanic sinks. Abstract, 6th Intern. Symp. Geochem. Earth's Surface, May 20-24, 2002, Honolulu, Hawaii, pp. 124-125. Hooper, P.R., 1997. The Columbia River Flood Basalt province: Current Status. In: J.J. Mahoney and M.E Coffin (Eds.), Large Igneous Provinces: Continental, Oceanic and Planetary Flood Volcanism. Amer. Geophys. Union, Geophys., Monograph 100, pp. 1-27. Hopkins, R.E, Eriksson, K.A. and Simpson, E.L., 2000. Annual cycles in tidal rhythmites: the ~ 2.9 Ga Coronation Shale, West Rand Group, South Africa. Geol. Soc. Amer., Abstracts with Programs, 32: A307. Horan, M.E, Smollar, M.I. and Walker, R.J., 1998. 182W and 187Re --187Os systematics of iron meteorites: chronology for melting, differentiation, and crystallization in asteroids. Geochim. Cosmochim. Acta., 62: 545-554. Horneck, G. and Baumstark-Khan, C. (Eds.), 2001. Astrobiology: the Quest for the Conditions of Life. Springer-Verlag, Berlin, 411 pp. Horodyski, R.J., 1977a. Environmental influences on columnar stromatolite branching patterns: examples from the Middle Proterozoic Belt Supergroup, Glacier National Park, Montana. J. Paleontol., 51: 661-671. Horodyski, R.J., 1977b. Lyngbya Mats at Laguna Mormona, Baja California, Mexico: comparison with Proterozoic stromatolites. J. Sediment. Petrol., 47: 1305-1320. Horodyski, R.J., 1982. Impressions of algal mats from the Middle Proterozoic Belt Supergroup, northwestern Montana, U.S.A. Sedimentology, 29: 285-289. Horodyski, R.J., 1993. Paleontology of Proterozoic shales and mudstones: examples from the Belt Supergroup, Chuar Group and Pahrump Group, western USA. Precambrian Res., 61: 241-278. Horodyski, R.J. and Knauth, L.P., 1994. Life on land in the Precambrian. Science, 263: 494-498. Horodyski, R.J., Bloeser, B. and Von der Haar, S., 1977. Laminated algal mats from a coastal lagoon, Laguna Mormona, Baja California, Mexico. J. Sediment. Petrol., 47: 680-696. Horstwood, M.S.A., Nesbitt, R.W., Noble, S.R. and Wilson, J.E, 1999. U-Pb zircon evidence for an extensive early Archean craton in Zimbabwe: a reassessment of the timing of craton formation, stabilization, and growth. Geology, 27(8): 707-710. Horwitz, R.C., 1982. Geological history of the early Proterozoic Paraburdoo Hinge Zone, Western Australia. Precambrian Res., 19: 191-200. Hough, R.M., Gilmour, I., Pillinger, C.T., Arden, J.W., Gilkes, K.W.R., Yuan, J. and Milledge, H.J., 1995. Diamond and silicon carbide in impact melt rock from the Ries impact crater. Nature, 378: 41-44. Houghton, B.F., Wilson, C.J.N., McWilliams, M.O., Lanphere, M.A., Weaver, S.D., Briggs, R.M. and Pringle, M.S., 1995. Chronology and dynamics of a large silicic magmatic system: central Taupo Volcanic Zone, New Zealand. Geology, 23: 13-16. Houl6, M., Lesher, C.M., Gibson, H.L., Fowler, A.D. and Sproule, R.A., 2001. Project unit 99-021. Physical volcanology of komatiites in the Abitibi greenstone belt. In: J. Ayer, C. Baker, R. Kelly and J. Parker (Eds.), Summary of Field Work and Other Activities, Open File Rep. 6070, Ontario Geol. Surv. 13, 17 pp. House, C.H., Schopf, J.W., McKeegan, K.D., Coath, C.D., Harrison, T.M. and Stetter, K.O., 2000. Carbon isotopic composition of individual Precambrian microfossils. Geology, 28: 707-710.

References

833

Houseman, G. and Molnar, E, 1997. Gravitational (Rayleigh-Taylor) instability of a layer with non-linear viscosity and convective thinning of continental lithosphere. Geophys. J. Int., 128: 125-150. Houston, R.S., Karlstrom, K.E., Graft, EJ. and Flurkey, A.J., 1992. New stratigraphic subdivisions and redefinition of subdivisions of Late Archean and early Proterozoic metasedimentary and metavolcanic rocks of the Sierra Madre and Medicine Bow Mountains, Southern Wyoming, U.S. U.S. Geol. Surv. Prof. Paper, 1520. Government Printing Office, Washington, 50 pp. Hsua, W., Wasserburg, G.J. and Huss, G.R., 2000. High time resolution by use of the 26A1chronometer in the multistage formation of a CAI. Earth Planet. Sci. Lett., 182: 15-29. Hudak, G.J., Morton, R.L., Franklin, J.M. and Peterson, D.M., 2002a. Morphology, distribution, and estimated eruption volumes for intracaldera felsic tuff deposits in the Archean Sturgeon Lake subaqueous caldera complex, Northwestern Ontario. Abstract, AGU-Chapman conference, Explosive Subaqueous Volcanism, 21-25th January 2002, Dunedin, New Zealand, p. 30. Hudak, G.J., Morton, R.L., Franklin, J.M. and Peterson, D.M., 2002b. Morphology, distribution, and estimated eruption volumes for intracaldera tufts associated with volcanic-hosted massive sulfide deposits in the Archean Sturgeon Lake Caldera Complex, northwestern Ontario. In: J. White, D. Clague and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph (in press). Huddleston, J., 1976. Early deformational history of Archaean rocks in the Vermillion District, northeastern Minnesota. Can. J. Earth Sci., 13: 579-592. Hudleston, EJ. and Schwerdtner, W.M., 1997. Strain. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Sci. Publ., Oxford, pp. 296-308. Hughes, C.J., 1977. Archean variolites---quenched immiscible liquids: discussion. Can. J. Earth Sci., 14: 137-139. Hughes, D.W., 1998. The mass distribution of crater producing bodies. In: M.M. Grady, R. Hutchinson, G.J.H. McCall and D.A. Rothery (Eds.), Meteorites: Flux with Time and Impact Effects. Geol. Soc. London Spec. Vol., 140, pp. 31-42. Huhma, H. and Meril~iinen, K., 1991. Provenance of paragneisses from the Lapland Granulite Belt. Res Terrae, Ser. A, Univ. Oulu, Finland, 5: 26. Humphreys, H.C. and Van Bever Donker, J.M., 1987. Aspects of deformation along the Namaqua Province eastern boundary, Kenhardt District, South Africa. Precambrian Res., 46: 39-63. Hunt, B.G., 1982. The impact of large variations of the Earth's obliquity on the climate. J. Meteorol. Soc. Japan, 60:309-318. Hunt, D. and Tucker, M.E., 1992. Stranded parasequences and the forced regressive wedge systems tract: deposition during base-level fall. Sediment. Geol., 81: 1-9. Hunt, D. and Tucker, M.E., 1995. Stranded parasequences and the forced regressive wedge systems tract: deposition during base-level fall-reply. Sediment. Geol., 95: 147-160. Hunten, D.M., 1993. Atmospheric evolution of the terrestrial planets. Science, 259: 915-920. Hunter, M.A., Bickle, M.J., Nisbet, E.G., Martin, A. and Chapman, H.J., 1998. Continental extensional setting for the Archaean Belingwe Greenstone Belt, Zimbabwe. Geology, 26: 883-886. Hunter, R.E., 1977. Basic types of stratification in small aeolian dunes. Sedimentology, 24:361-387. Hunter, R.E., 1981. Stratification styles in eolian sandstones: Some Pennsylvanian to Jurassic examples from the western interior U.S.A. In: EG. Etheridge and R.M. Flores (Eds.), Recent and Ancient Nonmarine Depositional Environments: Models for Exploration. Soc. Econ. Paleont. Mineral., Norman, Oklahoma, Spec. Publ., 31, pp. 315-329. Huppert, H.E. and Sparks, R.S.J., 1985a. Komatiites I: eruption and flow. J. Petrol., 26: 694-725.

834

References

Huppert, H.E. and Sparks, R.S.J., 1985b. Cooling and contamination of mafic and ultramafic magmas during ascent through continental crust. Earth Planet. Sci. Lett., 74:371-386. Hurtgen, M.T., Arthur, M.A., Suits, N. and Kaufman, A.J., 2002. The sulfur isotopic composition of Neoproterozoic seawater sulfate: Implications for a snowball Earth? Earth Planet. Sci. Lett., 203: 413-429. Huston, D.L., Sun, S.-S., Blewett, R., Hickman, A., Van Kranendonk, M., Phillips, D., Baker, D. and Brauhart, C.A., 2002. The timing of mineralization in the Archaean North Pilbara terrain, Western Australia. Econ. Geol., 97: 733-755. Hutcheon, I.D. and Hutchison, R.H., 1989. Evidence from the Semarkona ordinary chondrite for the 26A1 heating of small planets. Nature, 337:238-241. Hutcheon, I.D., Krot, A.N., Kell, K., Phiney, D.L. and Scott, E.R.D., 1998. 53Mn-S3Cr dating of fayalite formation in the CV3 chondrite Mokoia: evidence for asteroidal alteration. Science, 282: 1865-1867. Hutchison, R., Williams, I.P. and Russell, S.S., 2001. Theories of planetary formation: constraints from the study of meteorites. Phil. Trans. R. Soc. London A, 359: 2077-2093. Hutton, D.H.W., 1996. The "space problem" in the emplacement of granite. Episodes, 19:114-119. Hyde, R.S., 1980. Sedimentary facies in the Archean Timiskaming Group and their tectonic implications, Abitibi greenstone belt, northeastern Ontario, Canada. Precambrian Res., 12: 161-195. Hynes, A. and Francis, D.M., 1982. A transect of the Early Proterozoic Cape Smith foldbelt, New Quebec. Tectonophysics, 88: 23-59. Idnurm, M. and Giddings, J.W., 1988a. Australian Precambrian polar wander: a review. Precambrian Res., 40/41:61-88. Idnurm, M. and Giddings, J.W., 1998b. Paleoproterozoic-Neoproterozoic North America-Australia link: new evidence from paleomagnetism. Geology, 23: 149-152. Imai, A. and Geshi, N., 1998. Spinifex texture of native sulfur; a cooling product of sulfur flow eruptions at Shiretoko-lwozan Volcano, Hokkaido, Japan. Res. Geol., 49: 99-104. Ireland, T.R., Rudnick, R.L. and Spetsius, Z., 1994. Trace elements in diamond inclusions from eclogites reveal link to Archaean granites. Earth Planet. Sci. Lett., 128:199-213. Irwin, M.L., 1965. General theory of epeiric clear water sedimentation. Bull. Amer. Assoc. Petrol. Geol., 49: 445-459. Ishibashi, J. and Urabe, T., 1995. Hydrothermal activity related to arc-backarc magmatism in the western Pacific. In: B. Taylor (Ed.), Backarc Basins, Tectonism and Magmatism. Plenum Press, New York, pp. 451-495. Isley, A.E., 1995. Hydrothermal plumes and the delivery of iron to banded iron formation. J. Geol., 103: 161-185. Isley, A.E. and Abbott, D.H., 1999. Plume-related mafic volcanism and the deposition of banded iron formation. J. Geophys. Res., B, Solid Earth and Planets, 104: 15,461-15,477. Isley, A.E. and Abbott, D.H., 2002. Implications of the temporal distribution of high-Mg magmas for mantle plume volcanism through time. J. Geology, 110: 141-158. Ito, T., Masuda, K., Hamano, Y. and Matsui, T., 1995. Climate friction: a possible cause for secular drift of Earth's obliquity. J. Geophys. Res., 100: 15,147-15,161. Jackson, C., 1999. Characterization of Mesoproterozoic to Palaeozoic crustal evolution of western Dronning Maud Land. Study 3: Deformational history and thermochronology of the central Kirwanveggen. Unpubl. report, Department of Environmental Affairs and Tourism, Pretoria, 80 pp. Jackson, G.D. and Berman, R.G., 2000. Precambrian metamorphic and tectonic evolution of northern Baffin Island, Ninavut, Canada. Can. Mineral., 38: 399-421.

References

835

Jackson, J.M., Simpson, E.L. and Eriksson, K.A., 1990. Facies and sequence stratigraphic analysis in an intracratonic, thermal-relaxation basin: the middle Proterozoic, lower Quilalar Formation, Mount Isa Orogen, Australia. Sedimentology, 37: 1053-1078. Jackson, M.P.A., Cornelius, R.R., Craig, C.H., Gansser, A., Stocklin, J. and Talbot, C.J., 1990. Salt diapirs of the Great Kavir, Central Iran. Geol. Soc. Amer. Mem., 177: 139. Jackson, R.G. II, 1978. Preliminary evaluation of lithofacies models for meandering alluvial streams. In: A.D. Miall (Ed.), Fluvial Sedimentology. Can. Soc. Petrol. Geol., Calgary, Alberta, Mem., 5, pp. 543-576. Jackson, T.A., 1965. Power-spectrum analysis of two "varved" argillites in the Huronian Cobalt Series (Precambrian) of Canada. J. Sediment. Petrol., 35: 877-886. Jacob, D. and Foley, S.E, 1999. Evidence for Archaean ocean crust with low high field strength element signature from diamondiferous eclogite xenoliths. Lithos, 48: 317-336. Jacobs, J., Thomas, R.J. and Weber, K., 1993. Accretion and indentation tectonics along the southern margin of the Kaapvaal Craton during the Kibaran (Grenville) Orogeny. Geology, 21: 203-206. Jacobs, J., Falter, M., Thomas, R.J., Kunz, J. and Jel3berger, E.K., 1997.40Ar/39Ar thermochronological constraints on the structural evolution of the Mesoproterozoic Natal Metamorphic Province, SE Africa. Precambrian Res., 86: 71-92. Jacobs, J., Fanning, C.M., Henjes-Kunst, E, Olesch, M. and Paech, H.-S., 1998. Continuation of the Mozambique Belt into East Antarctica: Grenville-age metamorphism and polyphase Pan-African high-grade events in central Dronning Maud Land. J. Geol., 106: 385-406. Jacobs, J., Hansen, B.T., Henjes-Kunst, E, Thomas, R.J., Bauer, W., Weber, K., Armstrong, R.A. and Cornell, D.H., 1999. New age constraints on the Proterozoic/Lower Palaeozoic evolution of Heimefrontfjella, East Antarctica, and its bearing on Rodinia/Gondwana correlations. Terra Antarctica, 6: 377-389. Jacobs, J., Thomas, R.J., Fanning, C.M., Bauer, W. and Klemd, R., 2002. Collision, collapse and differential foreland exhumation of the East African/Antarctic Orogen. Abstract, 11th Quadrennial IAGOD Syposium and Geocongress 2002, Windhoek, July 22-26, 2002 (CD-ROM). Jacobsen, S.B., 1988. Isotopic constraints on crustal growth and recycling. Geochim. Cosmochim. Acta, 52: 1341-1350. Jacobsen, S.B. and Kaufman, A.J., 1999. The Sr, C and O isotopic evolution of Neoproterozoic seawater. Chem. Geol., 161: 37-57. Jacobsen, S.B. and Pimentel-Klose, M.R., 1988. A Nd isotopic study of the Hamersley and Michipicoten banded iron formations: the source of REE and Fe in Archaean oceans. Earth Plan. Sci. Lett., 87: 856-867. Jaeckel, E, Kr6ner, A., Kamo, S.L., Brandl, G. and Wendt, J.I., 1997. Late Archaean to early Proterozoic granitoid magmatism and high-grade metamorphism in the central Limpopo belt, South Africa. J. Geol. Soc. London, 154: 25--44. Jahn, B.M., Gruau, G. and Glickson, A.Y., 1982. Komatiites of the Onverwacht Group, South Africa: REE chemistry, Sm-Nd age and mantle evolution. Contrib. Mineral. Petrol., 80: 25-40. Jahnke, L. and Klein, H.E, 1979. Effects of low levels of oxygen on Sacchroyces cerevisial. Origins of Life 9: 329-334. Jahren, A.H., 2002. The biochemical consequences of the Mid-Cretaceous superplume. J. Geodynam., 34: 177-191. Jakosky, B.M., 1998. The Search for Life on Other Planets. Cambridge Univ. Press, Cambridge, 326 pp. Jakosky, B.M., 1999. Atmospheres of terrestrial planets. In: J.K. Beattie, C.C. Petersen and A. Chaikin (Eds.), The New Solar System. Cambridge Univ. Press, Cambridge, pp. 175-192.

836

References

Jakosky, B.M. and Phillips, R.J., 2001. Mars' volatile and climate history. Nature, 412: 237-244. Jakosky, B.M. and Shock, E.L., 1998. The biological potential of Mars, the early Earth, and Europa. J. Geophys. Res., 103: 19359-19364. Jakosky, B.M., Henderson, B.G. and Mellon, M.T., 1995. Chaotic obliquity and the nature of the Martian climate. J. Geophys. Res., 100: 1579-1584. James, H.L., 1954. Sedimentary facies of iron-formation. Econ. Geol., 49: 235-293. James, H.L., 1958. Stratigraphy of pre-Keweenawan rocks in parts of northern Michigan. U.S. Geol. Surv. Prof. Paper, 314-C: C22-C24. James, H.L., 1966. Chemistry of the iron-rich sedimentary rocks. In: M. Fleischer (Ed.), Data of Geochemistry, 6th edn. U.S. Geol. Surv. Prof. Paper, 440W, 61 pp. James, H.L., 1969. Comparison between Red Sea Deposits and older ironstone and iron formation. In: E.T. Degans and D.A. Ross (Eds.), Hot Brines and Recent Heavy Metal Deposits in the Red Sea. Springer-Verlag, New York, pp. 525-532. James, H.L., 1983. Distribution of banded iron-formation in space and time. In: A.E Trendall and R.C. Morris (Eds.), Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 471-490. James, H.L. and Sims, P.K. (Eds.), 1973. Precambrian iron-formations of the world. Econ. Geol., 68: 913-1179. James, H.L. and Trendall, A.E, 1982. Banded iron formation: distribution in time and paleoenvironmental significance. In: H.D. Holland and M. Schidlowski (Eds.), Mineral Deposits and the Evolution of the Biosphere. Springer-Verlag, New York. pp. 199-218. Jefferson, C.W., Delaney, G. and Olson, R.A., 2002. EXTECH IV Athabasca uranium multidisciplinary study: mid-year 2002-03 overview. Summary of Investigations 2002, Vol. 2, Saskatchewan Geological Survey, Saskatchewan Industry and Resources, Miscellaneous Report 2002-4.2 (CD-ROM). Jeffery, P.M. and Reynolds, J.H., 1961. Origin of excess 129Xe in stone meteorites. J. Geophys. Res., 66: 3582-3583. Jelsma, H.A., Van der Beek, P.A. and Vinyu, M.L., 1993. Tectonic evolution of the Bindura-Shamva greenstone belt (northern Zimbabwe): progressive deformation around diapiric batholiths. J. Struct. Geol., 15: 163-176. Jenkins, G.S., 2000. Global climate model high-obliquity solutions to the ancient climate puzzles of the Faint-Young Sun Paradox and low-latitude Proterozoic glaciation. J. Geophys. Res., 105: 7357-7370. Jenkins, G.S., 2001. High-obliquity simulations for the Archean Earth: implications for climatic conditions on early Mars. J. Geophys. Res., 106: 32,903-32,913. Jenkyns, H.C., 1980. Cretaceous anoxic events: from continents to oceans. J. Geol. Soc. London, 137: 171-188. Jensen, S., Saylor, B.Z., Gehling, J.G. and Germs, G.J.B., 2000. Complex trace fossils from the terminal Proterozoic of Namibia. Geology, 28: 143-146. Jephcoat, A.P., 1998. Rare-gas solids in the Earth's deep interior. Nature, 393: 355-358. Jervey, M.T., 1988. Quantitative geological modeling of siliciclastic rock sequences and their seismic expression. In: C.K. Wilgus, B.S. Hastings, C.G. St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Eds.), Sea Level Changes--An Integrated Approach. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 42, pp. 47-69. Jirsa, M.A., Southwick, D.L. and Boerboom, T.J., 1992. Structural evolution of Archean rocks in the Western Wawa subprovince, Minnesota: refolding of pre-cleavage nappes during D2 transpression. Can. J. Earth Sci., 29: 2146-2155.

References

837

Jochum, K.R, Amdt, N.T. and Hofmann, A.W., 1990. Nb-Th-La in komatiites and basalts: constraints on komatiite petrogenesis and mantle evolution. Earth Planet. Sci. Lett., 107: 272-289. Jochum, K.R, Polat, A., Stoll, B. and Hofmann, A.W., 2001. Low Nb/Ta in the Archean mantle: Implications for the missing Niobium in the silicate Earth. Eos Trans., 82: 1214. Johnson, H.D., 1977. Shallow marine sand bar sequences: an example from the late Precambrian of North Norway. Sedimentology, 24: 245-270. Johnson, N.M. and Fegley, B., 2000. Water on Venus: new insights from tremolite decomposition. Icarus, 146:301-306. Johnson, S.Y., 1993. Ice wedge casts in the Daltradian of South Donegal--evidence for subaerial exposure of the Boulder Bed. Irish J. Earth Sci., 12: 13-26. Johnsson, M.J., 1993. The system controlling the composition of clastic sediments. In: M.J. Johnsson and A. Basu (Eds.), Processes Controlling the Composition of Clastic Sediments. Geol. Soc. Amer., Spec. Paper, 284, pp. 1-19. Johnsson, M.J. and Stallard, R.E, 1989. Physiographic controls on the composition of sediments derived from volcanic and sedimentary terrains on Barro Colorado Island, Panama. J. Sediment. Petrol., 59:768-781. Johnston, S.T., Armstrong, R., Heaman, L., McCourt, S., Mitchell, A., Bisnath, A. and Arima, M., 2001. Preliminary U-Pb geochronology of the Tugela terrane, Natal belt, eastern South Africa. Memoirs of National Institute of Polar Research, 55: 40-58. Johnston, S.T., McCourt, S., Bisnath, A. and Mitchell, A.A., 2002. The Tugela Terrane, Natal belt: Kibaran magmatism and tectonism along the southeast margin of the Kaapvaal craton. S. Afr. J. Geol., 105: 1-14. Jolivet, L., Goffe, B., Monie, E, Truffert-Luxey, C., Patriat, M. and Bonneau, M., 1996. Miocene detachment in Crete and exhumation P-T-t paths of high-pressure metamorphic rocks. Tectonics, 15:1129-1153. Jolly, W.T. and Hallberg, J.A., 1990. Role of crustal contamination in heterogeneous Archean volcanics from the Leonora region, Western Australia. Precambrian Res., 48: 75-98. Jones, A.E, Claeys, P. and Heuschkel, S., 2000. Impact melting of carbonates from the Chicxulub crater. Lecture Notes in Earth Sciences, 91: 343-361. Jones, J.H. and Palme, H., 2000. Geochemical constraints on the origin of the Earth and Moon. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 197-216. Jordan, T., Isacks, B., Allmendinger, R.W., Brewer, J.A., Ramos, V.A. and Ando, C.J., 1983. Andean tectonics related to geometry of subducted Nazca plate. Geol. Soc. Amer. Bull., 94:341-361. Jordan, T.H., 1988. Structure and formation of the continental tectosphere. J. Petrol., 29:11-37. Kah, L.C., Lyons, T.W. and Chesley, J.T., 2001. Geochemistry of a 1.2 Ga carbonate-evaporite succession, northern Baffin and Bylot islands: implications for Mesoproterozoic marine evolution. Precambrian Res., 111: 203-234. Kah, L.C., Lyons, T.W. and Frank, T.D., submitted. Rapid sulphur isotope variation records low sulphate concentration in the Proterozoic ocean. Nature. Kah, L.C., Sherman, A.G., Narbonne, G.M., Knoll, A.H. and Kaufman, A.J., 1999. ~ 13C stratigraphy of the Proterozoic Bylot Supergroup, Baffin Island, Canada: implications for regional lithostratigraphic correlations. Can. J. Earth Sci., 36:313-332. K~ihk6nen, Y., 1987. Geochemistry and tectonomagmatic affinities of the metavolcanic rocks of the Early Proterozoic Tampere schist belt, southern Finland. Precambrian Res., 35:295-311. K~ihk6nen, Y., 1989. Geochemistry and petrology of the metavolcanic rocks of the early Proterozoic Tampere schist belt, southern Finland. Geol. Surv. Finland, Bull. 345, 107 pp.

838

References

Kakegawa, T. and Ohmoto, H., 1999. Sulfur isotope evidence for the origin of 3.4 to 3.1 Ga pyrite at the Princeton gold mine, Barberton Greenstone Belt, South Africa. Precambrian Res., 96: 209-224. Kakegawa, T., Kawai, H. and Ohmoto, H., 1998. Origins of pyrite in the ~ 2.5 Ga Mt. McRae Shale, the Hamersley District, Western Australia. Geochim. Cosmochim. Acta, 62: 3205-3220. Kalkowsky, E., 1908. Oolith and Stromatolith im Norddeutschen Buntsandstein. Z. dt. geol. Ges., 60: 68-125. Kalliokoski, J., 1986. Calcium carbonate cement (caliche) in Keweenawan sedimentary rocks (ca. 1.1 Ga), upper Peninsula of Michigan. Precambrian Res., 32: 243-259. Kalsbeek, E and Nutman, A.P., 1996. Anatomy of the Early Proterozoic Nagssugtoqidian Orogen, West Greenland, explored by reconnaissance SHRIMP U-Pb dating. Geology, 24:515-518. Kalsbeek, E, Pulvertaft, T.C.R. and Nutman, A.P., 1998. Geochemistry, age and origin of metagreywackes from the Palaeoproterozoic Karrat Group, Rinkian Belt, West Greenland. Precambrian Res., 91: 383-399. Kalugin, A.S., 1969. Formation of the volcanogenic-sedimentary banded iron ores in the Devonian deposits of the Altai. In: "Problemy obrazovaniya zhelezistych porod dokembriya" (Problems of formation of the Precambrian iron-formations). Izd. Naukova Dumka, Kiev, pp. 89-104 (in Russian). Kalugin, A.S., 1973. Geology and genesis of the Devonian banded iron-formation in Altai, western Siberia and eastern Kazakhstan. Proceedings, Kiev Symposium: Genesis of Precambrian Iron and Manganese Deposits, 1970. UNESCO, Paris, Earth Sci. 9, pp. 159-165. Kamber, B.S., Blenkinsop, T.G., Villa, I.M. and Dahl, P.S., 1995. Proterozoic transpressive deformation in the Northern Marginal Zone, Limpopo Belt, Zimbabwe. J. Geol., 103: 493-508. Kamber, B.S., Moorbath, S. and Whitehouse, M.J., 2001. The oldest rocks on Earth: time constraints and geological controversies. In: L.C. Lewis and S.J. Knell (Eds.), The Age of the Earth: From 4004 BC to AD 2002. Geol. Soc. London, Spec. Publ., 190, pp. 177-203. Kamenetsky, V.S., Sobolev, A.V., Joron, J.-L. and Semet, M.P., 1995. Petrology and geochemistry of Cretaceous ultramafic volcanics from eastern Kamchatka. J. Petrol., 36: 637-662. Kano, K., 2003. Subaqueous pumice eruptions and their products. In: J. White, D. Clague and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph (in press). Kano, K., Orton, G.J. and Kano, T., 1994. A hot Miocene subaqueous scoria-flow deposit in the Shimane Peninsula, SW Japan. J. Volcanol. Geotherm. Res., 60: 1-14. Kano, K., Yamamoto, T. and Ono, K., 1996. Subaqueous eruption and emplacement of the Shinjima Pumice, Shinjima (Moeshima) Island, Kagoshima Bay, SW Japan. J. Volcanol. Geotherm. Res., 71: 187-206. Kaplan, I.R. and Rittenberg, S.C., 1964. Microbiological fractionation of sulphur isotopes. J. Gen. Microbiol., 34: 195-212. Karato, S. and Wu, P., 1993. Rheology of the upper mantle--a synthesis. Science, 260:771-778. Karhu, J.A. and Holland, H.D., 1996. Carbon isotopes and the rise of atmospheric oxygen. Geology, 24: 867-870. Karl, D.M., Bird, D.E, Bj6rkman, K., Houlihan, T., Shackelford, R. and Tupas, L., 1999. Microorganisms in the accreted ice of Lake Vostok, Antarctica. Science, 286:2144-2147. Karlstrom, K.E. and Bowring, S.A., 1988. Early Proterozoic assembly of tectonostratigraphic terranes in southwestern North America. J. Geol., 96: 561-576.

References

839

Karlstrom, K.E., Flurkey, A.T. and Houston, R.S., 1984. Stratigraphy and depositional setting of Proterozoic metasedimentary rocks of southeastern Wyoming: record of an Early Proterozoic Atlantic-type cratonic margin. Geol. Soc. Amer. Bull., 94: 1257-1294. Karpeta, W.P., 1989. Bedded cherts in the Rietgat Formation, Hartbeesfontein, South Africa: a late Archaean to early Proterozoic magadiitic alkaline playa lake deposit? S. Afr. J. Geol., 92: 29-36. Karpeta, W.P. and Els, B.G., 1999. The auriferous late Archean Central Rand Group of South Africa: sea-level control of sedimentation. Precambrian Res., 97:191-214. Karsakov, L.P., 1978. Deep Seated Granulites. Nauka, Moscow, 150 pp. (in Russian). Karsten, J.L., Klein, E.M. and Sherman, S.B., 1996. Subduction zone geochemical characteristics in ocean ridge basalts from the southern Chile ridge: implications of modem ridge subduction systems for the Archean. Lithos, 37: 143-161. Karte, J., 1983. Periglacial phenomena and their significance as climatic and edaphic indicators. GeoJournal, 7: 329-340. Kasting, J.E, 1987. Theoretical constraints on oxygen and carbon dioxide concentrations in the Precambrian atmosphere. Precambrian Res., 34: 205-229. Kasting, J.E, 1988. Runaway and moist greenhouse atmospheres and the evolution of Earth and Venus. Icarus, 74: 472-494. Kasting, J.E, 1993. Earth's early atmosphere. Science, 259: 920-926. Kasting, J.E, 1997. Warming early Earth and Mars. Science, 276: 1213-1215. Kasting, J.E, 2001. Earth history: the rise of atmospheric oxygen. Science, 293: 819-820. Kasting, J.E and Ackerman, T.P., 1986. Climatic consequences of very high carbon dioxide levels in the Earth's early atmosphere. Science, 234: 1383-1385. Kasting, J.E and Brown, L.L., 1998. The early atmosphere as a source of biogenic compounds. In: A. Brack (Ed.), Molecular Origins of Life. Cambridge Univ. Press, Cambridge, pp. 35-56. Kasting, J.E and Siefert, J.L., 2002. Life and the evolution of Earth's atmosphere. Science, 296: 1066-1068. Kasting, J.E, Pavlov, A.A. and Siefert, J.L., 2001. A coupled ecosystem-climate model for predicting the methane concentration in the Archean atmosphere. Orig. Life Evolut. Biosphere, 31:271-285. Kasting, J.E, Zahnle, K.J. and Walker, J.C.G., 1983. Photochemistry of methane in the Earth's early atmosphere. Precambrian Res., 20:121-148. Kato, Y., Kano, T. and Kunugiza, K., 2002. Negative Ce anomaly in the Indian Banded Iron Formations: evidence for the emergence of oxygenated deep-sea at 2.9 ~ 2.7 Ga. Resource Geology, 52:101-110. Kaufman, A.J. and Knoll, A.H., 1995. Neoproterozoic variations in the C-isotope composition of seawater: stratigraphic and biochemical implications. Precambrian Res., 73: 27-49. Kaufman, A.J., Hayes, J.M., Knoll, A.H. and Germs, G.J.B., 1991. Isotopic composition of carbonates and organic carbon from upper Proterozoic successions in Namibia: stratigraphic variation and the effects of diagenesis and metamorphism. Precambrian Res., 4: 301-327. Kaufman, A.J., Knoll, A.H. and Narbonne, G.M., 1997. Isotopes, ice ages, and Neoproterozoic Earth history. Proc. Nat. Acad. Sci. USA, 94: 6600-6605. Kaulina, T.V., 1996. U-Pb zircon dating from key objects of Belomorian-Lapland belt (northwestern Belomorie). Ph.D. Thesis (unpubl.), Saint-Petersburg, IGGD RAS, 18 pp. (in Russian). Kaulina, T.V., 1999. U-Pb zircon ages for rocks of the Tanaelv belt, Kola Peninsula, Russia. Abstracts, EUG- 10 Conference, Strasbourg, 4(1), p. 134. Kaulina, T.V. and Bogdanova, M.N., 2000. Main stages of evolution in the northwestern Belomorian area (U-Pb isotope data). Lithosphere, 12:85-94 (in Russian).

840

References

Kay, R.W. and Kay, S.M., 1993. Delamination and delamination magmatism. Tectonophys., 219: 177-189. Kazmierczak, J. and Altermann, W., 2002. Neoarchean biomineralisation by benthic cyanobacteria. Science, 298:2351. Kazmierczak, J. and Degens, E.T., 1986. Calcium and the early eukaryotes. Mitt. Geol.-Pal~iont. Inst. Univ. Hamburg, 61: 1-20. Kazmierczak, J. and Kremer, B., 2002. Thermal alteration of the Earth's oldest fossils. Nature, 420: 477-478. Kazmierczak, J., Coleman, M., Gruszczynski, M. and Kempe, S., 1996. Cyanobacterial key to the genesis of micritic and peloidal limestones in ancient seas. Acta Palaeont. Polonica, 41:319-338. Kellogg, L.H., Hager, B.H. and Van der Hilst, R.D., 1999. Compositional stratification in the deep mantle. Science, 283: 1881-1884. Kempe, A., Schopf, J.W., Altermann, W., Kudryavtsev, A.B. and Heckl, W.M., 2002. Atomic force microscopy of Precambrian fossils. Proc. Natl. Acad. Sci. USA, 99:9117-9120. Kempe, S., 1977. Hydrographie, Warvenchronologie und organische Geochemie des Van Sees, Ostt~irkei. Dissertation, Mitt. Geol.-Pal~iont. Inst. Univ. Hamburg, 47: 125-228. Kempe, S., 1979. Carbon in the fresh water cycle. In: B. Bolin, E.T. Degens, S. Kempe and P. Ketner (Eds.), The Global Carbon Cycle, SCOPE Report 13. John Wiley and Sons, New York, pp. 317-342. Kempe, S. and Degens, E.T., 1985. An early soda ocean? Chem. Geol., 53: 95-108. Kempe, S. and Kazmierczak, J., 1990a. Calcium carbonate supersaturation and the formation of in situ calcified stromatolites. In: V.A. Ittekkot, S. Kempe, W. Michaelis and A. Spitzy (Eds.), Facets of Modern Biogeochemistry. Festschrift for E.T. Degens on occasion of his 60th birthday. Springer-Verlag, Berlin, pp. 255-278. Kempe, S. and Kazmierczak, J., 1990b. Chemistry and stromatolites of the sea-linked Satonda Crater Lake, Indonesia: a recent model for the Precambrian sea? Chem. Geol., 81:299-310. Kempe, S. and Kazmierczak, J., 1993. Satonda Crater Lake, Indonesia: Hydrogeochemistry and biocarbonates. Facies, 28: 1-32. Kempe, S. and Kazmierczak, J., 1994. The role of alkalinity in the evolution of ocean chemistry, organization of living systems and biocalcification processes. In: E Doumenge (Ed.), Past and Present Biomineralization Processes. Considerations about the Carbonate Cycle. Bull. Inst. Oceanogr. Monaco, Spec. Issue, No. 13, pp. 61-117. Kempe, S. and Kazmierczak, J., 1997. A terrestrial model for an alkaline Martian hydrosphere. Planet. Space Sci., 45(11): 1493-1499. Kempe, S. and Kazmierczak, J., 2002. Biogenesis and early life on Earth and Europa: favored by an alkaline ocean? Astrobiology, 2: 123-130. Kempe, S., Kazmierczak, J. and Degens, E.T., 1989. The soda ocean concept and its bearing on biotic and crustal evolution. In: R.E. Crick (Ed.), Origin, Evolution and Modern Aspects of Biomineralization in Plants and Animals. Proc. 5th Intern. Symp. Biomineralization, Arlington, Texas, May, 1986. Plenum Press, New York, pp. 29-43. Kempe, S., Kazmierczak, J., Landmann, G., Konuk, T., Reimer, A. and Lipp, A., 1991. Largest known microbialites discovered in Lake Van, Turkey. Nature, 349: 605-608. Kempton, P.D., Downes, H., Sharkov, E.V., Vetrin, V.R., Ionov, D.A., Carswell, D.A. and Beard, A., 1995. Petrology and geochemistry of xenoliths from the Northern Baltic shield: evidence for partial melting and metasomatism in the lower crust beneath an Archean terrane. Lithos, 36: 157-184.

References

841

Kempton, ED., Downes, H., Neymark, L.A., Wartho, J.A., Zartman, R.E. and Sharkov, E.V., 2001. Garnet granulite xenoliths from the northern Baltic Shieldwthe underplated lower crust of a Palaeoproterozoic large igneous province? J. Petrol., 42(4): 731-763. Kennard, J.M. and James, N.P., 1986. Thrombolites and stromatolites: Two distinct types of microbial structures. Palaios, 1(5): 492-503. Kennedy, M.J., Runnegar, B., Prave, A.R., Hoffman, K.-H. and Arthur, M.A., 1998. Two or four Neoproterozoic glaciations. Geology, 26: 1059-1063. Kennedy, M.J., Christie-Blick, N. and Prave, A.R., 2001. Carbon isotopic composition of Neoproterozoic glacial carbonates as a test of paleoceanographic models for snowball Earth phenomena. Geology, 29:1135-1138. Kent, D.V. and Smethurst, M.A., 1998. Shallow bias of paleomagnetic inclinations in the Paleozoic and Precambrian. Earth Planet. Sci. Lett., 160: 391-402. Kerr, A.C., 1994. Lithospheric thinning during the evolution of continental large igneous provinces. Geology, 22: 1027-1030. Kerr, A.C., 1998. Oceanic plateau formation: a cause of mass extinction and black shale deposition around the Cenomanian-Turonian boundary. J. Geol. Soc., London, 155: 619--626. Kerr, A.C. and Arndt, N.T., 2001. A note on the IUGS reclassification of the high-Mg and picritic volcanic rocks. J. Petrol., 42:2169-2171. Kerr, A., Tarney, J., Marriner, G., Nivia, A. and Saunders, A., 1997. The Caribbean-Colombian Cretaceous igneous province: The internal anatomy of an oceanic plateau. In: J.J. Mahoney and M.E Coffin (Eds.), Large Igneous Provinces. Amer. Geophys. Union Monograph Ser., 100, pp. 123-144. Kerr, D., Ye, L., Bahar, A., Kelkar, B.G. and Montgomery, S., 1999. Glenn Pool field, Oklahoma: a case of improved prediction from a mature reservoir. Amer. Assoc. Petrol. Geol. Bull., 83(1): 1-18. Kerr, D.J. and Gibson, H.L., 1993. A comparison of the Horne volcanogenic massive sulfide deposit and intracauldron deposits of the Mine Sequence, Noranda, Quebec. Econ. Geol., 88: 1419-1442. Kerr, D.J. and Mason, R., 1990. A Re-appraisal of the geology and ore deposits of the Horne Mine complex at Rouyn-Noranda, Quebec. In: M. Rive, E Verpaelst, Y. Gagnon, J.M. Lulin, G. Riverin and A. Simard (Eds.), La Ceinture polym6tallique du Nord-Ouest qu6b6cois. Can. Inst. Min. Metall., Spec. Vol. 43, pp. 153-165. Kerrich, R. and Fryer, B.J., 1979. Archean precious metal hydrothermal systems, Dome Mine, Abitibi Greenstone belt, 2. REE and oxygen isotope relations. Can. J. Earth Sci., 16: 440-458. Kerrich, R., Wyman, D., Fan, J. and Bleeker, W., 1998. Boninite series: low Ti-tholeiite associations from the 2.7 Ga Abitibi greenstone belt. Earth Planet. Sci. Lett., 164: 303-316. Key, R.M., 1983. The Geology of the Area Around Gaborone and Lobatse, Kweneng, Kgatleng, Southern and South East Districts. Geol. Surv. Botswana, Gaberone, District Memoir 5, 230 pp. Khakhaev, B.N., Masaitis, V.L. and Pevsner, L.A., 1994. Deep scientific drilling in the PuchezhKatunki impact structure. Abstracts, 7th Intern. Symposium on the Observation of the Continental Crust through Drilling, Santa Fe, New Mexico, 1994, pp. 208-210. Kimberley, M.M., 1974. Origin of iron ore by diagenetic replacement of calcareous oolite. Nature, 250:319-320. Kimberley, M.M., 1978. Paleoenvironmental classification of iron formations. Econ. Geol., 73: 215-229. Kimberley, M.M., 1989. Exhalative origins of iron formations. Ore Geol. Rev., 5: 13-145. King, J. and Helmstaedt, H., 1997. The Slave Province, North-West Territories, Canada. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 459--479.

842

References

Kinoshita, M., Von Herzen, R.E, Matsubayashi, O. and Fujioka, K., 1998. Tidally-driven effluent detected by long-term temperature monitoring at the TAG hydrothermal mound, Mid-Atlantic Ridge. Phys. Earth Planet. Interiors, 108: 143-154. Kirkham, R.V. and Roscoe, S.M., 1993. Atmospheric evolution and ore deposit formation. Resource Geology, Special Issue, 15:1-17. Kirschvink, J.L., 1992. Late Proterozoic low-latitude global glaciation: the snowball Earth. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere: A multidisciplinary study. Cambridge Univ. Press, Cambridge, pp. 51-52. Kirschvink, J.L., Ripperdan, R.L. and Evans, D.A., 1997. Evidence for a large-scale reorganization of Early Cambrian continental landmasses by inertial interchange true polar wander. Science, 277:541-545. Kislitsyn, R.V., Balagansky, V.V., Manttari, I., Timmerman, M.J. and Daly, J.S., 1999. Age of accretion and collision in the Palaeoproterozoic Lapland-Kola orogen: new isotope evidence from the Kolvitsa belt and Umba granulite terrane. Abstracts, SVEKALAPKO, A Europrobe Project, Oulu, Finland, p. 33. Kisters, A.EM. and Anhaeusser, C.R., 1995. Emplacement features of Archaean TTG plutons along the southern margin of the Barberton greenstone belt, South Africa. Precambrian Res., 75: 1-15. Kisters, A.EM., Charlesworth, E.G., Gibson, R.L. and Anhaeusser, C.R., 1996. Steep structure formation in the Okiep Copper District, South Africa: bulk inhomogeneous shortening of a highgrade metamorphic granite-gneiss sequence. J. Structural Geol., 18:735-751. Kitajima, K., Maruyama, S., Utsonomiya, S. and Liou, J.G., 2001. Seafloor hydrothermal alteration at an Archaean mid-ocean ridge. J. Met. Geol., 19:581-597. Kitsul, V.I., Glebovitsky, V.A., Vapnic, Ye.A. and Frisch, T., 2000. Gneisses from the granulite terrane of the Central Boothia uplift, Arctic Canada. Can. Mineral., 38: 443-454. Kivelson, M., Khurana, K.K., Russell, C.T., Volwerk, M. and Zimmer, C., 2000. Galileo magnetometer measurements: a stronger case for a subsurface ocean at Europa. Science, 289:1340-1343. Kiyokawa, S. and Taira, A., 1998. The Cleaverville Group in the West Pilbara coastal granitoidgreenstone terrain of Westem Australia: an example of a Mid-Archaean immature oceanic islandarc succession. Precambrian Res., 88: 109-142. Klein, C., 1983. Diagenesis and metamorphism of Precambrian iron-formations. In: A.E Trendall and R.C. Morris (Eds.), Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 417-480. Klein, C. and Beukes, N.J., 1989. Geochemistry and sedimentology of a facies transition from limestone to iron-formation deposition in the early Proterozoic Transvaal Supergroup, South Africa. Econ. Geol., 84: 1733-1774. Klein, C. and Beukes, N.J., 1992. Time distribution, stratigraphy, and sedimentologic setting and geochemistry of Precambrian iron formations. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, pp. 139-146. Klein, C. and Beukes, N.J., 1993. Sedimentology and geochemistry of the glaciogenic late Proterozoic Rapitan iron-formation, in Canada. Econ. Geol., 80: 542-565. Klein, C., Beukes, N.J. and Schopf, J.W., 1987. Filamentous microfossils in the early Proterozoic Transvaal Supergroup: their morphology, significance, and palaeoenvironmental setting. Precambrian Res., 36:81-94. Klein, G. de V., 1971. A sedimentary model for determining paleotidal range. Geol. Soc. Amer. Bull., 82: 2585-2592. Klemm, D.D., 1979. A biogenic model for the formation of Banded Iron Formation in the Transvaal Supergroup/South Africa. Mineral. Deposita, 14:381-385.

References

843

Klemm, D.D., 2000. The formation of Palaeoproterozoic banded iron formations and their associated Fe and Mn deposits, with special reference to the Griqualand West deposits, South Africa. J. Afr. Earth Sci., 30: 1-24. Kline, T., Mtinker, C., Mezger, K. and Palme, H., 2002. Rapid accretion and early core formation on asteroids and the terrestrial planets from Hf-W chronometry. Nature, 418: 952-955. Klootwijk, C.T., Gee, J.S., Peirce, J.W., Smith, G.M. and McFadden, P.L., 1992. An early IndiaAsia contact; paleomagnetic constraints from Ninetyeast Ridge, ODP Leg 121; with Suppl. Data 92-15. Geology, 20: 395-398. Kloppenburg, A., White, S.H. and Zegers, T.E., 2001. Structural evolution of the Warrawoona greenstone belt and adjoining granitoid complexes, Pilbara craton, Australia; implications for Archaean tectonic processes. Precambrian Res., 112: 107-147. Knoll, A.H., 1992. The early evolution of eukaryotes: a geological perspective. Science, 256: 622-627. Knoll, A.H., 1996. Archean and Proterozoic Paleontology. In: J. Jansonius and D.C. McGregor (Eds.), Palynology: Principles and Applications. AASP Foundation, Salt Lake City, Vol. 1, pp. 51-80. Knoll, A.H. and Barghoorn, E.S., 1977. Archean microfossils showing cell division from the Swaziland System of South Africa. Science, 198: 396-398. Knoll, A.H. and Canfield, D.E., 1998. Isotopic inferences on early ecosystems. In: R.D. Norris and R.M. Corfield (Eds.), Isotopic Paleobiology and Paleoecology. Paleontolog. Soc. Papers, 4, pp. 212-243. Knoll, A.H. and Walter, M.R., 1992. Latest Proterozoic stratigraphy and Earth history. Nature, 356: 673-678. Knoll, A.H., Grotzinger, J.P., Kaufman, A.J. and Kolosov, P.N., 1995a. Integrated approaches to terminal Proterozoic stratigraphy: an example from the Olenek Uplift, northeastern Siberia. Precambrian Res., 73: 251-270. Knoll, A.H., Hayes, J.M., Kauffman, A.J., Swett, K. and Lambert, I.B., 1986. Secular variation in carbon isotope ratios from Upper Proterozoic successions of Svalbard and east Greenland. Nature, 321: 832-838. Knoll, A.H., Kaufman, A.J. and Semikhatov, M.A., 1995b. The carbon-isotope composition of Proterozoic carbonates: Riphean succession from northwestern Siberia (Anabar Massif, Turukhansk Uplift). Amer. J. Sci., 295: 823-850. Knoll, A.H., Strother, P.K. and Rossi, S., 1988. Distribution and diagenesis of microfossils from the Lower Proterozoic Duck Creek Dolomite, Western Australia. Precambrian Res., 38: 257-279. Knorre, H.V. and Krumbein, W.E., 2000. Bacterial calcification. In: R.E. Riding and S.M. Awramik (Eds.), Microbial Sediments. Springer-Verlag, Berlin, pp. 25-31. Kober, B., Pigeon, R.T. and Lippolt, H.J., 1989. Single-zircon dating by stepwise Pb-evaporation constrains the Archean history of detrital zircons from the Jack Hills, Western Australia. Earth Planet. Sci. Lett., 91: 286-296. Kocurek, G., 1996. Desert aeolian system. In: H.G. Reading (Ed.), Sedimentary Environments: Processes, Facies, and Stratigraphy, 3rd edn. Blackwell, Oxford, pp. 125-153. Kocurek, G., 1998. Aeolian system response to external forcing factors; a sequence stratigraphic view of the Saharan region. Quaternary deserts and climatic change. In: A.S. Alsharhan, K.W. Glennie, G.L. Whittle and C.G.St.C. Kendall (Eds.), Quaternary Deserts and Climatic Change. Balkema, Rotterdam, pp. 327-337. Kocurek, G. and Dott, R.H., Jr., 1981. Distinction and uses of stratification type in the interpretation of eolian sand. J. Sediment. Petrol., 51: 579-595.

844

References

Kocurek, G. and Fielder, G., 1982. Adhesion structures. J. Sediment. Petrol., 51: 1229-1241. Kocurek, G. and Havholm, K., 1993. Aeolian sequence stratigraphyma conceptual framework. In: R Weimer and H. Posamentier (Eds.), Siliciclastic Sequence Stratigraphy. Amer. Assoc. Petrol. Geol., Tulsa, Oklahoma, Mere. 58, pp. 393-409. Kocurek, G. and Nielson, J., 1986. Conditions favorable for the formation of warm-climate aeolian sand sheets. Sedimentology, 33:795-816. Koeberl, C., 1992. Tektite origin by hypervelocity asteroidal or cometary impact; the quest for the source craters. LPI Contribution, 790: 42-43. Koeberl, C., 1998. Identification of meteoritical components in impactites. In: M.M. Grady, R. Hutchison, G.J.H. McCall and D.A. Rothery (Eds.), Meteorites: Flux with Time and Impact Effects. Geol. Soc. London, Spec. Publ. 140, pp. 133-153. Kojima, S., Hanamuro, T., Hayashi, K., Haruna, M. and Ohmoto, H., 1998. Sulphide minerals in Early Archean chemical sedimentary rocks of the eastern Pilbara district, Western Australia. Mineral. Petrol., 64:219-235. Kokelaar, B.P. and Busby, C.J., 1992. Subaqueous explosive eruption and welding of pyroclastic deposits. Science, 257: 196-201. Komar, V.A., 1989. Classification of microstructures of the USSR Precambrian stromatolites. Himalayan Geology, 13: 229-238. Kominz, M.A. and Bond, G.C., 1991. Unusually large subsidence and sea-level events during middle Paleozoic time: new evidence supporting mantle convection models for supercontinent assembly. Geology, 19: 56-60. Komiya, T., Hayashi, M., Maruyama, S. and Yurimoto, H., 2002. Intermediate-P/T type Archean metamorphism of the Isua supracrustal belt: implications for secular change of geothermal gradients at subduction zones and Archean plate tectonics. Amer. J. Sci., 302: 806-826. Komiya, T., Maruyama, S., Masuda, T., Nohda, S., Hayashi, M. and Okamoto, K., 1999. Plate tectonics at 3.8-3.7 Ga: field evidence from the Isua Accretionary Complex, southern West Greenland. J. Geol., 107: 515-554. Konhauser, K.O. and Ferris, F.G., 1996. Diversity of iron and silica precipitation by microbial mats in hydrothermal waters, Iceland: Implications for Precambrian iron formations. Geology, 24: 323-326. Konhauser, K.O., Hamade, T., Morris, R.C., Ferris, F.G., Southam, G., Raiswell, R. and Canfield, D.E., 2002. Could bacteria have formed the Precambrian banded iron formations? Geology, 30:1079-1082. Kontinen, A., 1987. An Early Proterozoic ophioliteBthe Jormua mafic-ultramafic complex, northeastern Finland. Precambrian Res., 35(1): 313-341. Korsman, K., Korja, T., Pajunen, M., Virransalo, R and GGT/SVEKA Working Group, 1999. The GGT/SVEKA transect: structure and evolution of the continental crust in the Palaeoproterozoic Svecofennian orogen in Finland. Intem. Geol. Review, 41: 287-333. Koster, E.H. and Steel, R.J. (Eds.), 1984. Sedimentology of Gravels and Conglomerates. Can. Soc. Petrol. Geol., Calgary, Mem., 10, 441 pp. Kotov, A.B., Kovach, V.P., Sal'nikova, E.B., Glebovitsky, V.A., Yakovleva, S.Z., Berezhnaya, N.G. and Myskova, T.A., 1995. Stages of continental crust formation in the central Aldan granulitegneiss domain: U-Pb and Sm-Nd isotopic data on granitoids. Petrology, 3(1): 99-110 (official English translation). Kovach, V.R, Kotov, A.B., Beryozkin, V.I., Sal' nikova, E.B., Velikoslavinsky, S.D., Smelov, A.P. and Zagornaya, N.Yu., 1999. Age boundaries of the high-grade supracrustal complexes of the central

References

845

Aldan Shield: Sm-Nd isotopic data. Stratigr. and Geol. Correlation, 7(1): 3-17 (official English translation). Kramers, J.D. and Tolstikhin, I.N., 1997. Two terrestrial lead isotope paradoxes, forward transport modelling, core formation and the history of the continental crust. Chem. Geol., 139:75-110. Kramers, J.D., N~igler, T.F. and Tolstikhin, I.N., 1998. Perspectives from global modeling of terrestrial Pb and Nd isotopes on the history of the continental crust: Schweiz. Min. Petr. Mitt., 78: 169-174. Kramers, J.D., Kreissig, K. and Jones, M.Q.W., 2001. Crustal heat production and style of metamorphism: a comparison between two high grade provinces in the Limpopo Belt, southern Africa. Precambrian Res., 112: 149-163. Krapez, B., 1993. Sequence stratigraphy of the Archaean supracrustal belts of the Pilbara Block, Western Australia. Precambrian Res., 60: 1-45. Krapez, B., 1996. Sequence-stratigraphic concepts applied to the identification of basin-filling rhythms in Precambrian successions. Austral. J. Earth Sci., 43: 355-380. Krapez, B., 1997. Sequence-stratigraphic concepts applied to the identification of depositional basins and global tectonic cycles. Austral. J. Earth Sci., 44: 1-36. Krapez, B. and Barley, M.E., 1987. Archaean strike-slip faulting and related ensialic basins: evidence from the Pilbara Block, Australia. Geol. Mag., 124: 555-567. Krapez, B. and Eisenlohr, B., 1998. Tectonic settings of Archaean (3325-2775 Ma) crustalsupracrustal belts in the West Pilbara Block. Precambrian Res., 88: 173-205. Krapez, B., Barley, M.E. and Pickard, A.L., 2002. Banded iron formations: ambient pelagites, hydrothermal muds or metamorphic rocks? Extended Abstracts, 4th International Archaean Symposium, 2001. AGSO-Geoscience Australia, Record 2001/37: 247-248. Kreisa, R.D., Moiola, R.J. and Ncttvedt, A., 1986. Tidal sand wave facies, Rancho Rojo Sandstone (Permian), Arizona. In: R.J. Knight and J.R. McLean (Eds.), Shelf Sands and Sandstones. Can. Soc. Petrol. Geologists, Calgary, Mem., 11, pp. 277-291. Kreissig, K., Holzer, L., Frei, R., Villa, I.M., Kramers, J.D., Kr6ner, A., Smit, C.A. and Van Reenen, D.D., 2001. Geochronology of the Hout River Shear Zone and the metamorphism in the Southern Marginal Zone of the Limpopo Belt, Southern Africa. Precambrian Res., 109: 145-173. Krogh, T.E., 1982. Improved accuracy of U-Pb zircon ages by the creation of more concordant systems using an air abrasion technique. Geochim. Cosmochim. Acta, 46: 637-649. Kr6ner, A., 1985a. Evolution of the Archaean continental crust. Annu. Rev. Earth Planet. Sci., 13: 49-74. Kr6ner, A., 1985b. Ophiolites and the evolution of tectonic boundaries in the late Proterozoic Arabian-Nubian Shield of northeast Africa and Arabia. Precambrian Res., 27: 277-300. Kr6ner, A., 1991. Tectonic evolution in the Archaean and Proterozoic. Tectonophysics, 187: 393-410. Kr6ner, A., 2001. The Mozambique belt of East Africa and Madagascar: significance of zircon and Nd model ages for Rodinia and Gondwana supercontinent formation and dispersal. S. Afr. J. Geol., 104:151-166. Kr6ner, A. and Layer, P.W., 1992. Crust formation and plate motion in the Early Archaean. Science, 256:1405-1411. Kr6ner, A., Greiling, R., Reischmann, T., Hussein, I.M., Stem, R.J., Diirr, S., KriJger, J. and Zimmer, M., 1987. Pan-African crustal evolution in the Nubian segment of northeast Africa. In: A. Kr6ner (Ed.), Proterozoic Lithospheric Evolution. Geodyn. Ser., 17. Amer. Geophys. Union, pp. 235-257.

846

References

Kr6ner, A., Linnebacher, R, Stem, R.J., Reischmann, T., Manton, W. and Hussein, I.M., 1991. Evolution of Pan-African island arc assemblages in the southern Red Sea Hills, Sudan, and in southwestern Arabia as exemplified by geochemistry and geochronology. Precambrian Res., 53: 99-118. Kr6ner, A., Jaeckel, E, Brandl, G., Nechin, A.A. and Pidgeon, R.T., 1999. Single zircon ages for granitoid gneisses in the Central Zone of the Limpopo Belt, southern Africa and geodynamic significance. Precambrian Res., 93: 299-337. Kropp, J., Block, A., Von Bloh, W., Klenke, T. and Schellnhuber, H.J., 1997. Multifractal characterization of microbially induced magnesian calcite formation in recent tidal flat sediments. Sediment. Geol., 109:37-51. Krumbein, W.E., 1974. On precipitation of aragonite on the surface of marine bacteria. Naturwissenschaften, 61: 167. Krumbein, W.E., 1979. Calcification by bacteria and algae. In: P.A. Trudinger and D.J. Swaine (Eds.), Biogeochemical Cycling of Mineral--Forming Elements. Elsevier, Amsterdam, pp. 47-68. Krumbein, W.E., 1983. Stromatolites--the challenge of a term in space and time. Precambrian Res., 20:493-531. Krumbein, W.E., 1986. Biotransfer of minerals by microbes and microbial mats. In: B.S.C. Leadbeater and R. Riding (Eds.), Biomineralization in lower plants and animals. Proc. Systematics Assoc. Symposium, Univ. Birmingham, April 1985. Oxford Univ. Press, Oxford, pp. 55-72. Krumbein, W.E. and Cohen, Y., 1977. Primary production, mat formation and lithification: contribution of oxygenic and facultative anoxygenic cyanobacteria. In: E. Fltigel (Ed.), Fossil Algae: Recent Results and Development. Springer-Verlag, Berlin, pp. 37-56. Krupp, R., Oberthur, T. and Hirdes, W., 1994. The Early Precambrian atmosphere and hydrosphere: thermodynamic constraints from mineral deposits. Econ. Geol., 89: 1581-1598. Krylov, I.N., 1976. Approaches to the classification of stromatolites. In: M.R. Walter (Ed.), Stromatolites. Developments in Sedimentology, Vol. 20. Elsevier, Amsterdam, pp. 31-44. Krynauw, J.R., Watters, B.R., Hunter, D.R. and Wilson, A.H., 1991. A review of the field relations, petrology and geochemistry of the Borgmassivet intrusions in the Grunehogna Province, western Dronning Maud Land, Antarctica. In: M.R.A. Thomson, J.A. Crame and J.W. Thomson (Eds.), Geological Evolution of Antarctica. Cambridge Univ. Press, Cambridge, pp. 33-39. Kudryavtsev, A.B., Schopf, J.W., Agresti, D.G. and Wdowiak, T.J., 2001. In situ laser-Raman imagery of Precambrian microscopic fossils. Proc. Natl. Acad. Sci. USA, 98: 823-826. Kuecher, G.J., Woodland, B.G. and Broadhurst, EM., 1990. Evidence of deposition from individual tides and of tidal cycles from the Francis Creek Shale (host rock to the Mazon Creek Biota), Westphalian D (Pennsylvanian), northeastern Illinois. Sediment. Geol., 68:211-221. Ktihl, M., Fenchel, T. and Kazmierczak, J., in press. Structure, function, and calcification potential of an artificial cyanobacterial mat. In: W.E. Krumbein, D. Paterson and G. Zavarzin (Eds.), Fossil and Recent Biofilms, a Natural History of the Impact of Life on Planet Earth. Kluver Academic Publishers, Dordrecht. Kuhn, W.R. and Atreya, S.K., 1979. Ammonia photolysis and the greenhouse effect in the primordial atmosphere of the Earth. Icarus, 37: 207-213. Kuhn, W.R., Walker, J.C.G. and Marshall, H.G., 1989. The effect on Earth's surface temperature from variations in rotation rate, continent formation, solar luminosity, and carbon dioxide. J. Geophys. Res., 94:11,129-11,136. Kuklis, C.A., Eriksson, K.A. and Simpson, E.L., 2000. Quantification of an Archean tidal record: the Witwatersrand Supergroup, South Africa. Geol. Soc. Amer., Abstracts with Programs, 32: A307.

References

847

Kump, L.R., 1991. Interpreting carbon-isotope excursions: strangelove oceans. Geology, 19: 299-302. Kump, L.R., 1993. Bacteria forge a new link. Nature, 362:790-791. Kump, L.R. and Arthur, M.A., 1999. Interpreting carbon-isotope excursions: carbonates and organic matter. Chem. Geol., 161: 181-198. Kump, L.R., Kasting, J.F. and Barley, M.E., 2001. Rise of atmospheric oxygen and the "upsidedown" Archean mantle. Geochem. Geophys. Geosyst., 2, Paper No. 2000GC000114. Kumpulainen, R.A., Mansfeld, J., Sunblad, K., Neymark, L. and Bergman, T., 1996. Stratigraphy, age, and Sm-Nd isotope systematics of the country rocks to Zn-Pb sulfide deposits,/~mmeberg district, Sweden. Econ. Geol., 91: 1009-1021. Kunz, J., Staudacher, T. and All~gre, C.J., 1998. Plutonium-fission Xenon found in Earth's mantle. Science, 280: 877-880. Kusky, T.M., 1993. Collapse of Archaean orogens and the generation of late to post-kinematic granitoids. Geology, 21: 925-928. Kusky, T.M., 1997. Principles of Precambrian Geology (book review). GSA Today, 7(5): 29-34. Kusky, T.M., 1998. Tectonic setting and terrane accretion of the Archaean Zimbabwe craton. Geology, 26:163-166. Kusky, T.M. and Hudleston, P.J., 1999. Growth and demise of an Archean carbonate platform, Steep Rock Lake, Ontario, Canada. Can. J. Earth Sci., 36: 565-584. Kusky, T.M. and Kidd, W.S.F., 1992. Remnants of an Archean oceanic plateau, Belingwe greenstone belt, Zimbabwe. Geology, 99: 60-62. Kusky, T.M. and Polat, A., 1999. Growth of granite-greenstone terranes at convergent margins, and stabilization of Archean cratons. Tectonophysics, 305: 43-73. Kusky, T.M. and Vearncombe, J.R., 1997. Structural aspects. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Science Publications, Oxford, pp. 91-123. Kusky, T.M. and Winsky, P.A., 1995. Structural relationships along a greenstone/shallow water shelf contact, Belingwe greenstone belt, Zimbabwe. Tectonics, 14:448-471. Kusky, T.M., Li, J.-H. and Tucker, R.D., 2001. The Archean Dongwanzi ophiolite complex, North China Craton, 2.505-Billion-year-old oceanic crust and mantle. Science, 292:1142-1145. Kvale, E.P., Johnson, H.W., Sonett, C.P., Archer, A.W. and Zawitoski, A., 1999. Calculating lunar retreat rates using tidal rhythmites. J. Sediment. Res., 69:1154-1168. Kyser, K., Hiatt, E., Renac, C., Durocher, K., Holk, G. and Deckart, K., 2000. Diagenetic fluids in Paleo- and Meso-Proterozoic sedimentary basins and their implications for long protracted fluid histories. In: K. Kyser (Ed.), Fluids and Basin Evolution. Short Course Ser. 28. Mineral. Assoc. Can., pp. 225-262. Kyte, ET., 2002. Composition of impact melt debris from the Eltanin impact strewn field, Bellingshausen Sea. Deep Sea Research II, 49: 1029-1047. Kyte, ET., Shukolyukov, A., Lugmair, G.W., Lowe, D.R. and Byerly, G.R., 2003. Early Archaean spherule beds: chromium isotopes confirm origin through multiple impacts of carbonaceous chondrite-type projectiles. Geology, 31: 283-286. Kyte, ET., Zhou, L. and Lowe, D., 1992. Noble metal abundances in an Early Archaean impact deposit. Geochim. Cosmochim. Acta, 56: 1365-1372. LaBerge, G.L., 1973. Possible biological origin of Precambrian iron formations. Econ. Geol., 68: 1098-1019. Lacroix, S. and Sawyer, E.W., 1995. An Archean fold-thrust belt in the northwestern Abitibi greenstone belt: structural and seismic evidence. Can. J. Earth Sci., 32:97-112.

848

References

Lafrance, B., 2003. Reconstruction d'un environnement de sulfures massifs volcanog~nes d6form6s: Exemple arch6en de Norm6tal, Abitibi. Ph.D. Thesis (unpubl.), Universit6 du Qu6bec ?a Chicoutimi, Canada (in press). Lafrance, B., Mueller, W.U., Daigneault, R. and Dupras, N., 2000. Evolution of a submerged composite arc volcano: volcanology and geochemistry of the Normetal volcanic complex, Abitibi greenstone belt, Quebec, Canada. Precambrian Res., 101: 277-311. Lahaye, Y., Arndt, N.T., Byerly, G., Chauvel., C., Fourcade, S. and Gruau, G., 1995. The influence of alteration on the trace element and Nd isotopic compositions of komatiites. Chem. Geol., 126: 43-64. Lajoie, J. and G61inas, L., 1978. Emplacement of Archean peridotitic komatiites in La Motte township, Quebec. Can. J. Earth Sci., 15: 672-677. Lamb, S.H., 1986. Archean synsedimentary tectonic deformation--a comparison with the Quaternary. Geology, 15: 565-568. Lambeck, K., 1980. The Earth's Variable Rotation: Geophysical Causes and Consequences. Cambridge Univ. Press, Cambridge, 449 pp. Lambert, D.D., Shirey, S.B. and Bergman, S.C., 1995. Proterozoic lithospheric mantle source for the Prairie Creek lamproites: Re-Os and Sm-Nd isotopic evidence. Geology, 23: 273-276. Lambert, I.B. and Donnelly, T.H., 1992. Global oxidation and a supercontinent in the Proterozoic: evidence from stable isotopic trends. In: M. Schidlowski (Ed.), Early Organic Evolution: Implications for Mineral and Energy Resources. Springer-Verlag, Heidelberg, pp. 408-414. Lambert, I.B., Donnelly, T.H., Dunlop, J.S.R. and Groves, D.I., 1978. Stable isotopic compositions of early Archean sulphate deposits of probable evaporitic and volcanogenic origins. Nature, 276: 808-810. Lambert, K.C., Walter, M.R., Wenlong, Z., Lu, S. and Ma, G., 1987. Paleoenvironmental and carbon isotope stratigraphy of upper Proterozoic carbonates of the Yangtze Platform. Nature, 325: 140-142. Lambert, M.B., Beaumont-Smith, C. and Paul, D., 1992. Structure and stratigraphic succession of an Archean stratovolcano, Slave Province, Northwest Territories. Current Res., Part C, Geol. Surv. Can. Paper, 92-1 C, pp. 189-200. Lambert, M.B., Burbridge, G.J., Beaumont-Smith, C. and Lustwerk, R., 1990. Stratigraphy, facies and structure in volcanic and sedimentary rocks of the Archean Back River volcanic complex, N.W.T. Current Res., Part C, Geol. Surv. Can. Paper, 90-1C, pp. 151-165. Landais, P. and Dereppe, J.M., 1985. A chemical study of the carbonaceous material from the Carswell structure. In: R. Lain6, D. Alonso and M. Svab (Eds.), The Carswell structure uranium deposits, Saskatchewan. Geol. Assoc. Can. Spec. Paper, 29, pp. 165-174. Lane, A.L., Nelson, R.M. and Matson, P.M., 1981. Evidence for sulfur implantation in Europa's UV band. Nature, 291: 38-39. Langenhorst, E and Masaitis, V.L., 1996. Microstructural characteristics of impact diamonds from the Popigai Crater, Russia. Meteoritics Planet. Sci., 31: 77. Langford, F.E, 1981. The Martin Group in the greater Beaverlodge area. Saskatchewan Geol. Surv., Summary of Investigations 1981, Misc. Report 1981(4), pp. 38-43. Langford, EE and Morin, J.A., 1976. The development of the Superior Province of the northwest Ontario by merging island arcs. Amer. J. Sci., 276: 1023-1034. Lanier, W.P., 1986. Approximate growth rates of Early Proterozoic microstromatolites as deduced by biomass productivity. Palaios, 1: 525-542. Lanier, W.P. and Lowe, D.R., 1982. Sedimentology of the Middle Marker (3.4 Ga) Onverwacht Group, Transvaal, South Africa. Precambrian Res., 18: 237-260.

References

849

Larsen, T.B. and Yuen, D.A., 1997. Fast plumeheads: temperature-dependent versus non-Newtonian rheology. Geophys. Res. Lett., 24: 1995-1998. Larson, J.E. and Hutchinson, R.W., 1993. The Selbaie Zn-Cu-Ag deposits Quebec: an example of evolution from subaqueous to subaerial volcanism and mineralization in an Archean caldera environment. Econ. Geol., 88: 1460-1482. Larson, R.L., 1991a. Latest pulse of Earth: evidence for a mid-Cretaceous superplume. Geology, 19: 547-550. Larson, R.L., 199 lb. Geological consequences of superplumes. Geology, 19: 963-966. Larson, R.L. and Kincaid, C., 1996. Onset of mid-Cretaceous volcanism by elevation of the 670 km thermal boundary layer. Geology, 24:551-554. Lasaga, A.C., 1998. Kinetic Theory in the Earth Sciences. Princeton Univ. Press, Princeton, New Jersey, 811 pp. Lasaga, A.C. and Ohmoto, H., 2002. The oxygen geochemical cycle: dynamics and stability. Geochim. Cosmochim. Acta, 66: 361-381. Laskar, J. and Robutel, E, 1993. The chaotic obliquity of the planets. Nature, 361 : 608-612. Laskar, J., Joutel, F. and Robutel, E, 1993. Stabilization of the Earth's obliquity by the Moon. Nature, 361: 615-617. Lawrie, K.C., 1992. Geochemical characterisation of a polyphase deformed, altered, and high grade metamorphosed volcanic terrane: implications for the tectonic setting of the Svecofennides, south-central Finland. Precambrian Res., 59:171-205. Le Bas, M.J., 2000. IUGS Reclassification of the high-Mg and picritic volcanic rocks. J. Petrol., 41: 1467-1470. Le Bas, M.J., 2001. Reply to Comment by Kerr and Arndt. J. Petrol., 42:2173-2174. Le Messurier, P., Williams, B.T. and Blake, D.H., 1990. Tennant Creek Inlier--regional geology and mineralisation. In: EE. Hughes (Ed.), Geology of the Mineral Deposits of Australia and Papua New Guinea. The Australian Institute of Mining and Metallurgy, Monographs, 14, pp. 829-838. Le Roux, J.P. and Gresse, EG., 1983. The sedimentary-tectonic realm of the Kango Group. In: A.P.G. S6hnge and I.W. H~ilbich (Eds.), Geodynamics of the Cape Fold Belt. Geol. Soc. S. Afr., Johannesburg, Spec. Publ. 12, pp. 33-46. Leather, J., Allen, EA., Brasier, M.D. and Cozzi, A., 2002. Neoproterozoic snowball Earth under scrutiny: evidence from the Fiq glaciation of Oman. Geology, 30:891-894. LeCheminant, A.N., Miller, A.R. and LeCheminant, G.M., 1987. Early Proterozoic alkaline igneous rocks, District of Keewatin, Canada: Petrogenesis and mineralization. In: T.C. Pharaoh, R.D. Bekinsale and D. Rickard (Eds.), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. Soc. London, pp. 219-240. Leckie, D.A., 1994. Canterbury Plains, New Zealand--implications for sequence stratigraphic models. Amer. Assoc. Petrol. Geol. Bull., 78: 1240-1256. Ledru, E, Johan, V., Milesi, J.-E and Tegyey, M., 1994. Markers of the last stages of the Palaeoproterozoic collision: evidence for a 2 Ga continent involving circum-South Atlantic provinces. Precambrian Res., 69: 169-191. Lee, D.-C. and Halliday, A.N., 1995. Hafnium-tungsten chronometry and the timing of terrestrial core formation. Nature, 378: 771-774. Lee, D.-C. and Halliday, A.N., 2000a. Hf-W internal isochrons for ordinary chondrites and the initial 182HfflS~ of the solar system. Chem. Geol., 169: 35-43. Lee, D.-C. and Halliday, A.N., 2000b. Accretion of primitive planetesimals: Hf-W isotopic evidence from enstatite chondrites. Science, 288:1629-1631.

850

References

Lee, D.-C., Halliday, A.N., Snyder, G.A. and Taylor, L.T., 1997. Age and origin of the Moon. Science, 278:1098-1103. Lee, D.-C., Halliday, A.N., Leya, I., Wieler, R. and Wiechert, U., 2002. Cosmogenic tungsten and the origin and earliest differentiation of the Moon. Earth Planet. Sci. Lett., 198: 267-274. Lee, J.K.W., Williams, I.S. and Ellis, D.J., 1997. Pb, U and Th diffusion in natural zircon. Nature, 390: 159-162. Lee, T. and Papanastassiou, D.A., 1974. Mg isotopic anomalies in the Allende meteorite and correlation with O and Sr effects. Geophys. Res. Lett., 1: 225-228. Lee, T., Papanastassiou, D.A. and Wasserburg, G.J., 1976. Demonstration of 26Mg excess in Allende and evidence for 26A1. Geophys. Res. Lett., 3:109-112. Legault, M., Daigneault, R. and Couture, J.-F., 1997. Contexte structural et mttallog6nique des indices aurif'eres du couloir de d6formation de Fancamp. Ministtre des Ressources naturelles du Qutbec, MB-97-32, 60 pp. Legault, M., Gauthier, M., Jebrak, M., Davis, D.W. and Baillargeon, E, 2002. Evolution of the subaqueous to near-emergent Joutel volcanic complex. Northern Volcanic Zone, Abitibi Subprovince, Quebec, Canada. Precambrian Res. 115:187-221. Leith, C. K., 1903. The Mesabi Iron-Beating District of Minnesota. U.S. Geol. Surv. Monograph, 43, 316 pp. Lepp, H. and Goldich, S.S., 1964. Origin of Precambrian iron-formations. Econ. Geol., 59: 1025-1060. Lesher, C.M. and Amdt, N.T., 1995. REE and Nd isotope geochemistry, petrogenesis and volcanic evolution of contaminated komatiites at Kambalda, Western Australia. Lithos, 34: 127-157. Lesher, C.M., Arndt, N.T. and Groves, D.I., 1984. Genesis of komatiite-associated nickel sulphide deposits at Kambalda, Western Australia: a distal volcanic model. In: D.L. Buchanan and M.J. Jones (Eds.), Sulphide Deposits in Mafic and Ultramafic rocks. Inst. Min. Metal., London, pp. 70-80. Levin, G.V., 2001. The oxides of Mars. Instruments, Methods, and Missions for Astrobiology, SPIE Proceedings, p. 4416. Lewry, J.E and Collerson, K.D., 1990. The Trans-Hudson Orogen: extent, subdivision, and problems. In: J.E Lewry and M.R. Stauffer (Eds.), The Early Proterozoic Trans-Hudson Orogen of North America. Geol. Assoc. Can. Spec. Paper, 37, pp. 1-14. Lewry, J.F. and Stauffer, M.R., 1997. The Reindeer zone of the southern Trans-Hudson Orogen, Canada. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 739-746. Leybourne, M.I., Van Wagoner, N.A. and Ayres, L.D., 1997. Chemical stratigraphy and petrogenesis of the Early Proterozoic Amisk Lake volcanic sequence, Flin Flon-Snow Lake greenstone belt, Canada. J. Petrol., 38: 1541-1564. Leybourne, M.I., Van Wagoner, N. and Ayres, L.D., 1999. Partial melting of a refractory subducted slab in a Paleoproterozoic island arc: implications for global chemical cycles. Geology, 27:731-734. Li, C.-W., Chen, J.-Y. and Hua, T.-E., 1998. Precambrian sponges with cellular structures. Science, 279: 879-882. Li, X.H., Li, Z.X., Zhou, H., Liu, Y. and Kinny, P.D., 2002. U-Pb zircon geochronology, geochemistry and Nd isotopic study of Neoproterozoic bimodal volcanic rocks in the Kangdian Rift of South China: implications for the initial rifting of Rodinia. Precambrian Res., 113: 135-154. Li, Z.X., Zhang, L. and Powell, C.M., 1996. Positions of the east Asian cratons in the Neoproterozoic supercontinent Rodinia. Austral. J. Earth Sci., 43: 593-604.

References

851

Li, Z.X., Li, X.H., Kinny, RD. and Wang, J., 1999. The breakup of Rodinia: did it start with a mantle plume beneath South China? Earth Planet. Sci. Lett., 173: 171-181. Li, Z.X., Li, X.H., Kinny, ED., Wang, J., Zhang, S. and Zhou, H., 2003. Geochronology of Neoproterozoic syn-rift magmatism in the Yangtze Craton, South China and correlations with other continents: evidence for a mantle superplume that broke up Rodinia. Precambrian Res. (in press). Lichtblau, A., 1989. Stratigraphy and facies at the south margin of the Archean Noranda caldera. M.Sc. Thesis (unpubl.), Univ. Qudbec ?aChicoutimi, Qu6bec, 135 pp. (1 map). Lichtblau, A. and Dimroth, E., 1980. Stratigraphy and facies at the south margin of the Archean Noranda caldera, Noranda, Quebec. Current Res., Part I A, Geol. Surv. Canada, Paper, 80-1A, pp. 67-75. Light, M.ER., 1982. The Limpopo Belt: a result of continental collision. Tectonics, 1: 325-342. Lindsay, J.E, 1987. Sequence stratigraphy and depositional controls in Late Proterozoic-Early Cambrian sediments of Amadeus Basin, central Australia. Amer. Assoc. Petrol. Geol. Bull., 71: 1387-1403. Lindsay, J.E, 1989. Depositional controls on glacial facies associations in a basinal setting, Late Proterozoic, Amadeus Basin, central Australia. Palaeogeog. Palaeoclimatol. Palaeoecol., 73: 205-232. Lindsay, J.E (Ed.), 1993. Geological Atlas of the Amadeus Basin. Australian Geological Survey Organisation, 25 plates. Lindsay, J.E, 1998. The Broadmere Structure: a window into Palaeoproterozoic mineralisation, McArthur Basin, Northern Australia. Australian Geological Survey Organisation, Record 98/38 (CD-ROM). Lindsay, J.E, 1999. The Heavitree Quartzite, a Neoproterozoic (c. 800-760 Ma), high-energy, tidally-influenced, ramp association, Amadeus Basin, Central Australia. Austral. J. Earth Sci., 46: 127-149. Lindsay, J.E, 2001. Basin dynamics and mineralisation, McArthur Basin, Northern Australia. Austral. J. Earth Sci., 48: 703-720. Lindsay, J.E, 2002. Supersequences, superbasins, supercontinents--evidence from the Neoproterozoic basins of central Australia. Basin Res., 14: 204-223. Lindsay, J.E and Brasier, M.D., 2000. A carbon isotope reference curve for c. 1700 to 1575 Ma, McArthur and Mount Isa Basins, northern Australia. Precambrian Res., 99:271-308. Lindsay, J.E and Brasier, M.D., 2002. Did global tectonics drive early biosphere evolution? Carbon isotope record from 2.6 to 1.9 Ga carbonates of Western Australian basins. Precambrian Res., 114: 1-34. Lindsay, J.E and Korsch, R.J., 1989. Interplay of tectonics and sea-level changes in basin evolution; an example from the intracratonic Amadeus Basin, central Australia. Basin Res., 2: 3-25. Lindsay, J.E and Leven, J.E, 1996. Evolution of a Neoproterozoic to Palaeozoic intracratonic setting, Officer Basin, South Australia. Basin Res., 8: 403-424. Lindsay, J.E, Korsch, R.J. and Wilford, J.R., 1987. Timing the breakup of a Proterozoic supercontinent: evidence from Australian intracratonic basins. Geology, 15:1061-1064. Lindsey, D.A., 1969. Glacial sedimentation of the Precambrian Gowganda Formation, Ontario, Canada. Geol. Soc. Amer. Bull., 80: 1685-1704. Lindsey, K.A. and Gaylord, D.R., 1992. Fluvial, coastal, nearshore, and shelf deposition in the Upper Proterozoic (?) to Lower Cambrian Addy Quartzite, northeastern Washington. Sediment. Geol., 77: 15-35. Lipman, E W., 1997. Subsidence of ash-flow calderasmrelation to caldera size and magma-chamber geometry. Bull. Volcanol., 59:198-218.

852

References

Lipman, EW., 2000. Calderas. In: H. Sigurdsson, B. Houghton, S. McNutt, H. Rymer and J. Stix (Eds.), Encyclopedia of Volcanoes. Academic Press, New York, pp. 643-662. Lofgren, G.E. and Russell, W.J., 1986. Dynamic crystallization of chondrule melts of porphyritic and radial pyroxene composition. Geochim. Cosmochim. Acta, 50:1715-1726. Lofgren, G.E., Donaldson, C.H., Williams, R.J., Mullins, O. and Ussehnan, T.M., 1974. Experimentally reproduced textures and mineral chemistry of Apollo 15 quartz-normative basalts. Proc. Lunar Sci. Conf., 6: 79-99. Logan, B.W., 1961. Cryptozoon and associated stromatolites from the Recent, Shark Bay, Western Australia. J. Geol., 69: 517-533. Logan, B.W., Rezak, R. and Ginsburg, R.N., 1964. Classification and environmental significance of algal stromatolites. J. Geol., 72: 68-83. Logan, G.A., Hayes, J.M., Hieshima, G.B. and Summons, R.E., 1995. Terminal Proterozoic reorganization of biogeochemical cycles. Nature, 376: 53-56. Logan, G.A., Calver, C.R., Gorjan, P., Summons, R.E., Hayes, J.M. and Walter, M.R., 1999. Terminal Proterozoic mid-shelf benthic microbial mats in the Centralian Superbasin and their environmental significance. Geochim. Cosmochim. Acta, 63: 1345-1358. Logan, G.A., Hinman, M.C., Walter, M.R. and Summons, R.E., 2001. Biogeochemistry of the 1640 Ma McArthur River (HYC) lead-zinc ore and host sediments, Northern Territory, Australia. Geochim. Cosmochim. Acta, 65: 2317-2336. Logan, J.A., Prather, M.J., Wofsy, S.C. and McElroy, M.B., 1981. Tropospheric chemistry: a global perspective. J. Geophys. Res., 86(C8): 7210-7254. Long, D.G.E, 1978. Proterozoic stream deposits: some problems of recognition and interpretation of ancient sandy fluvial systems. In: A.D. Miall (Ed.), Fluvial Sedimentology. Can. Soc. Petrol. Geol., Calgary, Alberta, Mem., 5, pp. 313-342. Long, D.G.F., 1987. Sedimentary framework of uranium deposits in the southern Cobalt Embayment, Ontario Canada. In: D. Pretorius (Ed.), Uranium deposits in Proterozoicquartz-pebble conglomerates. International Atomic Energy Agency, Tech. Doc., 427, pp. 155-188. Long, D.G.E, 2001a. Aspects of late Paleoproterozoic fluvial style: the Uair6n Formation, Roraima Supergroup, Venezuela. In: W. Alterman and P. Corcoran (Eds.), Precambrian Sedimentary Environments: a modern approach to ancient depositional systems. IAS Spec. Publ., 33. Blackwell, Oxford, pp. 323-338. Long, D.G.E, 200lb. Architecture of Late Paleoproterozoic Fluvial Strata in the Manitou Falls Formation at McClean Lake and Key Lake, Eastern Athabasca Basin, Northern Saskatchewan. Sask. Energy Mines, Geol. Sur., Misc. Rep., 2001-2 (CD). Loope, D.B., Dingus, L., Swisher, C.C., HI and Minjin, C., 1998. Life and death in a Late Cretaceous dunefield, Nemegt basin, Mongolia. Geology, 26: 27-30. Loope, D.B., Mason, J.A. and Dingus, L., 1999. Lethal landslides from eolian dunes. J. Geol., 107: 707-713. Lottermoser, B.G. and Ashley, P.M., 2000. Geochemistry, petrology and origin of Neoproterozoic ironstones in the eastern part of the Adelaide geosyncline, South Australia. Precambrian Res., 101: 49-67. Loutit, T.S., Hardenbol, J., Vail, P.R. and Baum, G.R., 1988. Condensed sections: the key to age-dating and correlation of continental margin sequences. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Eds.), Sea Level Changes--An Integrated Approach. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 42, pp. 183-213.

References

853

Lowe, D.R., 1982. Sediment gravity flows: II. Depositional models with special reference to the deposits of high-density turbidity currents. J. Sediment. Petrol., 52: 279-297. Lowe, D.R., 1983. Restricted shallow-water sedimentation of Early Archaean stromatolitic and evaporitic strata of the Strelley Pool Chert, Pilbara Block, Western Australia. Precambrian Res., 19: 239-283. Lowe, D.R., 1988. Suspended-load fallout rate as an independent variable in the analysis of current structures. Sedimentology, 35: 765-776. Lowe, D.R., 1994a. Archean greenstone-related sedimentary rocks. In: K.C. Condie (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 121-169. Lowe, D.R., 1994b. Abiological origin of described stromatolites older than 3.2 Ga. Geology, 22: 387-390. Lowe, D.R., 1999a. Petrology and sedimentology of cherts and related silicified sedimentary rocks in the Swaziland Supergroup. In: D.R. Lowe and G.R. Byerly (Eds.), Geological Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, pp. 83-114. Lowe, D.R., 1999b. Shallow-water sedimentation of accretionary lapilli-bearing units of the Msauli Chert: Evidence of explosive hydromagmatic komatiitic volcanism. In: D.R. Lowe and G.R. Byerly (Eds.), Geological Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, pp. 213-232. Lowe, D.R. and Byerly, G.R., 1986. Early Archaean silicate spherules of probable impact origin. Geology, 14: 83-86. Lowe, D.R. and Byerly, G.R., 1999a. Stratigraphy of the west-central part of the Barberton Greenstone Belt, South Africa. In: D.R. Lowe and G.R. Byerly (Eds.), Geological Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, pp. 1-36. Lowe, D.R. and Byerly, G.R., 1999b. Geological Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, 319 pp. Lowe, D.R. and Byerly, G.R., 2002. Environmental effects of large impacts in the 3.5-3.2 Ga Barberton Greenstone Belt, South Africa. Abstract, 16th Intern. Sedimentological Congress, Johannesburg, July 2002, p. 230. Lowe, D.R. and Fisher W.G., 1999. Sedimentology, mineralogy, and implications of silicified evaporites in the Kxomberg Formation, Barberton Greenstone Belt, South Africa. In: D.R. Lowe and G.R. Byerly (Eds.), Geologic Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, pp. 167-188. Lowe, D.R., Beukes, N.J., Grotzinger, J.P., Ingersoll, R.V., Kirschvink, J.L., Klein, C., Lambert, I.B. and Veizer, J., 1992. Geological evolution of the Proterozoic Earth. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Earth. Cambridge Univ. Press, Cambridge, pp. 43-79. Lowe, D.R., Byerly, G.R., Ransom, B.L. and Nocita, B.W., 1985. Stratigraphic and sedimentological evidence for repetition in Early Archaean rocks of the Barberton greenstone belt. Precambrian Res., 27: 165-186. Lowe, D.R., Byerly, G.R, Asaro, E and Kyte, ET., 1989. Geological and geochemical record of 3400-million-year-old terrestrial meteorite impacts. Science, 245: 959-962. Lowe, D.R., Byerly, G.R. and Heubeck, C., 1999. Structural divisions and development of the westcentral part of the Barberton Greenstone Belt. In: D.R. Lowe and G.R. Byerly (Eds.), Geologic Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, pp. 37-82. Lowe, D.R., Byerly, G.R., Kyte, ET., Shukolyukov, A., Asaro, E and Krull, A., 2003. Spherule beds 3.47-3.24 billion years old in the Barberton Greenstone Belt, South Africa: a record of large

854

References

meteorite impacts and their influence on early crustal and biological evolution. Astrobiology, 3: 7-48. Lowman, J.P. and Jarvis, G.T., 1996. Continental collisions in wide aspect ratio and high Rayleigh number two-dimensional mantle convection models. J. Geophys. Res., 101: 25,485-25,497. Lucas, S.B., Stem, R.A., Syme, E.C., Reilly, B.A. and Thomas, D.J., 1996. Intraoceanic tectonics and the development of continental crust: 1.92-1.84 Ga evolution of the Flin Flon belt, Canada. Geol. Soc. Amer. Bull., 108: 602-629. Ludden, J., Hubert, C., Barnes, A., Milkereit, B. and Sawyer, E.W., 1993. A three dimensional perspective on the evolution of the Earth's largest Archean greenstone belt: results from the AbitibiGrenville LITHOPROBE project. Lithos, 30: 357-372. Luepke, J.J. and Lyons, T.W., 2001. Pre-Rodinian (Mesoproterozoic) supercontinental rifting along the western margin of Laurentia: geochemical evidence from the Belt-Purcell Supergroup. Precambrian Res., 111: 79-90. Lugmair, G.W. and Galer, S.J.G., 1992. Age and isotopic relationships among the angrites Lewis Cliff 86010 and Angra dos Reis. Geochim. Cosmochim. Acta, 56: 1673-1694. Lugmair, G.W. and Shukolyukov, A., 2001. Early solar system events and timescales. Meteoritics Planet. Sci., 36: 1017-1026. Lugmair, G.W., Shukolyukov, A. and Maclssac, Ch., 1995. The abundance of 6~ in the early solar system. In: M. Busso, R. Gallino and C.M. Raiteri (Eds.), Nucleii and the Cosmos III. Proceedings, Third International Symposium on Nuclear Astrophysics, Assergi, Italy, 1994. Amer. Inst. Nuclear Physics, New York, p. 591. Luhr, J.F., 1997. Extensional tectonics and the diverse primitive volcanic rocks in the western Mexican Volcanic belt. Can. Mineral., 35: 473-500. Lunine, J.L., 1999. Earth: Evolution of a Habitable Planet. Cambridge Univ. Press, Cambridge, 319 pp. Lydon, J.W., 1996. Sedimentary exhalative sulphides (SEDEX). In: O.R. Eckstrand, W.D. Sinclair and R.I. Thorpe (Eds.), Geology of Canadian Mineral Deposit Types. Geol. Surv. Can., Geology of Canada, No. 8., pp. 130-152. Lyons, T.W., 1997. Sulfur isotopic trends and pathways of iron sulfide formation in upper Holocene sediments of the anoxic Black Sea. Geochim. Cosmochim. Acta, 61: 3367-3382. Lyons, T.W. and Gellatly, A.M., 2002. A survey of sulfur isotope trends in Proterozoic SEDEX deposits and inferred links to global ocean chemistry. Abstract, Annu. Mtg. Geol. Soc. Amer., Denver, 2002, 34, p. 517. Lyons, T.W. and Gellatly, A.M., in prep. Sulfur isotope trends in Proterozoic sedimentary exhalative (SEDEX) deposits and links to global ocean chemistry. In: H. Ohmoto and S.E. Kesler (Eds.), Evolution of the Early Atmosphere, Hydrosphere, and Biosphere: Constraints from Ore Deposits. Geol. Soc. Amer., Spec. Paper. Lyons, T.W., Luepke, J.J., Schreiber, M.E. and Zieg, G.A., 2000. Sulfur geochemical constraints on Mesoproterozoic restricted marine deposition: Lower Belt Supergroup, northwestern United States. Geochim. Cosmochim. Acta, 64: 427-437. Lyons, T.W., Gellatly, A.M. and Kah, L.C., 2002. Paleoenvironmental significance of trace sulfate in sedimentary carbonates. Proceedings, 6th Intern. Symp. Geochemist. Earth's Surface (GES-6), Honolulu, Hawaii, pp. 162-165. Lyons, T.W., Weme, J.P., Hollander, D.J. and Murray, R.W., 2003. Contrasting sulfur geochemistry and Fe/A1 and Mo/A1 ratios across the last oxic-to-anoxic transition in the Cariaco Basin, Venezuela. Chem. Geol., 195: 131-157.

References

855

Lyons, W.B., Long, D.T., Gaudette, H.E. and Armstrong, EB., 1984. Calcification of cyanobacterial mats in Solar Lake, Sinai. Geology, 12: 623-626. Maas, R., Kinny, ED., Williams, I.S., Froude, D.O. and Compston, W., 1992. The Earth's oldest known crust: a geochronological and geochemical study of 390(0-4200 Ma old detrital zircons from Mt. Narryer and Jack Hills, Western Australia. Geochim. Cosmochim. Acta, 56:1281-1300. Macgregor, A.M., 1951. Some milestones in the Precambrian of Southern Rhodesia. Trans. Geol. Soc. S. Afr., 54: 27-71. Machado, N., 1990. Timing of collisional events in the Trans-Hudson Orogen: evidence from U-Pb geochronology for the New Quebec Orogen, the Thompson Belt, and the Reindeer Zone (Manitoba and Saskatchewan). In: J.E Lewry and M.R. Stauffer (Eds.), The Early Proterozoic TransHudson Orogen of North America. Geol. Assoc. Can. Spec. Paper, 37, pp. 433-441. Machado, N. and Carneiro, M.A., 1992. U-Pb evidence of Late Archean tectonothermal activity in southern S~o Francisco shield, Brazil. Can. J. Earth Sci., 29: 2341-2346. Machado, N., Brooks, C. and Hart, S.R., 1986. Determination of intial 878r/86Sr and 143Nd]144Nd in primary minerals from mafic and ultramafic rocks: experimental procedure and implications for isotopic characteristics of the Archean mantle under the Abitibi greenstone belt, Canada. Geochim. Cosmochim. Acta, 50: 2335-2348. Machado, N., Schrank, A., Noce, C.M. and Gauthier, G., 1996. Ages of detrital zircon from Archean-Paleoproterozoic sequences: implications for Greenstone Belt setting and evolution of a Transamazonian foreland basin in Quadril~itero Ferrifero, southeast Brazil. Earth Planet. Sci. Lett., 141: 259-276. MacHattie, T.G., 2002. Petrogenesis of the Wathaman Batholith and La Ronge domain plutons in the Reindeer Lake area, Trans-Hudson Orogen, Saskatchewan. M.Sc. Thesis (unpubl.), Memorial University, St. Johns, Newfoundland, 265 pp. Machel, H.G., Krouse, H.R. and Sassen, R., 1995. Products and distinguishing criteria of bacterial and thermochemical sulfate reduction. Appl. Geochem., 10: 373-389. MacKenzie, D.B., 1972. Tidal sand deposits in lower Cretaceous Dakota Group near Denver, Colorado. The Mountain Geologist, 9: 269-277. MacPherson, G.J., Davis, A.M. and Zinner, E.K., 1995. The distribution of Aluminium-26 in the early solar system--a reappraisal. Meteoritics, 30: 365-386. MacRae, N.D., Armitage, A.E., Jones, A.L. and Miller, A.R., 1995. A diamondiferous lamprophyre dike, Gibson Lake area, Northwest Territories. Int. Geol. Rev., 37: 212-229. Madigan, M.T., Martinko, J.M. and Parker, J., 2000. Brock Biology of Microorganisms. PrenticeHall, London, 991 pp. Maher, K.A. and Stevenson, D.J., 1988. Impact frustration of the origin of life. Nature, 331:612-6 14. Maicher, D. and White, J.D.L., 2001. The formation of deep-sea Limu o Pele. Bull. Volcanol., 63: 482-496. Maicher, D., White, J.D.L. and Batiza, R., 2000. Sheet hyaloclastite: density-current deposits of quench and bubble-burst fragments from thin, glassy sheet lava flows, Seamount Six, Eastern Pacific Ocean. Marine Geol., 171: 75-94. Malin, M. and Edgett, K.S., 2000. Evidence for recent groundwater seepage and surface runoff on Mars. Science, 288: 2330-2335. Maloof, A.C., Kellogg, J.B. and Anders, A.M., 2002. Neoproterozoic sand wedges: crack formation in frozen soils under diurnal forcing during a snowball Earth. Earth Planet. Sci. Lett., 204: 1-15. Mandl, G. and Crans, W., 1981. Gravitational gliding in deltas. In: K.R. McClay and N.J. Price (Eds.), Thrust and Nappe Tectonics. Geol. Soc. London. Spec. Publ., 82, pp. 41-54.

856

References

Manikyamba, C., Balaram, V. and Naqvi, S.M., 1993. Geochemical signatures of polygenetic origin of a banded iron formation (BIF) of the Archaean Sandur greenstone belt (schist belt) Karnataka nucleus, India. Precambrian Res., 61: 137-164. Manley, C.R., 1992. Extended cooling and viscous flow of large, hot rhyolite lavas: implications of numerical modeling results. J. Volcanol. Geotherm. Res., 53: 27--46. Manville, V., White, J.D.L., Houghton, B.E and Wilson, C.J.N., 1998. The saturation behaviour of pumice and some sedimentological implications. Sediment. Geol., 119: 5-16. Manville, V.M. and White, J.D.L., 2003. Incipient granular mass flows at the base of sediment-laden floods, and the roles of flow competence and flow capacity in the deposition of stratified bouldery sands. Sediment. Geol., 155: 157-173. Mareschal, J.-C. and West, G.E, 1980. A model for Archean tectonism. Part 2. Numerical models of vertical tectonism in greenstone belts. Can. J. Earth Sci., 17: 60-71. Margulis, L., 1981. Symbiosis in Cell Evolution. Freeman and Co., San Francisco, 419 pp. Marmo, J.S. and Ojakangas, R.W., 1984. Proterozoic glaciogenic deposits, eastern Finland. Geol. Soc. Amer. Bull., 95: 1055-1062. Marsh, J.S., Bowen, M.E, Rogers, N.W. and Bowen, T.B., 1992. Petrogenesis of Late Archaean flood-type basic lavas from the Klipriviersberg Group, Ventersdorp Supergroup, South Africa. J. Petrol., 33: 817-847. Marshak, S., 1999. Deformation style way back when: thoughts on the contrasts between Archaean/Palaeoproterozoic and contemporary orogens. J. Struct. Geol., 21:1175-1182. Marshak, S., Tinkham, D., Alkmim, E, Brueckner, H. and Bomhost, T., 1997. Dome-and-keel provinces formed during Paleoproterozoic orogenic collapse--core complexes, diapirs, or neither?: Examples from the Quadrilatero Ferrifero and Penokean orogen. Geology, 25:415-418. Martel, A.T. and Gibling, M.R., 1991. Wave-dominated lacustrine facies and tectonically controlled cyclicity in the Lower Carboniferous Horton Bluff Formation, Nova Scotia, Canada. In: E Anadon, L. Cabrera and K. Kelts (Eds.), Lacustrine Facies Analysis. IAS Spec. Publ., 13. Blackwell, Oxford, pp. 223-243. Martin, A., Nisbet, E.G., Bickle, M.J. and Orpen, J.L., 1993. Rock units and stratigraphy of the Belingwe greenstone belt: the complexity of the tectonic setting. In: M.J. Bickle and E.G. Nisbet (Eds.), The Geology of the Belingwe Greenstone Belt, Zimbabwe: A Study of the Evolution of Archaean Continental Crust. Geol. Soc. Zimbabwe Spec. Publ. 2. Balkema, Rotterdam, pp. 13-37. Martin, D.M., 1999. Depositional setting and implications of Paleoproterozoic glaciomarine sedimentation in the Hamersley Province, Western Australia. Geol. Soc. Amer. Bull., 111:189-203. Martin, D.M., Powel, C.M. and George, A.D., 2000. Stratigraphic architecture and evolution of the early Palaeoproterozoic McGrath Trough, Western Australia. Precambrian Res., 99: 33-64. Martin, H., 1965. The Precambrian Geology of South West Africa and Namaqualand. Univ. of Cape Town, South Africa, Precambrian Res. Unit, 159 pp. Martin, H., 1986. Effect of steeper Archaean geothermal gradient on geochemistry of subductionzone magmas. Geology, 14: 753-756. Martin, H., 1994. The Archean grey gneisses and the genesis of continental crust. In: K.C. Condie (Ed.), Archaean Crustal Evolution. Developments in Precambrian Geology, 11. Elsevier, Amsterdam, pp. 259-295. Martin, H., 1999. Adakitic magmas: modern analogues of Archaean granitoids. Lithos, 46:411-429. Martin, H. and Moyen, J.E, 2002. Secular changes in tonalite-trondhjemite-granodiorite composition as markers of the progressive cooling Earth. Geology, 30: 319-322. Martin, U. and White, J.D.L., 2001. Depositional and eruptive mechanisms of density current deposits at Allans Beach, Otago Peninsula, New Zealand. In: W.D. McCaffrey, B.C. Kneller

References

857

and J. Peakall (Eds.), Particulate Gravity Currents. IAS Spec. Publ., 31. Blackwell, Oxford, pp. 245-259. Martini, J.E.J., 1990. An Early Proterozoic playa in the Pretoria Group, Transvaal, South Africa. Precambrian Res., 46:341-351. Martini, J.E.J., 1994. A late Archean-Paleoproterozoic (2.6 Ga) paleosol on ultramafics in the eastern Transvaal, South Africa. Precambrian Res., 67: 159-180. Martins-Neto, M.A., 1994. Braidplain sedimentation in a Proterozoic rift basin: the S~o Jo~o da Chapada Formation, southeastern Brazil. Sediment. Geol., 89: 219-239. Martins-Neto, M.A. and Costa, N.B.O.R., 1985. Geologia da borda oeste da Serra do Caraqa. Trabalho Geol6gico, IGC/UFMG, in6dito. Martins-Neto, M.A. and Hercos, C.M., 2002. Sedimentation and tectonic setting of Early Neoproterozoic glacial deposits in south eastern Brazil. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. Blackwell, Oxford, pp. 383-403. Martins-Neto, M.A., Pedrosa-Soares, A.C. and Lima, S.A.A., 2001. Tectono-sedimentary evolution of sedimentary basins from Late Paleoproterozoic to Late Neoproterozoic in the S~o Francisco craton and Araquai fold belt, eastern Brazil. Sediment. Geol., 141/142: 343-370. Maruyama, S., 1994. Plume tectonics. J. Geol. Soc. Japan, 100: 24-49. Masaitis, V.L. and Sysoev, A.G., 1975. Nickel-bearing iron sulphides and native nickel in suevites of Popagai crater (in Russian). Zapiski Vserossiyskogo Mineralogicheskogo Obshchestva, Proceedings of the Russian Mineralogical Society, 101:108-113. Masaitis, V.L., Shafranovsky, G.I., Grieve, R.A.F., Langenhorst, E, Peredery, W.V., Balmasov, E.L., Fedorova, I.G. and Therriault, A., 1997. Discovery of impact diamonds at the Sudbury Structure. LPI Contribution, 922: 33. Masaitis, V.L., Shafranovsky, G.I., Grieve, R.A.E, Langenhorst, E, Peredery, W.V., Therriault, A.M., Balmasov, E.L. and Fedorova, I.G., 1999. Impact diamonds in the suevitic breccias of the Black Member of the Onaping Formation, Sudbury Structure, Ontario, Canada. Geol. Soc. Amer. Spec. Paper, 339, pp. 317-321. Maslov, V.P., 1960. Stromatolites (Genesis, Study Methods, Facies Relation and Geological Significance). Trans. Geol. Inst. Akademia Nauk SSSR, 41. Akademia Nauk SSSR, Moscow, 188 pp. Matin, A. and Mukhopadhyay, D., 1992. Primary sedimentary structures in banded ferruginous quartzite of Sandur Schist Belt, Karnataka. J. Geol. Soc. India, 40:403-413. Matthews, P.E., 1972. Possible Precambrian obduction and plate tectonics in southeastern Africa. Nature, 240: 37-39. Maurette, M., Duprat, J., Engrand, C., Gounelle, M., Kurat, G., Matrajt, G. and Toppani, A., 2000. Accretion of neon, organics, CO2, nitrogen and water from large interplanetary dust particles on the early Earth. Planet. Space Sci., 48:1117-1137. Mawson, D., 1949. The Late Precambrian ice age and glacial record of the Bibliando dome. Journal and Proceedings of the Royal Society of New South Wales, 82:150-174. Maxeiner, R.O., Sibbald, T.I., Slimmon, W.L., Heaman, L.M. and Watters, B.R., 1999. Lithogeochemistry of volcano-plutonic assemblages of the southern Hanson Lake Block and southeastern Glennie Domain, Trans-Hudson Orogen: evidence for a single island arc complex. Can. J. Earth Sci., 36: 209-225. Maynard, J.B., Ritger, S.D. and Sutton, S.J., 1991. Chemistry of sands from the modern Indus River and the Archaean Witwatersrand basin: implications for the composition of the Archaean atmosphere. Geology, 19: 265-268.

858

References

Mazimhaka, R K. and Hendry, H.E., 1984. The Martin Group, Beaverlodge Area. Saskatchewan Geol. Surv., Summary of Investigations 1984, Saskatchewan Energy and Mines, Misc. Report 84-4, pp. 53-62. Mazimhaka, P.K. and Hendry, H.E., 1985. The Martin Group, Tazin Lake, Charlot Point and Jug Bay Areas. Saskatchewan Geol. Surv., Summary of Investigations 1985, Saskatchewan Energy and Mines, Misc. Report 85-4, pp. 67-80. Mazumder, R., Bose, EK. and Sarkar, S. 2000. A commentary on the tectono-sedimentary record of the pre-2.0 Ga continental growth of India vis-a-vis a possible pre-Gondowana Afro-lndian supercontinent. J. Afr. Earth Sci., 30: 201-217. McCall, J., 2001. Tektites in the Geologic Record: Showers of Glass from the Sky. Geol. Soc. London, 256 pp. McClendon, J.H., 1999. The origin of life. Earth Sci. Rev., 47:71-93. McCormick, D.S. and Grotzinger, J.E, 1993. Distinction of marine from alluvial facies in the Paleoproterozoic (1.9 Ga) Burnside Formation, Kilohigok Basin, N.W.T., Canada. J. Sediment. Petrol., 63: 398-419. McCourt, S. and Armstrong, R.A., 1998. SHRIMP U-Pb zircon chronology of granites from the Central Zone, Limpopo Belt, southern Africa: implications for the age of the Limpopo Orogeny. S. Afr. J. Geol., 101: 329-337. McCourt, S. and Vearncombe, J.R., 1987. Shear zones bounding the central zone of the Limpopo Mobile Belt, southern Africa. J. Struct. Geol., 9: 127-137. McCourt, S. and Vearncombe, J.R., 1992. Shear zones of the Limpopo Belt and adjacent granitoidgreenstone terranes: implications for late Archaean collision tectonics in southern Africa. Precambrian Res., 55: 553-570. McCulloch, M.T. and Bennett, V.C., 1994. Progressive growth of the Earth's continental crust and depleted mantle: geochemical constraints. Geochim. Cosmochim. Acta, 58: 4717-4738. McCulloch, M.T. and Bennett, V.C., 1998. Early differentiation of the Earth: an isotopic perspective. In: I. Jackson (Ed.), The Earth's Mantle. Cambridge Univ. Press, Cambridge, pp. 127-138. McDonald, I. and Simonson, B.M., 2002. PGE anomalies detected in two more 2.5-2.6 billion yearold spherule layers in the Hamersley basin of Western Australia. Abstract, 33rd Lunar and Planetary Science Conference, Houston, Texas, 2002, pp. 1250. McDonough, M.R., McNicoll, V.J. and Schetselaar, E.M., 1995. Age and kinematics of crustal shortening and escape in a two sided oblique slip collisional and magmatic orogen, Proterozoic Taltson Magmatic Zone, northeastern Alberta. In: G.M. Ross (Ed.), Alberta Basement Transects Workshop. Lithoprobe Report No. 47. LITHOPROBE Secretariat, University of British Columbia, pp. 264-239. McElhinny, M.W., Taylor, S.R. and Stevenson, D.J., 1978. Limits to the expansion of the Earth, Moon, Mars and Mercury and to changes in the gravitational constant. Nature, 271:316-321. McGlynn, J.C. and Irving, E., 1981. Horizontal motions and rotations in the Canadian Shield during the early Proterozoic. In: F.H.A. Campbell (Ed.), Proterozoic basins of Canada. Geol. Surv. Can. Paper, 81-10, pp. 183-190. McKay, D.S., Gibson, E.K., Thomas-Keprta, K.L., Vali, H., Romanek, C., Clemett, S.J., Chiller, X.D.E, Maechling, C.R. and Zare, R.N., 1996. Search for past life on Mars: possible relic biogenic activity in Martian meteorite ALH84001. Science, 273: 924-930. McKee, E.D., 1979. Ancient sandstones considered to be eolian. In: E.D. McKee (Ed.), A Study of Global Sand Seas. U.S. Geol. Surv. Prof. Paper, 1052, pp. 187-238. McKenzie, D. and Bickle, M.J., 1988. The volume and composition of melt generated by extension of the lithosphere. J. Petrol., 29: 625-679.

References

859

McKenzie, D. and Weiss, N., 1975. Speculations on the thermal and tectonic history of the Earth. Geophys. J. R. Astron. Soc., 42: 131-174. McKinnon, W.B., Zahnle, K.J., Ivanov, B.I. and Melosh, H.J., 1997. Cratering on Venus: models and observations. In: S.W. Brougher, D.M. Hunten and R.J. Phillips (Eds.), Venus II. Univ. Arizona Press, Tucson, Arizona, pp. 969-1014. McLennan, S.M. and Taylor, S.R., 1991. Sedimentary rocks and crustal evolution: tectonic setting and secular trends. J. Geol., 99: 1-21. McMechan, M.E., 2000. Vreeland diamictites--Neoproterozoic glaciogenic slope deposits, Rocky Mountains, northeast British Columbia. Bull. Can. Petrol. Geol., 48: 246-261. McMenamin, M.A.S. and McMenamin, D.L.S. 1990. The Emergence of Animals: The Cambrian Breakthrough. Columbia Univ. Press, New York, 217 pp. McNaughton, N.J., Compston, W. and Barley, M.E., 1993. Constraints on the age of the Warrawoona Group, eastern Pilbara Block, Western Australia. Precambrian Res., 60: 69-98. McNicoll, V.J., Th6riault, R.J. and McDonough, M., 2000. U-Pb and Nd isotopic constraints on the basement to the Paleoproterozoic Taltson magmatic zone, northeastern Alberta. Can. J. Earth Sci., 37: 1575-1596. McPherson, J.G., Shanmugam, G. and Moiola, R.J., 1987. Fan-deltas and braid deltas: varieties of coarse-grained deltas. Geol. Soc. Amer. Bull., 99:331-340. McPhie, J., Doyle, M. and Allen, R., 1993. Volcanic Textures: A Guide to the Interpretation of Textures in Volcanic Rocks. CODES Key Centre, Univ. Tasmania, Hobart, 198 pp. Meert, J.G., 1999. A palaeomagnetic analysis of Cambrian true polar wander. Earth Planet. Sci. Lett., 168: 131-144. Meert, J.G., 2002. Paleomagnetic evidence for a Paleo-Mesoproterozoic supercontinent, Columbia. Gondwana Res., 5:207-216. Meert, J.G. and Powell, C.M., 2001. Introduction to the special volume on the assembly and breakup of Rodinia. Precambrian Res., 110: 1-8. Meert, J.G. and Torsvik, T.H., 2003. The making and unmaking of a supercontinent: Rodinia revisited. Tectonophysics (in review). Meert, J.G. and Van der Voo, R., 1997. The assembly of Gondwana 800--550 Ma. J. Geodyn., 23: 223-235. Meert, J.G., Van der Voo, R., Powell, C.M., Li, Z.X., McElhinny, M.W., Chen, Z. and Symons, D.T.A., 1993. A plate tectonic speed limit? Nature, 363: 216-217. Meert, J.G., Torsvik, T.H., Eide, E.A. and Dahlgren, S., 1998. Tectonic significance of the Fen Complex, S. Norway: constraints from geochronology and palaeomagnetism. J. Geol., 106: 553-564. Meert, J.G., N6d61ec, A., Hall, C., Wingate, M.T.D. and Rakotondrazafy, M. 2001. Palaeomagnetism, geochronology and tectonic implications of the Cambrian-age Carion granite, central Madagascar. Tectonophysics, 340: 1-21. Meibom, A., Desch, S.J., Krot, A.N., Cuzzi, J.N., Petaev, M.I., Wilson, L. and Kell, K., 2000. Largescale thermal events in the solar nebula: evidence from Fe, Ni metal grains in primitive meteorites. Science, 288: 839-841. Meinster, B. and Tickell, S.J., 1975. Precambrian aeolian deposits in the Waterberg Supergroup. Trans. Geol. Soc. S. Afr., 78:191-200. Melezhik, V.A. and Sturt, B.E, 1994. General geology and evolutionary history of the Early Proterozoic Polmak-Pasvik-Pechenga-lmandra-Varzuga-Ust'Ponoy greenstone belt in the northeastern Baltic Shield. Earth Sci. Rev., 36: 205-241.

860

References

Melezhik, V.A., Fallick, A.E., Medvedev, E V. and Makarikan, V.V., 1999. Paleoproterozoic magnesite-stromatolite-dolostone-"red bed" association, Russian Karelia: paleoenvironmental constraints on the positive carbon isotope shift. Norsk Geologisk Tidskrift, 80:163-186. Melezhik, V.A., Fallick, A.E., Medvedev, EV. and Makarikan, V.V., 2000. Extreme ~3Ccarb enrichment in ca. 2.0 Ga magnesite-stromatolite-dolostone-"red beds" association in a global context: a case for the world-wide signal enhanced by local environment. Earth Sci. Rev., 48: 71-120. Mellor, G.L. and Ezer, T., 1995. Sea level variations induced by heating and cooling: an evaluation of the Boussinesq approximation in ocean models. J. Geophys. Res., 100: 20,565-20,577. Mellqvist, C., 13hlander, B., Ski61d, T. and Wikstr6m, A., 1999. The Archaean-Proterozoic palaeoboundary in the Lule~ area, northern Sweden: field and isotope geochemical evidence for a sharp terrane boundary. Precambrian Res., 96: 225-243. Mel' nik, Y.P., 1982. Precambrian Banded Iron-Formations. Elsevier, Amsterdam, 310 pp. Melosh, H.J., 1989. Impact Cratering: A Geologic Process. Oxford Univ. Press, New York, 245 pp. Melosh, H.J., Vickery, A.M. and Tonks, W.B., 1993. Impacts and the early environment and evolution of the terrestrial planets. In: E.H. Levy and J.I. Lunine, (Eds.), Protostars and Planets. Arizona Univ. Press, Tuscon, Arizona, pp. 1339-1370. Melosh, H.J., Ekholm, A.G., Showman, A.P. and Lorenz, R.D., 2002. Is Europa's subsurface water ocean warm? Abstract, Lunar and Planetary Sciences Conference, 33, Houston, Texas, 2002, abstract no. 1824. Mendelson, C.V. and Schopf, J.W., 1992. Proterozoic and selected Early Cambrian microfossils and microfossil-like objects. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere, A Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 865-951. Mengel, J.T., 1965. Precambrian taconite iron formation: a special type of sandstone. Abstracts, Annu. Meeting, Geol. Soc. Amer., Kansas City, Missouri, p. 106. Menzies, M. and Murthy, V.R., 1980. Nd and Sr isotope geochemistry of hydrous mantle nodules and their host alkali basalts: implications for local heterogeneities in metasomatically veined mantle. Earth Planet. Sci. Lett., 46: 323-334. Mertanen, S., Halls, H.C., Vuollo, J.L., Pesonen, L.J. and Stepanov, V.S., 1999. Paleomagnetism of 2.44 Ga mafic dykes in Russian Karelia, eastern Fennoscandian Shield--implications for continental reconstructions. Precambrian Res., 98:197-221. Merz, M., 1992. The biology of carbonate precipitation by cyanobacteria. Facies, 26: 81-102. Merz-Preil3, M., 2000. Calcification in cyanobacteria. In: R.E. Riding and S.M. Awramik (Eds.), Microbial Sediments. Springer-Verlag, Berlin, pp. 50-56. Merz-PreiB, M. and Riding, R., 1999. Cyanobacterial tufa calcification in two freshwater streams: ambient environment, chemical thresholds and biological processes. Sediment. Geol., 126: 103-124. Meunier, A. and Velde, B., 1976. Mineral reactions at grain contacts in early stages of granite weathering. Clay Minerals, 11: 235-240. Meyer, B.S., 1997. Supernova nucleosynthesis. In: T.J. Bernatowicz and E.K. Zinner (Eds.), Astrophysical Implications of the Laboratory Study of Presolar Materials. Proceedings, AIP conference, St. Louis, Missouri, 1997, Amer. Inst. Physics, Woodbury, New York, pp. 155-178. Meyer, C., 1985. Ore metals through geologic history. Science, 227: 1421-1428. Miall, A.D., 1977. A review of the braided river depositional environment. Earth Sci. Rev., 13: 1-62. Miall, A.D., 1978. Lithofacies types and vertical profile models in braided river deposits: a summary. In: A.D. Miall (Ed.), Fluvial Sedimentology. Can. Soc. Petrol. Geol. Mem., 5, pp. 597-604. Miall, A.D., 1985. Architectural-element analysis: a new method of facies analysis applied to fluvial deposits. Earth Sci. Rev., 22:261-308.

References

861

Miall, A.D., 1992. Alluvial deposits. In: R.G. Walker and N.E James (Eds.), Facies Models: Response to Sea-Level Change. Geol. Assoc. Can., Toronto, pp. 119-142. Miall, A.D., 1996. The Geology of Fluvial Deposits: Sedimentary Facies, Basin Analysis and Petroleum Geology. Springer-Verlag, Berlin, 582 pp. Miall, A.D., 1997. The Geology of Stratigraphic Sequences. Springer-Verlag, Berlin, 433 pp. Miall, A.D., 2000. Principles of Sedimentary Basin Analysis, 3rd edn. Springer-Verlag, Berlin, 616 pp. Miall, A.D. and Miall, C.E., 2001. Sequence stratigraphy as a scientific enterprise: the evolution and persistence of conflicting paradigms. Earth Sci. Rev., 54: 321-348. Michaelis, W., Seifert, R., Nauhaus, K., Treude, T., Thiel, V., Blumberg, M., Knittel, K., Gieske, A., Peterknecht, K., Pape, T., Boetius, A., Amann, R., Pimenov, N.V. and Gulin, M.B., 2002. Microbial reefs in the Black Sea fueled by anaerobic oxidation of methane. Science, 297: 1013-1015. Middleton, L.T. and Trujillo, A.E, 1984. Sedimentology and depositional setting of the Upper Proterozoic Scanlan Conglomerate, central Arizona. In: E.H. Koster and R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Can. Soc. Petrol. Geol., Calgary, Alberta, Mem., 10, pp. 189-201. Mil6si, J.E, Ledru, E, Feybesse, J.-L., Dommanget, A. and Marcoux, E., 1992. Early Proterozoic ore deposits and tectonics of the Birimian orogenic belt, West Africa. Precambrian Res., 58: 305-344. Miller, C., Schuster, R., Klotzli, U., Frank, W. and Purtscheller, E, 1999. Post-collisional potassic and ultrapotassic magmatism in SW Tibet: geochemical and Sr-Nd-Pb-O isotopic constraints for mantle source characteristics and petrogenesis. J. Petrol., 40: 1399-1424. Miller, D.J. and Eriksson, K.A., 1997. Late Mississippian prodeltaic rhythmites in the Appalachian basin: a hierarchical record of tidal and climatic periodicities. J. Sediment. Res., B67: 653-660. Miller, R.G. and O'Nions, R.K., 1985. Source of Precambrian and chemical clastic sediments. Nature, 314: 325-330. Miller, R.M., 1983. The Pan-African Damara orogen of South West Africa/Namibia. In: R.M. Miller (Ed.), Evolution of the Damara Orogen of South West Africa/Namibia. Geol. Soc. S. Afr., Johannesburg, Spec. Publ., 11, pp. 431-515. Miller, S.L., 1953. A production of amino acids under possible primitive Earth conditions. Science, 117: 528-529. Miller, S.L. and Urey, H.C., 1959. Organic compound synthesis on the primitive Earth. Science, 130: 245-251. Minter, W.E.L., 1976. Detrital gold, uranium and pyrite concentrations related to sedimentology of the Vaal Reef placer, Witwatersrand, South Africa. Econ. Geol., 71: 157-176. Minter, W.E.L., 1999. Irrefutable detrital origin of Witwatersrand gold and evidence of eolian signatures. Econ. Geol., 94: 665-670. Mints, M.V., 1998. The correlation between the Palaeoproterozoic orogens and granulite belts in the Baltic Shield and North America Craton: a suggested model of Palaeoproterozoic plate tectonics. Gondwana Res., 1: 235-246. Mints, M.V. and Konilov, A.N., 1998. Thermal strucrture of the crust during granulite metamorphism: petrological speculations and geodynamic implications. In: Y. Motoyoshi and K. Shiraishi (Eds.), Origin and Evolution of Continents. International Symposium, Tokyo, 1997. Mem. Nat. Inst. of Polar Research. Spec. Issue, 53, pp. 137-156. Mints, M.V., Glaznev, V.N., Konilov, A.N., Kunina, N.M., Nikitichev, A.E, Raevsky, A.B., Sedikh, Yu.N., Stupak, V.M. and Fonarev, V.I., 1996. The Early Precambrian of the Northeastern Baltic Shield: Paleogeodynamics, Crustal Structure and Evolution. Transactions of GIN RAS, Vol. 503. Scientific World, Moscow, 287 pp. (in Russian).

862

References

Mints, M.V., Konilov, A.N., Krotov, A.V., Vetrin, V.R. and Stupak, V.M., 2000a. Origin and evolution of the Palaeoproterozoic Lapland granulite belt, Baltic Shield: integrated results of geological and petrological studies of exposed granulites and deep crustal xenolites and seismic profiling. Abstracts, 31 th Int. Geol. Congress, Rio de Janeiro, August 2000 (CD-ROM). Mints, M.V., Rundquist, D.V., Larin, A.M., Nenakhov, V.M., Rytsk, E.Yu., Turchenko, S.I. and Chernyshov, N.M., 2000b. Early Precambrian geodynamics and metallogeny. In: N.V. Mezhelovsky, A.E Morozov, G.S. Gusev and V.S. Popov (Eds.), Geodynamics and Metallogeny. Moscow, Ministry of Nat. Res. of the RF, GEOKART, pp. 105-192. Mints, M.V., Berzin, R.G., Zamozhnyaya, N.G., Stupak, V.M., Suleimanov, A.K., Konilov, A.N. and Babarina, I.I., 2001. Collision structures in the Early Precambrian crust of the eastern Baltic Shield: a geological interpretation of seismic data along the 4B Profile. Doklady Earth Sci., 379(5): 515-520 (official English translation). Mints, M.V., Berzin, R.G., Zamozhnyaya, N.G., Zlobin, V.L., Kaulina, T.V., Konilov A.N. and Suleimanov, A.K., 2002. The Palaeoproterozoic collision structures in the Karelia craton deep crust: synthesis of the results of the CDP seismic profiling and interpretation of potential fields and geological data. Doklady Earth Sci., 385(5) (official English translation, in press). Mints, M.V., Konilov, A.N., Krotov, A.V., Kaulina, T.V. and Stupak, V.M. (submitted). The thermal and geodynamic evolution of the Lapland granulite belt: implications on the thermal structure of the lower crust during granulite-facies metamorphism. Precambrian Res. Mitchell, R.H. and Bergman, S.C., 1991. Petrology of Lamproites. Plenum Press, New York, 445 pp. Mitchum, R.M., Jr., Vail, P.R. and Thompson, S., III, 1977. Seismic stratigraphy and global changes of sea-level, part 2: the depositional sequence as a basic unit for stratigraphic analysis. In: C.E. Payton (Ed.), Seismic Stratigraphy--Applications to Hydrocarbon Exploration. Amer. Assoc. Petrol. Geol. Mem. 26, pp. 53-62. Mitrofanov, EP. and Smolkin, V.E (Eds.), 1995. Magmatism, Sedimentogenesis and Geodynamics of the Pechenga Paleorift. Kola Sci. Centre RAS, Apatity, 255 pp. (in Russian). Mitrofanov, EP., Balashov, Yu.A. and Balagansky, V.V., 1991. New geochronological data on lower Precambrian complexes of the Kola Peninsula. In: EP. Mitrofanov and V.V. Balagansky (Eds.), Correlation of Lower Precambrian Formations of the Karelia-Kola Region, USSR and Finland. Geol. Inst. Kola Sci. Centre, Apatity, pp. 12-16. Mitrofanov, EP., Balagansky, V.V., Balashov, Y.A., Gannibal, L.E, Dokuchaeva, V.S., Nerovich, L.I., Radchenko, M.K. and Ryungenen, G.I. 1995. U-Pb age of gabbro-anorthosite massifs in the Lapland Granulite Belt. Nor. Geol. Unders. Spec. Publ., 7: 179-183. Miura, Y., 1994. Two types of shocked quartz and graphite at Barringer impact crater. In: Abstracts, Lunar and Planetary Science Conference, Houston, Texas, 1994, 25, pp. 913-914. Modie, B.N., 1996. Depositional environments of the Meso- to Neoproterozoic Ghanzi-Chobe Belt, Northwest Botswana. J. Afr. Earth Sci., 22: 255-268. Mojzsis, S.J. and Harrison, T.M., 2000. Vestiges of a beginning: clues to the emergent biosphere recorded in the oldest known sedimentary rocks. GSA Today, 10: 1-6. Mojzsis, S.J., Arrhenius, G., McKeegan, K.D., Harrison, T.M., Nutman, A.P. and Friend, C.R.L., 1996. Evidence for life on Earth before 3,800 million years ago. Nature, 384: 55-59. Mojzsis, S.J., Harrison, T.M. and Pidgeon, R.T., 2001. Oxygen-isotope evidence from ancient zircons for liquid water at the Earth's surface 4,300 Myr ago. Nature, 409:178-181. Mojzsis, S.J., Coath, C.D., Greenwood, J.P., McKeegan, K.D. and Harrison, T.M., 2002. Confirmation of mass-independent isotope effects in Archean (2.5-3.8 Ga) sedimentary sulfides as determined by ion microprobe analysis. Abstract, Annu. Meeting Geol. Soc. Amer., Denver, 34, p. 516.

References

863

Mojzsis, S.J., Coath, C.D., Greenwood, J.E, McKeegan, K.D. and Harrison, T.M., 2003. Massindependent isotope effects in Archean (2.5 to 3.8 Ga) sedimentary sulfides determined by ion microprobe analysis. Geochim. Cosmochim. Acta, 67: 1635-1658. Monty, C.L.V., 1977. Evolving concepts on the nature of ecological significance of stromatolites, a review. In: E. FliJgel (Ed.), Fossil Algae. Springer-Verlag, Berlin, pp. 15-35. Mooney, W.D., Laske, G. and Masters, T.G., 1998. Crust 5.1, a global crustal model at 5 by 5 degrees square. J. Geophys. Res., 103: 727-747. Moorbath, S. and Taylor, P.N., 1981. Isotopic evidence for continental growth in the Precambrian. In: A. Kr6ner (Ed.), Precambrian Plate Tectonics. Elsevier, Amsterdam, pp. 491-525. Moorbath, S., O'Nions, R.K. and Pankhurst, R.J., 1973. Early Archaean age for the Isua Iron Formation, West Greenland. Nature, 240:138-139. Moorbath, S., Allaart, J.H., Bridgwater, D. and McGregor, V.R., 1977. Rb-Sr ages for early Archean supracrustal rocks and Amitsoq gneisses. Nature, 270: 4345. Moore, A.G., Cas, R.A.E, Beresford, S.W. and Sone, M., 2000. Geology of an Archaean metakomatiite succession, Tramways, Kambalda Ni province, Western Australia: assessing the extent to which volcanic facies architecture and flow emplacement mechanisms can be constructed. Aust. J. Earth Sci., 47: 659-673. Moore, E.S. and Maynard, J.E., 1929. Solution, transportation and precipitation of iron and silica. Econ. Geol., 24: 272-303; 365-402; 506-527. Moore, J.M., 1989. A comparative study of metamorphosed supracrustal rocks from the western Namaqualand Metamorphic Complex. Precambrian Res. Unit, Univ. Cape Town, Bulletin, 37: 1-370. Moore, J.M., Reid, D.L. and Watkeys, M.K., 1990. The regional setting of the Aggeneys/Gamsberg base metal deposits, Namaqualand, South Africa. In: P.G. Spry and L.T. Bryndzia (Eds.), Regional Metamorphism of Ore Deposits. Science Press, Rotterdam, pp. 77-95. Moores, E.M., 1981. Ancient suture zones within continents. Science, 213: 41-46. Moores, E.M., 1982. Origin and emplacement of ophiolites. Rev. Geophys. Space Phys., 20: 65-68. Moores, E.M., 1986. The Proterozoic ophiolite problem, continental emergence and the Venus connection. Science, 234: 65-68. Moores, E.M., 1993. Neoproterozoic oceanic crustal thinning, emergence of the continents, and origin of the Phanerozoic ecosytem: a model. Geology, 21: 5-8. Moores, E.M., 2002. Pre-I Ga (pre-Rodinian) ophiolites: their tectonic and environmental implications. Geol. Soc. Amer. Bull., 114: 80-95. Moores, E.M., Kellogg, L.H. and Dilek, Y., 2000. Tethyan ophiolites, mantle convection and the tectonic "historical contingency", a resolution of the ophiolite conundrum. Geol. Soc. Amer. Spec. Paper, 349, pp. 3-12. Morelli, R., Ashton, K.E. and Ansdell, K., 2001. A geochemical investigation of the Martin Group igneous rocks, Beaverlodge domain, northwestern Saskatchewan. Summary of Investigations, Saskatchewan Geological Survey, 2001, Vol. 2. Misc. Rep. 2001-4.2, pp. 64-75. Morey, G.B., 1983. Animikie Basin, Lake Superior Region, U.S.A. In: A.E Trendall and R.C. Morris (Eds.), Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 13-68. Morgan, J.W., Horan, M.E, Walker, R.J. and Grossman, J.N., 1995. Rhenium-osmium concentration and isotope systematics in group IIAB iron meteorites. Geochim. Cosmochim. Acta, 59: 2331-2344. Morita, R.Y., 1980. Calcite precipitation by marine bacteria. Geomicrobiology J., 2: 63-82. Morris, R.C., 1993. Genetic modelling for banded iron-formation of the Hamersley Group, Pilbara craton, Western Australia. Precambrian Res., 60: 243-286.

864

References

Morris, R.C. and Horwitz, R.C., 1983. The origin of the iron-formation-rich Hamersley Group of Western Australia--deposition on a platform. Precambrian Res., 21: 273-297. Morris, R.C. and Trendall, A.E, 1988. Discussion: a model for the deposition of the microbanded Precambrian iron formations. Amer. J. Sci., 288: 664-669. Mortensen, J.K., 1993a. U-Pb Geochronology of the eastern Abitibi Subprovince: Part I: Chibougamau-Matagami-Joutel. Can. J. Earth Sci., 30:11-28. Mortensen, J.K., 1993b. U-Pb Geochronology of the eastern Abitibi Subprovince: Part II: NorandaKirkland Lake area. Can. J. Earth Sci., 30:29-41. Morton, R.L., Walker, J.S., Hudak, G.J. and Franklin, J.M., 1991. The early development of an Archean submarine caldera complex with emphasis on the Mattabi ash-flow tuff and its relationship to the Mattabi massive sulfide deposit. Econ. Geol., 86:1002-1011. Moser, D.E., Heaman, L.M., Krogh, T.E. and Hanes, J.A., 1996. Intracrustal extension of an Archaean orogen revealed using single-grain U-Pb zircon geochronology. Tectonics, 15: 1093-1109. Moser, D.E., Flowers, R.M. and Hart, R.J., 2001. Birth of the Kaapvaal tectosphere 3.08 billion years ago. Science, 291: 465-468. Mosher, S., 1998. Tectonic evolution of the southern Laurentian Grenville orogenic belt. Geol. Soc. Amer. Bull., 110: 1357-1375. Mound, J.E., Mitrovica, J.X., Evans, D.A.D. and Kirschvink, J.L., 1999. A sea-level test for inertial interchange true polar wander events. Geophys. J. Int., 136: F5-F10. Moyes, A.B., Krynauw, J.R. and Barton, J.M.J., 1995. The age of the Ritscherflya Supergroup and Borgmassivet Intrusions, Dronning Maud Land, Antarctica. Antarctic Sci., 7: 87-97. Mueller, W., 1991. Volcanism and related slope to shallow marine volcaniclastic sedimentation: an Archean example, Chibougamau, Quebec, Canada. Precambrian Res., 49: 1-22. Mueller, W.U., 2003. A subaqueous eruption model for shallow-water, small volume eruptions: evidence from two Precambrian examples. In: J. White, D. Clague, and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph (in press). Mueller, W.U. and Corcoran, P.L., 1998. Late-orogenic basins in the Archaean Superior Province, Canada: characteristics and inferences. Sediment. Geol., 120:177-203. Mueller, W.U. and Corcoran, P.L., 2001. Volcano-sedimentary processes operating on a marginal continental arc: the Archean Raquette Lake Formation, Slave Province, Canada. Sediment. Geol., 141/142: 169-204. Mueller W.U. and Dimroth, E., 1987. A terrestrial-shallow marine transition in the Archaean Opemisca Group East of Chapis, Quebec. Precambrian Res., 37: 29-55. Mueller, W.U. and Donaldson, J.A., 1992a. Development of sedimentary basins in the Abitibi belt: an overview. Can. J. Earth Sci., 29: 2249-2265. Mueller, W.U. and Donaldson, J.A., 1992b. A felsic feeder dyke swarm formed under the sea: the Archean Hunter Mine Group, south-central Abitibi belt, Quebec, Canada. Bull. Volcanol., 54: 602-610. Mueller, W.U. and Mortensen, J.K., 2002. Age constraints and characteristics of subaqueous volcanic construction, the Archean Hunter Mine Group, Abitibi greenstone belt. Precambrian Res., 115: 119-152. Mueller, W.U. and White, J.D.L., 1992. Felsic fire-fountaining beneath Archean seas: pyroclastic deposits of the 2730 Ma Hunter Mine Group, Quebec, Canada. J. Volcanol. Geotherm. Res., 54: 117-134.

References

865

Mueller, W.U., Chown, E.H., Sharma, K.N.M., Tait, L. and Rocheleau, M., 1989. Paleogeographic evolution of a basement-controlled Archean supracrustal sequence, Chibougamau-Caopatina, Quebec. J. Geol., 97: 399-420. Mueller, W.U., Donaldson, J.A., Dufresne, D. and Rocheleau, M., 1991. The Duparquet formation: sedimentation in a late Archean successor basin, Abitibi greenstone belt, Quebec, Canada. Can. J. Earth Sci., 28: 1394-1406. Mueller, W.U., Chown, E.H. and Potvin, R., 1994a. Substorm wave base felsic hydroclastic deposits in the Archean Lac des Vents volcanic complex, Abitibi Belt, Canada. J. Volcanol. Geotherm. Res., 60: 273-300. Mueller, W.U., Donaldson, J. A. and Doucet, P., 1994b. Volcanic and tectono-plutonic influences on sedimentation in the Archean Kirkland Basin Abitibi greenstone belt, Canada. Precambrian Res., 68:201-230. Mueller, W.U., Daigneault, R., Mortensen, J.K. and Chown, E.H., 1996. Archean terrane docking: upper crust collision tectonics, Abitibi greenstone belt, Quebec, Canada. Tectonophysics, 265: 127-150. Mueller, W.U., Champagne, C., Barnes, S.-J. and Pilote, P., 1999. Physical volcanology of komatiite flows at Spinifex Ridge, La Motte Vassan Formation, Malartic Block, Quebec portion, Abitibi greenstone belt. In: P. Davis (Ed.), Classic Komatiite Localities and Magmatic Fe-Ni-Cu-(PGE) Sulfide Deposits of the Abitibi Greenstone Belt, Ontario-Qu6bec. MERC Guidebook Ser. 1, pp. 45-53. Mueller, W.U., Garde, A. and Stendal, H., 2000. Shallow-water, eruption-fed, mafic pyroclastic deposits along a Paleoproterozoic coastline: Kangerluluk volcano-sedimentary sequence, southeast Greenland. Precambrian Res., 101: 163-192. Mueller, W.U., White, J.D.L. and Corcoran, P.L., 2002a. Fieldtrip guide: dynamic evolution of the submarine Oamaru volcanic complex: syn-eruptive collapse, unconformities, eruption-fed density currents, and backset strata. In: J.D.L. White and B.E Houghton (Eds.), AGU-Chapman Conference, Explosive Subaqueous Volcanism, 21-25th January 2002, Dunedin, New Zealand, p. 28. Mueller, W.U., Corcoran, P.L. and Donaldson, J.A., 2002b. Sedimentology of a tide- and waveinfluenced high-energy Archaean coastline: the Jackson Lake Formation, Slave Province, Canada. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. Blackwell, Oxford, pp. 153-182. Muhling, P.C. and Brakel, A.T., 1985. Geology of the Bangemall Group: the evolution of an intracratonic Proterozoic basin. Western Aust. Geol. Surv. Bull., 128:1-219. Muir, M.D. and Grant, P.R., 1976. Micropaleontological evidence from the Onverwacht Group, South Africa. In: B.E Windley (Ed.), The Early History of the Earth. John Wiley and Sons, London, pp. 595-604. Mukhin, L.M. and Pinenov, K.Yu., 2002. Impact hot spots on the cold surface of the early Earth. Planet. Space Sci., 50: 41-43. Mukhopadhyay, D., 1994. Evolution of the Precambrian terrain of Singhbhum. In: Extended Abstracts, Symposium on Mantle Dynamics and Its Relation to Earthquake and Volcanism, Calcutta, 1994, pp. 45-48. Mialler, M.J., 1982. Selected Climatic Data for a Global Set of Standard Stations for Vegetation Science. Dr. W. Junk Publishers, The Hague, 306 pp. Murakami, T., Utsunomiya, S., Imazu, Y. and Prasad, N., 2001. Direct evidence of late Archean to early Proterozoic anoxic atmosphere from a product of 2.5 Ga old weathering. Earth Planet. Sci. Lett., 184: 523-528.

866

References

Murphy, J.B. and Nance, D., 1992. Mountain belts and the supercontinent cycle. Sci. Amer., 266: 84-91. Mustard, P.S. and Donaldson, J.A., 1987. Early Proterozoic ice-proximal glaciomarine deposition: the lower Gowganda Formation at Cobalt, Ontario, Canada. Geol. Soc. Amer. Bull., 98: 373-387. Myers, J.S., 2001a. Protoliths of the 3.8-3.7 Ga Isua greenstone belt, west Greenland. Precambrian Res., 105: 129-141. Myers, J.S., 200lb. Earth's oldest known volcanic and sedimentary rocks: origin and tectonic evolution of the > 3.8-3.7 Ga Isua Greenstone Belt, southwest Greenland. In: K.E Cassidy, J.M. Dunphy and M.J. Van Kranendonk (Eds.), Extended Abstract Volume, 4th Internat. Archaean Symp., Perth, 2001, pp. 67. Myers, J.S. and Crowley, J.L., 2000. Vestiges of life in the oldest Greenland rocks? A review of early Archean geology in the Godth~bsfjord region, and reappraisal of field evidence for > 3850 Ma life on Akilia. Precambrian Res., 103: 101-124. Myers, J.S. and Swager, C., 1997. The Yilgarn craton. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 640-656. Myers, J.S., Shaw, R.D. and Tyler, I.M., 1994. Proterozoic tectonic evolution of Australia. Geol. Soc. Aust., Abstract, 37:312. Myers, J.S., Shaw, R.D. and Tyler, I.M., 1996. Tectonic evolution of Proterozoic Australia. Tectonics, 15: 1431-1446. Myrow, EM. and Hiscott, R.N., 1991. Shallow-water gravity-flow deposits, Chapel Island Formation, southeast Newfoundland, Canada. Sedimentology, 38: 935-959. Naqvi, S.M., Venkatachala, B.S., Shukla, M., Kumar, B., Natarajan, R. and Sharma, M., 1987. Silicified cyanobacteria from the cherts of Archaean Sandur Schist Belt--Karnataka, India. J. Geol. Soc. India, 29: 535-539. Narbonne, G.M., Kaufman, A.J. and Knoll, A.H., 1994. Integrated chemostratigraphy and biostratigraphy of the upper Windermere Supergroup (Neoproterozoic), Mackenzie Mountains, northwestern Canada. Geol. Soc. Amer. Bull., 106: 1281-1292. Neale, K.L., 1993. Stratigraphy and geochemistry of Neoproterozoic iron formation, South Australia. Ph.D. Thesis (unpubl.), Univ. Western Ontario, 377 pp.Nealson, K.H., 1997. The limits of life on Earth and searching for life on Mars. J. Geophys. Res., 102: 23,675-23,686. Needham, R.S. and De Ross, G.J., 1990. Pine Creek InliermRegional geology and mineralisation. In: EE. Hughes (Ed.), Geology of the Mineral Deposits of Australia and Papua New Guinea. The Australian Institute of Mining and Metallurgy Monographs, 14, pp. 727-737. Nelson, B.K. and DePaolo, D.J., 1985. Rapid production of continental crust 1.7-1.9 b.y. ago: Nd and Sr isotopic evidence from the basement of the North America midcontinent. Geol. Soc. Amer. Bull., 96: 746-754. Nelson, D.R., 1992. Isotopic characteristics of potassic rocks: evidence for the involvement of subducted sediments in magma genesis. Lithos, 28: 403-420. Nelson, D.R., 1997. Evolution of the Archaean granite-greenstone terranes of the Eastern Goldfields, Western Australia: SHRIMP U-Pb zircon constraints. Precambrian Res., 83: 57-81. Nelson, D.R., 1998a. Granite-greenstone crust formation on the Archaean Earth--a consequence of two superimposed processes. Earth Planet. Sci. Lett., 158:109-119. Nelson, D.R., 1998b. Compilation of SHRIMP U-Pb zircon geochronology data, 1997. Geol. Surv. Western Aust., Record 1998/2, 242 pp. Nelson, D.R., 1999. Compilation of geochronology data, 1998. Geol. Surv. Western. Aust., Record 1999/2, 222 pp.

References

867

Nelson, D.R., 2000. Compilation of geochronology data, 1999. Geol. Surv. Western. Aust., Record 2000/2, 251 pp. Nelson, D.R., 2001a. Compilation of geochronology data, 2000. Geol. Surv. Western. Aust., Record 2001/2, 205 pp. Nelson, D.R., 2001 b. An assessment of the determination of depositional ages for Precambrian clastic sedimentary rocks by U-Pb dating of detrital zircons. Sediment. Geol., 141/142: 37-60. Nelson, D.R., 2002. Compilation of geochronology data, 2001. Geol. Surv. Western. Aust., Record 2002/2, 282 pp. Nelson, D.R., Trendall, A.E and Altermann, W., 1999. Chronological correlations between the Pilbara and Kaapvaal cratons. Precambrian Res., 97: 165-189. Nelson, D.R., Robinson, B.W. and Myers, J.S., 2000. Complex geological histories extending for > 4.0 Ga deciphered from xenocryst zircon microstructures. Earth Planet. Sci. Lett., 181: 89-102. Nemec, W. and Steel, R.J., 1984. Alluvial and coastal conglomerates; their significant features and some comments on gravelly mass-flow deposits. In: E.H. Koster and R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Can. Soc. Petrol. Geol. Mem., 10, pp. 1-31. Nerovich, L.I., 1999. Petrology and geochronology of anorthosites from the Lapland granulite belt. Ph.D. Thesis (unpubl.), Geol. Inst. of the Kola Sci. Centre RAS, Apatity, 23 pp. (in Russian). Nesbitt, H.W. and Markovics, G., 1997. Weathering of granodioritic crust, long-term storage of elements in weathering profiles, and petrogenesis of siliciclastic sediments. Geochim. Cosmochim. Acta, 61: 1653-1670. Nesbitt, H.W. and Wilson, R.E., 1992. Recent chemical weathering of basalts. Amer. J. Sci., 292: 740-777. Nesbitt, H.W. and Young, G.M., 1982. Early Proterozoic climates and plate motions inferred from major element chemistry of lutites. Nature, 299:715-717. Nesbitt, H.W. and Young, G.M., 1984. Prediction of some weathering trends of plutonic and volcanic rocks based on thermodynamic and kinetic considerations. Geochim. Cosmochim. Acta, 48: 1523-1534. Nesbitt. H.W. and Young, G.M., 1989. Formation and diagenesis of weathering profiles. J. Geol., 97: 129-147. Nesbitt, H.W. and Young, G.M., 1997. Sedimentation in the Venezuelan basin, circulation in the Caribbean Sea, and onset of Northern Hemisphere glaciation. J. Geol., 105:531-544. Nesbitt, R.W., 1971. Skeletal crystal forms in the ultramafic rocks of the Yilgarn Block, Western Australia: evidence for an Archaean ultramafic liquid. Geol. Soc. Aust., 3:331-347. Nesbitt, R.W., Sun, S.S. and Purvis, A.C., 1979. Komatiites: geochemistry and genesis. Can. Mineral., 17:165-186. Neumann, A.C., Gebelein, C.D. and Scoffin, T.P., 1970. The composition, structure and erodability of subtidal mats, Abaco, Bahamas. J. Sediment. Petrol., 40: 274-297. Newman, M.J. and Rood, R.T., 1977. Implications of solar evolution for the Earth's early atmosphere. Science, 198: 1035-1067. Neymark, L.A., Larin, A.M., Ovchinnikova, G.V. and Yakovleva, S.Z., 1992. U/Pb age of Dzhugdzhur anorthosites. Trans. Russian Acad. Sci., 323 (4-6): 514-518. Neymark, L.A., Nemchin, A.A., Vetrin, V.R. and Salnikova, E.V., 1993. Sm-Nd and Pb-Pb isotope systematics in deep crustal xenoliths and explosive pipes of the southern part of the Kola Peninsula. Dokl. Akad. Nauk., 326(6): 781-784 (in Russian). Nielson, S.G., Baker, J.A. and Krogsstad, E.J., 2002. Petrogenesis of an early Archaean (3.4 Ga) norite dyke, Isua, West Greenland: evidence for early Archaean crustal recycling? Precambrian Res., 118: 133-148.

868

References

Nijman, W., 1999. Growth structures: examples of integrated sedimentological and structuralgeological Basin Analysis. Geologica Ultraiectina, 173:45 (CD). Nijman, W. and de Vries, S.T., 2001. Collapse structures in the early Earth and possible planetary equivalents. Abstract, 32nd Lunar and Planetary Science Conference, Lunar and Planetary Institute, Houston, abstract no. 1289 (CD-ROM). Nijman, W., de Bruijne, C.H. and Valkering, M., 1998a. Growth fault control of Early Archaean cherts, barite mounds, and chert-barite veins, North Pole Dome, Eastern Pilbara, W. Australia. Precambrian Res., 88: 25-52. (Re-edition in Precambrian Res., 1999, 95: 247-274.) Nijman, W., Willigers, B.J.A. and Krikke, A., 1998b. Tensile and compressive growth structures: relationships between sedimentation, deformation and granite intrusion in the Archaean Coppin Gap greenstone belt, Eastern Pilbara, W. Australia. Precambrian Res., 88: 83-108. (Re-edition in Precambrian Res., 1999, 95: 277-302.) Nijman, W., de Bruijne, K.H. and Valkering, M. 1999. Growth fault control of Early Archaean cherts, barite mounds and chert-barite veins, North Pole Dome, Eastern Pilbara, Western Australia. Precambrian Res., 95: 247-274. Nikitina, L.P., Lokhov, K.I., Levsky, L.K., Beletsky, B.V., Zhuralev, V.A., Lepekhina, E.N., Antonov, A.V. and Putintzeva, E.V., 1999. Proterozoic lamproite magmatism of the east part of the Baltic Shield. Vestnik Sankt Petersburgskogo Universiteta Seriya Geologiya i Geografiya, 8(2): 11-25. Nio, S.D., and Yang, C., 1991. Diagnostic attributes of clastic tidal deposits: a review. In: D.G. Smith, G.E. Reinson, B.A. Zaitlin and R.A. Rahmani (Eds.), Clastic Tidal Sedimentology. Can. Soc. Petrol. Geologists, Calgary, Mem., 16, pp. 3-28. Nironen, M., 1997. The Svecofennian orogen: a tectonic model. Precambrian Res., 86: 21-44. Nisbet, E.G., 1982. The tectonic setting and petrogenesis of komatiites. In: N.T. Arndt and E. Nisbet (Eds.), Komatiites. George Allen and Unwin, London, pp. 501-526. Nisbet, E.G., 1987. The Young Earth: An Introduction to Archean Geology. George Allen and Unwin, Boston, 402 pp. Nisbet, E.G., 1995. Archaean ecology: a review of evidence for the early development of bacterial biomes and speculation on the development of a global-scale biosphere. In: M.P. Coward and A.C. Ries (Eds.), Early Precambrian Processes. Geol. Soc., London, pp. 27-52. Nisbet, E.G. and Sleep, N.H., 2001. The habitat and nature of early life. Nature, 409: 1083-1091. Nisbet, E.G., Arndt, N.T., Bickle, M.J., Cameron, W.E., Chauvel, C., Cheadle, M., Hegner, E., Kyser, T.K., Martin, A., Renner, R. and Roedder, E., 1987. Uniquely fresh 2.7 Ga komatiites from the Belingwe greenstone belt. Geology, 15:1147-1150. Nisbet, E.G. Cheadle, M.J., Arndt, N.T. and Bickle, M.J., 1993a. Constraining the potential temperature of the Archean mantle: a review of evidence from komatiites. Lithos, 30:291-307. Nisbet, E.G., Martin, A., Bickle, M.J. and Orpen, J.L., 1993b. The Ngezu Group: komatiites, basalts and stromatolites on continental crust. In: M.J. Bickle and E.G. Nisbet (Eds.), The Belingwe Greenstone Belt, Zimbabwe: A Study of Archean Continental Crust. Geol. Soc. Zimbabwe Spec. Publ., 2. Balkema, Rotterdam, pp. 121-165. Nixon, P.H., Pearson, D.G. and Condliffe, E., 1999. Harzburgites with spinifex texture from SE Spain; petrological and geochemical constraints on their origin. Proceedings, International Kimberlite Conference, Rondebosch, South Africa, 1998, 7, Vol. 2, pp. 605-615. Nockolds, S.R., 1954. Average chemical compositions of some igneous rocks. Geol. Soc. Amer. Bull., 65: 1007-1032.

References

869

Noffke, N., 1998. Multidirected ripple marks arising from bacterial stabilization counteracting physical reworking in modern sandy deposits (Mellum Island, southern North Sea). Geology, 26: 879-882. Noffke, N., Gerdes, G., Klenke, Th. and Knambein, W.E., 1997. A microscopic sedimentary succession indicating the presence of microbial mats in siliciclastic tidal flats. Sediment. Geol., 110: 1-6. Noldart A.J. and Wyatt, J.D., 1962. The geology of a portion of the Pilbara Goldfield, covering the Marble Bar and Nullagine 4-mile map sheets. West. Aust. Geol. Surv. Bull., 115: 199. Nowicki, T.E., Frimmel, H.E. and Waters, D.J., 1995. The occurrence of osumilite in pelitic granulites of the Namaqualand Metamorphic Complex, South Africa. S. Afr. J. Geol., 98: 191-201. Nozhkin, A.D., 1999. Early Proterozoic margin-continental complexes of the Angar fold belt and their metallogenic features. Geology and Geophysics, 40(11): 1524-1544 (in Russian). Nummedal, D., Riley, G.W. and Templet, P.L., 1993. High-resolution sequence architecture: a chronostratigraphic model based on equilibrium profile studies. In: H.W. Posamentier, C.P. Summerhayes, B.U. Haq and G.P. Allen (Eds.), Sequence Stratigraphy and Facies Associations. IAS Spec. Publ., 18. Blackwell, Oxford, pp. 55-68. Nutman, A.P., 1986. The early Archean to Proterozoic history of the Isukasia area, southern West Greenland. Geol. Survey Greenland Bull., 154: 80. Nutman, A.P., 1997. The Greenland sector of the North Atlantic Craton. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 665-674. Nutman, A.P. and Bridgwater, D., 1986. Early Archean Amitsoq tonalites and granites of the Isukasia area, southern West Greenland: development of the oldest-known sial. Contrib. Mineral. Petrol., 94: 137-148. Nutman, A.P., Allaart, J.H., Bridgwater, D., Dimroth, E. and Rosing, M., 1984. Stratigraphic and geochemical evidence for the depositional environment of the early Archaean Isua supracrustal belt, southern West Greenland. Precambrian Res., 25: 365-396. Nutman, A.P., McGregor, V.R., Friend, C.L., Bennett, V.C. and Kinny, P.D., 1996. The Itsaq Gneiss Complex of southern West Greenland; the world's most extensive record of early crustal evolution (3900-3600 Ma). Precambrian Res., 78: 1-39. Nutman, A.P., Kalsbeek, E, Marker, M., van Gool, J.A.M. and Bridgwater, D., 1999. U-Pb zircon ages of Kang~miut dykes and detrital zircons in metasediments in the Palaeoproterozoic Nagssugtoqidian Orogen (West Greenland). Clues to the precollisional history of the orogen. Precambrian Res., 93: 87-104. Nutman, A.P., Friend, C.L. and Bennett, V.C., 2001. Review of the oldest (4400-3600 Ma) geological and mineralogical record: glimpses of the beginning. Episodes, 24(2): 93-101. Nutman, A.P., Friend, C.R.L. and Bennett, V.C., 2002. Evidence for 3650-3600 Ma assembly of the northern end of the ltsaq Gneiss Complex, Greenland: implication for early Archaean tectonics. 9Tectonics, 21 : 1005, doi: 10.1029 12000TC001203. Nyblade, A.A. and Pollack, H.N., 1993. A global analysis of heat flow from Precambrian terrains: implications for the thermal structure of Archean and Proterozoic lithosphere. J. Geophys. Res., 98: 12,207-12,218. Nyquist, L.E., Weismann, H., Bansal, B., Shih, C.-Y., Keith, J.E. and Harper, C.L., 1995. 146Sm-142Nd formation interval for the lunar mantle. Geochim. Cosmochim. Acta, 59: 2817-2837. Nyquist, L.E., Lindstrom, D., Shih, C.-Y., Weismann, H., Mittlefehldt, D., Wentworth, S. and Martinez, R., 1997. Mn-Cr isotopic systematics of chondrules from the Bishunpur and Chainpur meteorites. Lunar Planet. Sci., 27: 1093-1094.

References

870

Nyquist, L., Shih, C.-Y., Weismann, H., Reese, W., Ulyanov, A.A. and Takeda, H., 1999. Toward a Mn-Cr timescale of the early solar system. Lunar Planet. Sci, 30: abstract #1604. O'Brien, N.R., 1990. Significance of lamination in the Toarcian (Lower Jurassic) shales from Yorkshire, Great Britain. Sediment. Geol., 67: 25-34. Occhipinti, S.A., Grey, K., Pirajno, E, Adamides, N.G., Bagas, L., Dawes, P. and Le Blanc Smith, G. 1997. Stratigraphic revision of Palaeoproterozoic rocks of the Yerrida, Bryah and Padbury Basins (former Glengarry Basin). West. Aust. Geol. Surv. Record, 1997/3, 57 pp. Occhipinti, S.A., Swager, C.P. and Pirajno, E, 1998. Structural-metamorphic evolution of the Palaeoproterozoic Bryah and Padbury Groups during the Capricorn orogeny, Western Australia. Precambrian Res., 90: 141-158. O'Dea, M.G., Lister, G.S., MacCready, T., Betts, P.G., Oliver, N.H.S., Pound, K.S., Huang, W. and Valenta, R.K., 1997. Geodynamic evolution of the Proterozoic Mount Isa terrain. In: J.P. Burg and M. Ford (Eds.), Orogeny Through Time. Geol. Soc. Spec. Publ., 121, pp. 99-122. Odette, J.B., 1969. Jadeite: shock induced formation from oligoclase: Ries crater, Germany. Science, 165: 1005-1008. Oglesby, R.J. and Ogg, J.G., 1998. The effect of large fluctuations in obliquity on climates of the Late Proterozoic. Paleoclimates, 2:293-316. O'Hara, M.J., 2000. Flood basalts, basalt floods or topless Bushvelds? Lunar petrogenesis revisited. J. Petrol., 41: 1545-1651. Ohmoto, H., 1992. Biogeochemistry of sulfur and the mechanisms of sulfide-sulfate mineralization in Archean oceans. In: M. Schidlowski (Ed.), Early Organic Evolution: Implications for Mineral and Energy Resources. Springer-Verlag, Heidelberg, pp. 378-397. Ohmoto, H., 1993. The banded iron formations in the Hamersley Basin, Australia: products of the oxygen-rich Archean atmosphere? Abstract, Geol. Soc. Amer. Annu. Meeting, Boston, p. A-89. Ohmoto, H., 1996a. Formation of volcanogenic massive sulfide deposits: the Kuroko perspective. Ore Geol. Rev., 10: 135-177. Ohmoto, H., 1996b. Evidence in pre-2.2 Ga paleosols for the early evolution of atmospheric oxygen and terrestrial biota. Geology, 24:1135-1138. Ohmoto, H., 1997. When did the Earth's atmosphere become oxic? The Geochemical News, 93: 12-13; 26-28. Ohmoto, H., 1999. Redox state of the Archean atmosphere: evidence from detrital heavy minerals in ca. 3250-2750 Ma sandstones from the Pilbara Craton, Australia: Comment. Geology, 27: 1150-1151.

Ohmoto, H., 2002. Banded iron formations as guides for the history of the lithosphere, atmosphere, and biosphere. Abstract, Geol. Soc. Amer., Vol. 34(6), pp. 125-129. Ohmoto, H. and Felder, R.E, 1987. Bacterial activity in the warmer, sulphate-bearing, Archean oceans. Nature, 328: 244-246. Ohmoto, H. and Goldhaber, M.B., 1997. Sulfur and carbon isotopes. In: H.L. Barnes (Ed.), Geochemistry of Hydrothermal Ore Deposits, 3rd edn. John Wiley and Sons, New York, pp. 517-611. Ohmoto, H., Kakegawa, T. and Lowe, D.R., 1993.3.4-billion-year-old biogenic pyrites from Barberton, South Africa: Sulfur isotope evidence. Science, 262: 555-557. Ohmoto, H., Yamaguchi, K.E., Ono, S., Farquhar, J., Bao, H., Thiemens, M.H., Hu, G. and Rumble, D., 2001. Questions regarding Precambrian sulfur isotope fractionation. Science, 292: 1959a. Ohta, H., Maruyama, S., Takahashi, E., Watanabe, Y. and Kato, Y., 1996. Field occurrence, geochemistry and petrogenesis of the Archean Mid-Oceanic Ridge Basalts (AMORBs) of the Cleaverville area, Pilbara Craton, Western Australia. Lithos, 37:199-221.

References

871

Ohtani, E., Kawabe, I., Moriyama, J. and Nagata, Y., 1989. Partitioning of elements between majorite garnet and melt and implications for petrogenesis of komatiite. Contrib. Mineral. Petrol., 103: 263-269. Ojakangas, R.W., 1985. Review of Archean clastic sedimentation, Canadian Shield: major felsic volcanic contributions to turbidite and alluvial fan-fluvial facies associations. In: L.D. Ayres, EC. Thurston, K.D. Card and W. Weber (Eds.), Evolution of Archean Supracrustal Sequences. Geol. Assoc. Can. Spec. Paper, 28, pp. 23-47. Ojakangas, G.W. and Stevenson, D.J., 1989. Thermal state of an ice shell on Europa. Icarus, 81: 220-241. Ojakangas, R.W., Marmo, J.S. and Heiskanen, K.I., 2001a. Basin evolution of the Paleoproterozoic Karelian Supergroup of the Fennoscandian (Baltic) Shield. Sediment. Geol., 141/142: 255-285. Ojakangas, R.W., Morey, G.B. and Southwick, D.L., 2001 b. Paleoproterozoic basin development and sedimentation in the Lake Superior region, North America. Sediment. Geol., 141/142:319-341. Ojakangas, R.W., Morey, G.B. and Green, J.C., 2001c. The Mesoproterozoic Midcontinent Rift System, Lake Superior region, USA. Sediment. Geol., 141/142: 421-442. Olafsson, M. and Eggler, D.H., 1983. Phase relations of amphibole, amphibole-carbonate, and phlogopite-carbonate peridotite: petrologic constraints on the asthenosphere. Earth Planet. Sci. Lett., 64: 305-315. Oliver, N.H.S. and Cawood, EA., 2001. Early tectonic dewatering and brecciation on the overturned sequence at Marble Bar, Pilbara Craton, Western Australia: dome-related or not? Precambrian Res., 105: 1-15. Olsen, EE., Remington, C.L., Corliss, B. and Thomson, K.S., 1978. Change in Late Triassic lacustrine communities. Science, 201: 729-733. O'Nions, K., 1992. The continents. In: G. Brown, C. Hawkesworth and C. Wilson (Eds.), Understanding the Earth. Cambridge Univ. Press, Cambridge, pp. 145-163. O'Nions, R.K., Evernsen, N.M., Hamilton, EJ. and Carter, S.R., 1978. Melting of the mantle past and present: isotopic and trace element evidence. Philos. Trans. Roy. Soc. London A, 258: 547-559. Oparin, A.I., 1938. Origin of Life. Macmillan, New York, 270 pp. Orphan, V.J., House, C.H., Hinrichs, K-U., McKeegan, K.D. and DeLong, E.E, 2001. Methaneconsuming Archaea revealed by directly coupled isotopic and phylogenetic analysis. Science, 293: 484-487. Ortoleva, EJ., 1994. Geochemical Self-Organization. Oxford Univ. Press, Oxford, 411 pp. Orton, G.J., 1996. Volcanic environments. In: H.G Reading (Ed.), Sedimentary Environments: Processes, Facies and Stratigraphy. Blackwell, Oxford, pp. 485-567. Ott, U., 2001. Presolar grains in meteorites: an overview and some implications. Planet. Space Sci., 49: 763-767. Owen, T., Cess, R.D. and Ramanathan, V., 1979. Early Earth: an enhanced carbon dioxide greenhouse to compensate for reduced solar luminosity. Nature, 277: 640--642. Owen, T.C. and Bar-Nun, A., 2000. Volatile contributions from icy planetesimals. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Arizona Univ. Press, Tucson, Arizona, pp. 459-471. Pacquet, A. and McNamara, S., 1985. The study of the basal Athabasca succession in the D, E, L, F, and S areas of the Carswell Structure. In: R. Lain6, D. Alonso and M. Svab (Eds.), The Carswell Structure Uranium Deposits, Saskatchewan. Geol. Assoc. Can. Spec. Paper, 29, pp. 81-86. Paerl, H.W., Steppe, T.E and Reid, R.E, 2001. Bacterially mediated precipitation in marine stromatolites. Environ. Microbiol., 3: 123-130.

872

References

Page, N.J., 1981. The Precambrian diamictite below the base of the Stillwater Complex, Montana. In: M.J. Hambrey and W.B. Harland (Eds.), Earth's pre-Pleistocene glacial record. Cambridge Univ. Press, Cambridge, pp. 821-823. Page, N.J. and Koski, R.A., 1973. A Precambrian diamictite below the base of the Stillwater Complex, southwestern Montana. U.S. Geol. Surv. J. Res., 1:403-414. Page, R.W. and Williams, I.S., 1988. Age of the Barramundi Orogeny in northern Australia by means of ion microprobe and conventional U-Pb zircon studies. Precambrian Res., 40/41:21-36. Pagel, M., 1975a. Cadre g6ologique des gisements d'uranium dans la structure Carswell (Saskatchewan, Canada). l~tude des phases fluides. 3~ Cycle Docteur de Sp6cialit6, Universit6 de Nancy, France, pp. 1-157. Pagel, M., 1975b. D6termination des conditions physicochimiques de la silicifications diag6n6tiques des gr~s Athabasca (Canada) au moyen des inclusions fluides. Comptes Rendus de l'Acad6mie des Sciences Paris S6ries D, 280: 2301-2304. 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 for continental accretion. Precambrian Res., 38: 1-54. Pappalardo, R.T., Sherman, N.D., Head, J.W., Collins, G.G., Greeley, R., Klemaszewski, J., Sullivan, R., Phillips, C.B., Ewan, A.S. and Geisler, P.E., 1998. Distribution of mottled terrain on Europa: a possible link to nonsynchronous rotation stresses. Abstract, Lunar and Planetary Science Conference, 29, Houston, Texas, 1998, abstract no. 1923. Park, A.F., 1991. Continental growth by accretion: a tectonostratigraphic terrane analysis of the evolution of the western and central Baltic Shield, 2.50-1.75 Ga. Geol. Soc. Amer. Bull., 103: 522-537. Park, J.K., 1997. Paleomagnetic evidence of low-latitude glaciation during deposition of the Neoproterozoic Rapitan Group, Mackenzie Mountains, N.W.T., Canada. Can. J. Earth Sci., 34: 34--49. Park, J.K., Buchan, K.L. and Harlan, S.S., 1995. A proposed giant radiating dyke swarm fragmented by the separation of Laurentia and Australia based on palaeomagnetism of the ca. 780 Ma mafic intrusions in western North America. Earth Planet. Sci. Lett., 132:129-139. Park, R. and Epstein, S., 1963. Carbon isotopic fractionation during photosynthesis. Geochim. Cosmochim. Acta, 21: 11 0-115. Parkes, R.J., Cragg, B.A. and Wellsbury, P., 2000. Recent studies on bacterial populations and processes in subseafloor sediments: a review. Hydrogeol. J., 8:11-28. Parman, S.W., Dann, J.C., Grove, T.L. and de Wit, M.J., 1997. Emplacement conditions of komatiite magmas from the 3.49 Ga Komati Formation, Barberton greenstone belt, South Africa. Earth Planet. Sci. Lett., 150: 303-323. Pasteris, J.D. and Wopenka, B., 1991. Raman spectra of graphite as indicators of degree of metamorphism. Can. Mineral., 2: 1-9. Pasteris, J.D., Wopenka, B., Schopf, J.W., Kudryavtsev, A.B., Agresti, D.G., Wdowiak, T.J. and Czaja, A.D., 2002. Images of the Earth's earliest fossils? Nature, 420: 476-477. Patchett, P.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. Patifio Douce, A.E. and Beard, J.S., 1995. Dehydration melting of biotite gneiss and quartz amphibolite from 3 to 15 kbar. J. Petrol., 36: 707-738. Pattullo, J.G. 1966. Mean sea level. In: R.W. Fairbridge (Ed.), The Encyclopedia of Oceanography. Reinhold, New York, pp. 475-479.

References

873

Paulsson, O. and Austrheim, H., 2001. 500 Ma migmatitisation event in Gjelsvikfjella, Dronning Maud Land--implications for continent-continent collision. Abstracts, EUG XI Conference, Strasbourg, 2001, pp. 375. Pavlov, A.A. and Kasting, J.E, 2002. Mass-independent fractionation of sulfur isotopes in Archean sediments: strong evidence for an anoxic Archean atmosphere. Astrobiology, 2:27-41. Pavlov, A.A., Brown, L.L. and Kasting, J.F., 2001a. UB shielding of NH3 and O2 by organic hazes in the Archean atmosphere. J. Geophys. Res., 106: 23,267-23,287. Pavlov, A.A., Kasting, J.E, Eigenbrode, J.L. and Freeman, K.H., 200lb. Organic haze in Earth's early atmosphere: source of low-13C late Archean kerogens? Geology, 29: 1003-1006. Pavlov, A.A., Kasting, J.E, Brown, L.L., Rages, K.A. and Freedman, R., 200 lc. Greenhouse warming by CH4 in the atmosphere of early Earth. J. Geophys. Res., 105:11,981-11,990. Pavlov, A.A., Hurtgen, M.T., Kasting, J.E and Arthur, M.A., 2003. Methane-rich Proterozoic atmosphere? Geology, 31: 87-90. Paytan, A., Kastner, M., Campbell, D. and Thiemens, M.H., 1998. Sulfur isotopic composition of Cenozoic seawater sulfate. Science, 282: 1459-1462. Payton, C.E., 1977. Seismic StratigraphymApplications to Hydrocarbon Exploration. Amer. Assoc. Petrol. Geol. Mem. 26, 516 pp. Pearce, J.A. and Peate, D.W., 1995. Tectonic implications of the composition of volcanic arc magmas. Ann. Rev. Earth Planet. Sci., 23:251-285. Pearce, J.A., Kempton, P.D., Nowell, G.M. and Noble, S.R., 1999. Hf-Nd element and isotope perspective on the nature of provenance of mantle and subduction components in Western Pacific arc-basin systems. J. Petrol., 40:1579-1611. Pearson, D.G., 1999. The age of continental roots. Lithos, 48: 171-194. Peccerillo, A., 1999. Multiple mantle metasomatism in central-southern Italy: geochemical effects, timing and geodynamic implications. Geology, 27:315-318. Peck, W.H., Valley, J.W., Wilde, S.W. and Graham, C.M., 2001. Oxygen isotope ratios and rare earth elements in 3.3 to 4.4 Ga zircons: ion microprobe evidence for high ~ 180 continental crust and oceans in the Early Archaean. Geochim. Cosmochim. Acta, 65: 4215-4229. P61oquin, A.S., Potvin, R., Paradis, S., Lafl~che, M.R., Verpaelst, P. and Gibson, H.L., 1990. The Blake River Group, Rouyn-Noranda area, Quebec: a stratigraphic synthesis. In: M. Rive, P. Verpaelst, Y. Gagnon, J.M. Lulin, G. Riverin and A. Simard (Eds.), La Ceinture polym6tallique du Nord-Ouest qu6b6cois. Can. Inst. Min. Metall. Spec. Vol., 43, pp. 107-118. Peltier, W.R., Butler, S. and Solheim, L.P., 1997. The influence of phase transformations on mantle mixing and plate tectonics. In: D.J. Crossley (Ed.), Earth's Deep Interior. Gordon and Breach, Amsterdam, pp. 405-430. Peltonen, P., Kontinen, A. and Huhma, H., 1996. Petrology and geochemistry of metabasalts from the 1.95 Ga Jormua ophiolite, northeastern Finland. J. Petrol., 3: 1359-1383. Pentecost, A. and Bauld, J., 1988. Nucleation of calcite on the sheaths of cyanobacteria using a simple diffusion cell. Geomicrobiology J., 6: 129-135. Pepin, R.O., 1998. Isotopic evidence for a solar Argon component in the Earth's mantle. Nature, 394: 664-667. Pe-Piper, G., 1998. The nature of Triassic extension-related magmatism in Greece: evidence from Nd and Pb isotope geochemistry. Geol. Mag., 135(3): 331-348. Peter, J.M., 2001. Ancient iron-rich metalliferous sediments (iron formations): their genesis and use in the exploration for stratiform base metal sulphide deposits, with examples from the Bathurst Mining Camp. In: D.R. Lentz (Ed.), Geochemistry of Sediments and Sedimentary Rocks:

874

References

Secular Evolutionary Considerations to Mineral Deposit-Forming Environments. Geol. Assoc. Can., pp. 1-38. Peterson, D.W., Holcomb, R.T., Tilling, R.I. and Christiansen, R.L., 1994. Development of lava tubes in the light of observations at Mauna Ulu, Kilauea Volcano, Hawaii. Bull. Volcanol., 56: 343-360. Peterson, T.D., 1994. Early Proterozoic ultrapotassic volcanism of the Keewatin Hinterland, Canada. In: H.O.A. Meyer and O.H. Leonardos (Eds.), Proceedings, 5th International Kimberlite Conference, Vol. 1: Kimberlites, related rocks and mantle xenoliths. Campanhia de Pesquisa de Recursos Minerais, Brasilia, pp. 221-235. Peterson, T.D. and van Breemen, O., 1999. Review and progress report of Proterozoic granitoid rock of the western Churchill Province, Northwest Territories (Nunavut). Current Research 1999-C, Geol. Surv. Can., pp. 119-127. Peterson, T.D. and LeCheminant, A.N., 1993. Glimmerite xenoliths in Early Proterozoic ultrapotassic rocks from the Churchill Province. Can. Mineral., 31:801-819. Peterson, T.D. and LeCheminant, A.N., 1996. Ultrapotassic rocks of the Dubawnt Supergroup, district of Keewatin, N.W.T. In: A.N. LeCheminant, D.G. Richardson, R.N.W. DiLabio and K.A. Richardson (Eds.), Searching for Diamonds in Canada. Geol. Surv. Can., Ottawa, Open File, 3228, pp. 97-100. Peterson, T.D. and Rainbird, R.H., 1990. Tectonic and petrological significance of regional lamproite-minette volcanism in the Thelon and Trans-Hudson hinterlands, Northwest Territories. Current Research 90-1C, Geol. Surv. Can., pp. 69-79. Peterson, T.D., Esperanca, S. and LeCheminant, A.N., 1994. Geochemistry and origin of the Proterozoic ultrapotassic rocks of the Churchill Province, Canada. Contr. Mineral. Petrol., 51:251-276. Peterson, T.D., van Breemen, O., Sandeman, H.A. and Rainbird, R.H., 2000. Proterozoic (1.85-1.75 Ga) granitoid plutonism and tectonics of the Western Churchill Province. GAC-MAC Annual Meeting, Program with Abstracts (CD-ROM). Peterson, T.D., van Breemen, O., Sandeman, H.A. and Cousens, B.L., 2003. Proterozoic (1.9-1.7 Ga) igneous suites of the western Churchill Province: granitoid and ultrapotassic magmatism in an Archean hinterland. Precambrian Res. (in press). Petford, N. and Gallagher, K., 2001. Partial melting of mafic (Amphibolitic) lower crust by periodic influx of basaltic magma. Earth Planet. Sci. Lett., 193: 483-499. Petford, N., Cruden, A.R., McCaffrey, K.J.W. and Vigneresse, J.L., 2000. Granite magma formation, transport and emplacement in the Earth's crust. Nature, 408: 669-673. Pettijohn, F.J., 1934. Conglomerate of Abram Lake, Ontario, and its extensions. Geol. Soc. Amer. Bull., 45:479-505 (incl. plate 51 map, scale 1 : 126720). Pettijohn, F.J., 1984. Memoirs of an Unrepentant Field Geologist: A Candid Profile of Some Geologists and Their Science, 1921-1981. Univ. Chicago Press, 288 pp. Pettijohn, F.J. and Bastron, A., 1959. Chemical composition of argillites of the Cobalt Series (Precambrian) and the problem of soda-rich sediments. Geol. Soc. Amer. Bull., 70: 593-600. Pettijohn, EJ. and Potter, P.E., 1964. Atlas and Glossary of Primary Sedimentary Structures. Springer-Verlag, New York, 370 pp. P6w6, T.L., 1959. Sand-wedge polygons (tesselations) in the McMurdo Sound region, Antarctica-a progress report. Amer. J. Sci., 257: 545-552. Pflug, H.D., 1978. Frtiheste bisher bekannte Lebewesen; Isuasphaera isua n. gen. N. spec. Aus der Isua-Serie von Groenland (ca. 3800 Mio. J.). Oberhessische naturwissenschaftlische Zeitschrift, 44: 131-145. Pflug, H.D., 2001. Earliest organic evolution. Essay to the memory of Bartholomew Nagy. Precambrian Res., 106:79-91.

References

875

Pflug, H.D. and Jaeschke-Boyer, H., 1979. Combined structural and chemical analysis of 3800-Myrold microfossils. Nature, 280: 483-486. Pfliiger, E, 1999. Matground Structures and Redox Facies. Palaios, 14: 25-39. Pfltiger, E and Gresse, P.G., 1996. Microbial sand-chips, a non-actualistic sedimentary structure. Sediment. Geol., 102: 263-274. Pfltiger, E and Seilacher, A., 1991. Flash flood conglomerates, In: G. Einsele, W. Richen and A. Seilacher (Eds.), Cycles and Events in Stratigraphy. Springer-Verlag, Berlin, pp. 383-391. Pharaoh, T.C. and Brewer, T.S., 1990. Spatial and temporal diversity of early Proterozoic volcanic sequences---comparisons between the Baltic and Laurentian schields. Precambrian Res., 47: 169-189. Phillips, G.N., Law, J.D.M. and Myers, R.E., 2001. Is the redox state of the Archean atmosphere constrained? Soc. Econ. Geol. Newsletter, 47: 1; 9-18. Phillips, R.J. and Hansen, V.L., 1994. Tectonic and magmatic evolution of Venus. Annu. Rev. Earth Planet. Sci., 22: 597-654. Phillips, R.J. and Hansen, V.L., 1998. Geological evolution of Venus: rises, plains, plumes and plateaus. Science, 279: 1492-1497. Philpotts, A.R., 1977. Archean variolites---quenched immiscible liquids: discussion. Can. J. Earth Sci., 14: 139-144. Phinney, W.C., Morrison, D.A. and Maczuga, D.E., 1988. Anorthosites and Related Megacrystic Units in the Evolution of Archean Crust. J. Petrol., 29: 1283-1323. Picard, C. and Piboule, M., 1986a. Prtrologie des roches volcaniques du sillon de roches vertes archrennes de Matagami-Chibougamau ?a l'ouest de Chapais (Abitibi est, Qurbec). 2. Le groupe hautement potassique d'Oprmisca. Can. J. Earth Sci., 23:1169-1189. Picard, C. and Piboule, M., 1986b. Prtrologie des roches volcaniques du sillon de roches vertes archrennes de Matagami-Chibougamau a l'ouest de Chapais (Abitibi est, Qurbec). 1. Le groupe basal de Roy. Can. J. Earth Sci., 23:561-578. Picard, M.D. and High, L.R. Jr., 1972. Criteria for recognizing lacustrine rocks. In: J.K. Rigby and W.K. Hamblin (Eds.), Recognition of Ancient Sedimentary Environments. SEPM, Tulsa, Oklahoma, Spec. Publ., 16, pp. 108-145. Piche, M., Guha, J. and Daigneault, R., 1993. Stratigraphic and structural aspects of the volcanic rocks of the Matagami mining camp, Quebec: implications for the Norita ore deposit. Econ. Geol., 88: 1542-1558. Pickard, A.L., 2002. SHRIMP U-Pb zircon ages of tuffaceous mudrocks in the Brockman Iron Formation of the Hamersley Range, Western Australia. Austral. J. Earth Sci., 49:491-507. Pickett, C., 2002. A sedimentary facies analysis of the > 2.8 Ga Beniah Lake and Bell Formations, Slave Province, Northwest Territories. M.Sc. Thesis (unpubl.), Univ. Qurbec a Chicoutimi, Qurbec, 136 p (with 1 map). Pickett, C. and Mueller, W.U., 2000. Lithofacies characteristics and petrology of a stable Archean shelf sequence: slave province, NWT. In: E Cook and P. Erdmer (Compilers), Slave-Northern Cordillera Lithospheric Evolution (SNORCLE) Transect and Cordilleran Tectonics Workshop meeting (Feb. 25-27), Univ. Calgary, Lithoprobe Report No. 72, pp. 11-24. Pidgeon, R.T., Nemchin, A.A. and Williams, I.S., 2001. Internal structures of > 3.9 Ga detrital zircons from the Jack Hills, Western Australia. In: K.K. Cassidy, J.M. Dunphy and M.J. van Kranendonk (Eds.), Extended Abstracts, 4th International Archaean Symposium 2001. Geoscience Australia, Record 2001/37, pp. 75-77. Pierazzo, E., Vickery, A.M. and Melosh, H.J., 1997. A reevaluation of impact melt production. Icarus, 127: 408-423.

876

References

Pike, G., 2001. The facies architecture of two contrasting volcano-sedimentary basin successions from the Archaean Whim Creek belt, North Pilbara terrain, Western Australia: the intra-continental arc-related Whim Creek Group and plume-related, continental rift-hosted Bookingarra Group. Ph.D. Thesis (unpubl.), Monash Univ., Australia, 465 pp. Pike, G. and Cas, R., 2002. Stratigraphic evolution of Archaean volcanic rock-dominated rift basins from the Whim Creek Belt, west Pilbara Craton, Western Australia. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. Blackwell Science, Oxford, pp. 213-234. Pike, G., Cas, R. and Smithies, R.H., 2002. Geologic constraints on base metal mineralization of the Whim Creek greenstone belt, Pilbara craton, Western. Australia. Econ. Geol., 97: 827-845. Pilkington, M. and Grieve, R.A.E, 1992. The geophysical signature of terrestrial impact structures. Rev. Geophys., 30:161-181. Pilote, P., Moorhead, J. and Mueller, W.U., 1998. D6veloppement d'un arc volcanique, la r6gion de Val d'Or ceinture de l'Abitibi--volcanologie physique et 6volution m6tallog6nique. Guide d'excursion A2, AGC/GAC, Qu6bec, Canada, 104 pp. Piper, J.D.A., 1983. Proterozoic palaeomagnetism and single continent plate tectonics. Geophys. J. Roy. Astr. Soc., 74: 163-197. Piper, J.D.A., 2000. The Neoproterozoic supercontinent: Rodinia or Palaeopangaea? Earth Planet. Sci. Lett., 176: 131-146. Pirajno, E and Adamides, N.G., 2000. Geology and mineralisation of the Palaeoproterozoic Yerrida basin, Western Australia. Geol. Surv. Western Australia, Report 60, 43 pp. Pirajno, E, Occhipinti, S.A. and Swager, C.P., 1998. Geology and tectonic evolution of the Palaeoproterozoic Bryah, Padbury and Yerrida Basins (formerly Glengarry Basin), Western Australia: implications for the history of the south-central Capricorn Orogen. Precambrian Res., 90:119-140. Pisarevsky, S.A., Li Z.-X., Grey, K. and Stevens, M.K., 2001. A palaeomagnetic study of Empress 1A, a stratigraphic drillhole in the Officer basin: evidence for a low-latitude position of Australia in the Neoproterozoic. Precambrian Res., 110: 93-108. Playford, P.E., 1980. Devonian "Great Barrier Reef" of the Canning Basin, Western Australia. Amer. Assoc. Petrol. Geol. Bull., 64: 814-840. Plint, A.G., 1988. Sharp-based shoreface sequences and "offshore bars" in the Cardium Formation of Alberta; their relationship to relative changes in sea level. In: C.K. Wilgus, B.S. Hastings, C.G. St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. van Wagoner (Eds.), Sea Level ChangesmAn Integrated Approach. Soc. Econ. Paleontol. Mineral. Spec. Publ., 42, pp. 357-370. Plint, A.G. and Nummedal, D., 2000. The falling stage systems tract: recognition and importance in sequence stratigraphic analysis. In: D. Hunt and R.L. Gawthorpe (Eds.), Sedimentary Response to Forced Regression. Geol. Soc. London Spec. Publ., 172, pp. 1-17. Plumb, K.A., Derrick, G.M. and Wilson, I.H., 1980. Precambrian geology of the McArthur RiverMount Isa region, northern Australia. In: R.A. Henderson and P.J. Stephenson (Eds.), The Geology and Geophysics of Northeastern Australia. Geol. Soc. Australia (Queensland Division), pp. 71-88. Plumb, K.A., Ahmad, M. and Wygralak, A.S., 1990. Mid-Proterozoic basins of the northern Australian craton--regional geology and mineralisation. In: EE. Hughes (Ed.), Geology of the Mineral Deposits of Australia and Papua New Guinea. Australasian Institute of Mining and Metallurgy, Melbourne, pp. 881-902. Podosek, EA. and Cassen, P., 1994. Theoretical, observational and isotopic estimates of the lifetime of the solar nebula. Meteoritics, 29: 6-25.

References

877

Podosek, EA. and Nichols Jr., R.H., 1997. Short-lived Radionuclides in the solar nebula. In: T.J. Bernatowicz and E. Zinner (Eds.), Astrophysical Implications of the Laboratory Study of Presolar Materials. Proceedings of AlP conference, St. Louis, Missouri, Amer. Inst. Physics, Woodbury, New York, pp. 617-648. Podosek, EA. and Ozima, M., 2000. The Xenon age of the Earth. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 63-72. Polat, A. and Kerrich, R., 1999. Formation of an Archean tectonic melange in the Schreiber-Hemlo greenstone belt, Superior Province, Canada; implications for Archean subduction-accretion process. Tectonics, 18: 733-755. Polat, A. and Kerrich, R., 2000. Archean greenstone belt magmatism and the continental growthmantle evolution connection; constraints from Th-U-Nb-LREE systematics of the 2.7 Ga Wawa Subprovince, Superior Province, Canada. Earth Planet. Sci. Lett., 175:41-54. Polat, A. and Kerrich, R., 2001a. Geodynamic processes, continental growth, mantle evolution in late Archean greenstone belts of the southern Superior Province, Canada. Precambrian Res., 112: 5-25. Polat, A. and Kerrich, R., 200lb. Magnesian andesites, Nb-enriched basalt-andesites, and adakites from late Archean 2.7 Ga Wawa greenstone belts, Superior Province, Canada: implications for late Archean subduction zone petrogenetic processes. Contrib. Mineral. Petrol., 141: 36-52. Polat, A., Kerrich, R. and Wyman, D.A., 1998. The late Archean Schreiber-Hemlo and White RiverDayohessarah greenstone belts, Superior Province: collages of oceanic plateaus, oceanic arcs, and subduction-accretion complexes. Tectonophysics, 289: 295-326. Polat, A., Hofmann, A.W. and Rosing, M.T., 2002. Boninite-like volcanic rocks in the 3.7-3.8 Ga Isua greenstone belt, West Greenland: geochemical evidence for intra-oceanic subduction zone processes in the early Earth. Chem. Geol., 184: 231-254. Polat, A., Hofmann, A.W., Munker, C., Regelous, M. and Appel, P.W.U., 2003. Contrasting geochemical patterns in the 3.7-3.8 Ga pillow basalt cores and rims, Isua greenstone belt, Southwest Greenland: implications for post magmatic alteration processes. Geochim. Cosmochim. Acta, 67: 441-457. Polet, J. and Anderson, D.L., 1995. Depth extent of cratons as inferred from tomographic studies. Geology, 23: 205-208. Pollack, H.N., 1997. Thermal characteristics of the Archaean. In M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 223-293. Porada, H. and Behr, H.J., 1988. Setting and sedimentary facies of late Proterozoic alkali lake (playa) deposits in the southern Damara Belt of Namibia. Sediment. Geol., 58:171-194. Porcelli, D., Woolum, D. and Cassen, P., 2001. Deep Earth rare gases: initial inventories, capture from the solar nebula, and losses during Moon formation. Earth Planet. Sci. Lett., 193: 237-251. Porter, S.M. and Knoll, A.H., 2000. Testate amoebae in the Neoproterozoic Era: evidence from vaseshaped microfossils in the Chuar Group, Grand Canyon. Paleobiol., 26: 360-385. Posamentier, H.W., 2001. Lowstand alluvial bypass systems: incised vs. unincised. Amer. Assoc. Petrol. Geol. Bull., 85(10): 1771-1793. Posamentier, H.W., 2002. Ancient shelf ridges--A potentially significant component of the transgressive systems tract: case study from offshore northwest Java. Amer. Assoc. Petrol. Geol. Bull., 86(1): 75-106. Posamentier, H.W. and Allen, G.P., 1993. Variability of the sequence stratigraphic model: effects of local basin factors. Sediment. Geol., 86: 91-109.

878

References

Posamentier, H.W. and Allen, G.P., 1999. Siliciclastic Sequence Stratigraphy: Concepts and Applications. Concepts in Sedimentology and Paleontology No. 7. Soc. Econ. Paleont. Mineral., Tulsa, Oklahoma, 210 pp. Posamentier, H.W. and Morris, W.R., 2000. Aspects of the stratal architecture of forced regressive deposits. In: D. Hunt and R.L. Gawthorpe (Eds.), Sedimentary Responses to Forced Regressions. Geol. Soc. London Spec. Publ., 172, pp. 19-46. Posamentier, H.W. and Vail, P.R., 1988. Eustatic controls on clastic deposition ll--sequence and systems tract models. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross, and J.C. Van Wagoner (Eds.), Sea Level Changes--An Integrated Approach. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 42, pp. 125-154. Posamentier, H.W. and Walker, R.G., 2002. Turbidite facies models; integrating subsurface and outcrop. Abstract, Canadian Society of Petroleum Geologists Diamond Jubilee Convention, June 3-7, 2002, p. 283. Posamentier, H.W., Jervey, M.T. and Vail, P.R., 1988. Eustatic controls on clastic deposition I - conceptual framework. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. van Wagoner (Eds.), Sea Level Changes--An Integrated Approach. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 42, pp. 110-124. Posamentier, H.W., Allen, G.P., James, D.P. and Tesson, M., 1992. Forced regressions in a sequence stratigraphic framework: concepts, examples, and exploration significance. Amer. Assoc. Petrol. Geol. Bull., 76: 1687-1709. Postma, G., Nemec, W. and Kleinspehn, K.L., 1988. Large floating clasts in turbidites: a mechanism for their emplacement. Sediment. Geol., 58: 47-61. Potter, P.E., 1986. South America and a few grains of sand: Part 1--beach sands. J. Geol., 94: 301-319. Poulsen, C.J., Pierrehumbert, R.T. and Jacob, R.L., 2001. Impact of ocean dynamics on the simulation of the Neoproterozoic "snowball Earth". Geophys. Res. Lett., 28: 1575-1578. Poulsen, K.H., Borradaile, G.J. and Kehlenbeck, M.M., 1980. An inverted Archean succession at Rainy Lake, Ontario. Can. J. Earth Sci., 17: 1358-1369. Powell, C.M., Preiss, W.V., Gatehouse, C.G., Krapez, B. and Li, Z.-X., 1994. South Australian record of a Rodinian epicontinental basin and its mid-Neoproterozoic breakup (~ 700 Ma) to form the Paleo-Pacific Ocean. Tectonophysics, 237:113-140. Powell, C.M., Oliver, N.H.S., Li, Z.-X., Martin, D.M. and Ronaszeki, J., 1999. Synorogenic hydrothermal origin for giant Hamersley iron oxide ore bodies. Geology, 27: 175-178. Powell, C.M., Jones, D.L., Pisarevsky, S. and Wingate, M.T.D., 2001. Palaeomagnetic constraints on the position of the Kalahari craton in Rodinia. Precambrian Res., 110: 33-46. Prasad, N. and Roscoe, S.M., 1996. Evidence of anoxic to oxic atmospheric change during 2.45-2.22 Ga from lower and upper sub-Huronian paleosols, Canada. Catena, 27: 105-121. Pratt, B.R., 2001. Oceanography, bathymetry and syndepositional tectonics of a Precambrian intracratonic basin: integrating sediments, storms, earthquakes and tsunamis in the Belt Supergroup (Helena Formation, ca. 1.45 Ga), western North America. Sediment. Geol., 141/142: 371-394. Pratt, B.R. and James, N.P., 1986. The St. George Group (Lower Ordovician) of western Newfoundland: Tidal flat island model for carbonate sedimentation in shallow epeiric seas. Sedimentology, 33: 313-343. Prave, A.R., 1999. The Neoproterozoic Dalradian Supergroup of Scotland: an alternative hypothesis. Geol. Mag., 136:609-617. Preiss, W.V., 1976. Basic field and laboratory methods for the study of stromatolites. In: M.R. Walter (Ed.), Stromatolites. Developments in Sedimentology, Vol. 20, Elsevier, Amsterdam, pp. 5-13.

References

879

Preiss, W.V. (Compiler), 1987. The Adelaide Geosyncline. Geol. Surv. S. Aust. Bull., 53, 438 pp. Priscu, J.C., Adams, E.E., Lyons, W.B., Voytek, M.A., Mogk, D.W., Brown, R.L., McKay, C.P., Takacs, C.D., Welch, K.A., Wolf, C.F., Ki_rshtein, J.D. and Avci, R., 1999. Geomicrobiology of subglacial ice above Lake Vostok, Antarctica. Science, 286:2141-2144. Prokoph, A. and Barthelmes, F., 1996. Detection of nonstationarities in geological time series: wavelet transform of chaotic and cyclic sequences. Computer & Geosciences, 22:1097-1108. Prokoph, A., Ernst, R.E. and Buchan, K.L., 2003. Time-series analysis of Large Igneous Provinces: 3500 Ma to Present. J. Geology (in press). Puchtel, I.S. and Humayun, M., 2001. Platinum group element fractionation in a komatiitic basalt lava lake. Geochim. Cosmochim. Acta, 65: 2979-2993. Puchtel, I.S., Hofmann, A.W., Mezger, K., Shchipansky, A.A., Kulikov, V.S. and Kulikova, V.V., 1996. Petrology of a 2.41 Ga remarkably fresh komatiitic basalt lava lake in Lion Hills, central Vetreny Belt, Baltic Shield. Contrib. Miner. Petrol., 124: 273-290. Puchtel, I.S., Hofmann, A.W., Mezger, K., Jochum, K.P., Shchipansky, A.A. and Samsonov, A.V., 1998a. Oceanic plateau model for continental crustal growth in the Archaean: a case study from the Kostomuksha greenstone belt, NW Baltic Shield. Earth Planet. Sci. Lett., 155: 57-74. Puchtel, I.S., Arndt, N.T., Hofmann, A.W., Haase, K.M., Kr6ner, A., Kulikov, V.S., Kulikova, V.V., Garbe-Sch6neberg, C.-D. and Nemchin, A.A., 1998b. Petrology of mafic lavas within Onega plateau, central Karelia: evidence for 2.0 Ga plume-related continental crustal growth in the Baltic Shield. Contrib. Miner. Petrol., 130: 134-153. Pyke, D.R., Naldrett, A.J. and Eckstrand, O.R., 1973. Archean ultramafic flows in Munro Township, Ontario. Geol. Soc. Amer. Bull., 84: 955-978. Pysklywec, R,N. and Mitrovica, J.X., 1998. Mantle flow mechanisms for the large scale subsidence of continental interiors. Geology, 26: 687-690. Qian, Y.-Z. and Wasserburg, G.J., 2000. Stellar abundances in the early galaxy and two r-process components. Phys. Rep., 333/334: 77-108. Qiu, Y. and Groves, D.I., 1999. Late Archaean collision and delamination in the southwest Yilgarn craton: the driving force for Archaean orogenic lode gold mineralization? Econ. Geol., 94: 115-122. 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. Quitte, G., Birck, J.L. and All~gre, C.J., 2000. 182Hf-182W systematics in eucrites: the puzzle of iron segregation in the early solar system. Earth Planet. Sci. Lett., 184: 83-94. Raaben, M.E., 1969. Columnar stromatolites and late Precambrian stratigraphy. Amer. J. Sci., 267: 1-18. Radhakrishna, B.P. (Ed.), 1986. Banded Iron-formation of India. Geol. Soc. India J. (Spec. Issue), 28, 270 pp. Rainbird, R.H., 1992. Anatomy of a large-scale braid-plain quartzarenite from the Neoproterozoic Shaler Group, Victoria Island, Northwest Territories, Canada. Can. J. Earth Sci., 29: 2537-2550. Rainbird, R.H. and Hadlari, T., 2000. Revised stratigraphy and sedimentology of the Paleoproterozoic Dubawnt Supergroup at the northern margin of Baker Lake Basin, Nunavut. Current Research Paper 2000-C8, Geol. Surv. Can., pp. 1-9. Rainbird, R.H. and Peterson, T.D., 1990. Physical volcanology and sedimentology of lower Dubawnt Group strata, Dubawnt Lake, District of Keewatin, N.W.T. Current Research, Paper, 11990-1C, Geol. Surv. Canada, pp. 207-217. Rainbird, R.H., Nesbitt, H.W. and Donaldson, J.A., 1990. Formation and diagenesis of a subHuronian saprolith: comparison with a modem weathering profile. J. Geol., 98:801-822.

880

References

Rainbird, R.H., Hadlari, T. and Donaldson, J.A., 1999. Stratigraphy and paleogeography of the Paleoproterozoic Baker Lake Group in the eastern Baker Lake Basin, Northwest Territories (Nunavut). Current Res. Paper 1999-C., Geol. Surv. Can., pp. 43-54. Rainbird, R.H., Hadlari, T. and Donaldson, J.A., 2001a. Evolution of the Paleoproterozoic intracratonic Baker Lake and Thelon basins, Western Churchill Province, Canada. In: K.E Cassidy, J.M. Dunphy and M.J. van Kranendonk (Eds.), 4th International Archean Symposium. AGSOGeoscience Australia Record 2001/37, Perth, Australia, pp. 259-260. Rainbird, R.H., Hadlari, T. and Davis, W.J., 200lb. Sequence stratigraphy and tectonic setting of the 1.78-1.74 Ga Wharton Group, Baker Lake and Thelon basins, Western Churchill Province, Nunavut. Abstract, Geol. Assoc. Can., Mineral. Assoc. Can., Joint Annual Meeting, Vol. 26., 123 pp. Rainbird, R.H., Davis, W.J. and Stem, R.A, 2002a. Intergrated geochronology of the late Paleoprotereozoic Baker Lake Group (Dubawnt Supergroup), Baker Lake basin. Abstract, Geol. Assoc. Canada, Mineral. Assoc. Canada, Joint Annual Meeting, Vol. 27, p. 95. Rainbird, R.H., Stem, R.H., and Jefferson, C.W., 2002b. Summary of detrital zircon geochronology of the Athabasca Group, Northern Saskatchewan and Alberta. Saskatchewan Geol. Surv., Summary of Investigations 2002, Vol. 2, Sask. Industry Resources, Misc. Rep. 2002-4.2 (CD-ROM). Rainbird, R.H., Hadlari, T., Aspler, L.B., Donaldson, J.A., LeCheminant, A.N. and Peterson, T.D. (submitted). Evolution of the Paleoproterozoic Baker Lake and Thelon basins, western Churchill Province, Nunavut, Canada. Precambrian Res. Raiswell, R. and Canfield, D.E., 1998. Sources of iron for pyrite formation in marine sediments. Amer. J. Sci., 298: 219-245. Raith, J.G. and Meisel, T., 2001. Metabasites along the amphibolite-granulite facies transition in the Okiep Copper District, South Africa. S. Afr. J. Geol., 104: 77-100. Raith, J.G., Cornell, D.H., Frimmel, H.E. and de Beer, C.H., 2003. New insights into the geology of the Namaqua Tectonic Province, South Africa, from ion probe dating of detrital and metamorphic zircon. J. Geol. (in press). Raith, M.M., Srikantappa, C., Ashamanjari, K.G. and Spiering, B., 1990. The granulite terrane of the Nilgiri Hills (Southern India): characterization of high-grade metamorphism. In: D.Vielzeuf and P.Vidal (Eds.), Granulites and Crustal Evolution. Kluwer Academic Publishers, Dordrecht, pp. 339-365. Raith, M.M., Srikantappa, C., Buhl, D. and Koeler, H., 1999. The Nilgiri enderbites, South India: nature and age constraints on protolith formation, high-grade metamorphism and cooling history. Precambrian Res., 98: 129-150. Ramaekers, P., 1979. Stratigraphy of the Athabasca Basin. Saskatchewan Geol. Surv., Summary of Investigations 1979, Saskatchwan Mineral Resources, Misc. Rep. 79/10, pp. 154-160. Ramaekers, P., 1980. Stratigraphy and tectonic history of the Athabasca Group (Helikian) of Northern Saskatchewan. Saskatchewan Geol. Surv., Summary of Investigations 1980, Saskatchwan Mineral Resources, Misc. Rep. 80/4, pp. 99-106. Ramaekers, P., 1981. Hudsonian and Helikian basins of the Athabasca Region, Northern Saskatchewan. In: EH.A. Campbell (Ed.), Proterozoic Basins of Canada. Geol. Surv. Can. Paper 81/10, pp. 219-234. Ramaekers, P., 1990. The geology of the Athabasca Group (Helikian) in Northern Saskatchewan. Rep. 195, Saskatchewan Energy and Mines, pp. 1-49. Ramaekers, P., 2003. Phases 1 to 4 Ex-tech IV Study of the Early Proterozoic Athabasca Group, Northeastern Alberta. Energy Utilities Board/Alberta Geol. Surv., Spec. Rep. 61, pp. 1-36. Ramberg, H., 1967. Gravity, Deformation and the Earth's Crust. Academic Press, London, 214 pp.

References

881

Rampino, M.R. and Stothers, R.B., 1999. Periodicity in extinctions and impacts; where do we stand after 15 years? In: Anonymous (Ed.), Geological Society of America, 1999 annual meeting., Vol. 31; 7, Abstracts with Programs. Geol. Soc. Amer., Boulder, Colorado, pp. 398. Ramsay, J.G. and Huber, M.I., 1987. The Techniques of Modern Structural Geology, Vol. 2: Folds and Fractures. Academic Press, London, 700 pp. Ramsay, W.R.H., Crawford, A.J. and Foden J.D., 1984. Field setting, mineralogy, chemistry, and genesis of arc picrites, New Georgia, Solomon Islands. Contrib. Mineral. Petrol., 88: 386-402. Rapp, R.P., 1997. Heterogeneous source regions for Archaean granitoids: experimental and geochemical evidence. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 267-279. Rasmussen, B., 2000. Filamentous microfossils in a 3,235-million-year-old volcanogenic massive sulphide deposit. Nature, 405: 676-679. Rasmussen, B. and Buick, R., 1999. Redox state of the Archean atmosphere: evidence from detrital heavy minerals in ca. 3250-2750 Ma sandstones from the Pilbara Craton, Australia. Geology, 27: 115-118. Ratcliff, D.W. and Weber, D.J., 1997. Geophysical imaging of subsalt geology. The Leading Edge, Feb. 1997:115-142. Rautenbach, C.J. de W., 2001. A hypothetical approach to determining the effect of palaeorotational rates on Earth's Neoproterozoic palaeoclimate. J. Afr. Earth Sci., 33: 463-473. Ray, J.S., Martin, M.W., Veizer, J. and Bowring, S.A., 2002. U-Pb zircon dating and Sr-isotope systematics of the Vindhyan Supergroup, India. Geology, 30:131-134. Reading, H.G. (Ed.), 1996. Sedimentary Environments: Processes, Facies and Stratigraphy, 3rd edn. Blackwell, Oxford, 688 pp. Reeves, C.C., 1968. Introduction to Paleolimnology. Developments in Sedimentology, 11. Elsevier, Amsterdam, 228 pp. Reid, D.L., 1997. Sm-Nd age and REE geochemistry of Proterozoic arc-related igneous rocks in the Richtersveld Subprovince, Namaqua Mobile Belt, Southern Africa. J. Afr. Earth Sci., 24: 621-633. Reid, D.L., Welke, H.J., Erlank, A.J. and Betton, P.J., 1987. Composition, age and tectonic setting of amphibolites in the central Bushmanland Group, western Namaqua Province, South Africa. Precambrian Res., 36: 99-12. Reid, D.L., Smith, C.B., Watkeys, M.K., Welke, H.J. and Petton, P.J., 1997. Whole rock radiometric age patterns in the Aggeneys-Gamsberg ore district, central Bushmanland, South Africa. S. Afr. J. Geol., 100:11-22. Reilly, B.A., 1993. Revision bedrock geological mapping of the northwest Amisk Lake area (parts of NTS 63L-9 and-16). Summary of Investigations 1993, Saskatchewan Energy and Mines, Sask. Geol. Surv., Misc. Rep. 93-4, pp. 12-20. Reimer, T.O., 1983. Accretionary lapilli in volcanic ash falls: physical factors governing their formation. In: T. Peryt (Ed.), Coated Grains. Springer-Verlag, Berlin, pp. 56-68. Reimold, W.U. and Colliston, W.P., 1994. Pseudotachylites of the Vredefort Dome and the surrounding Witwatersrand Basin, South Africa. Geol. Soc. Amer. Spec. Paper, 293, pp. 177-196. Reimold, W.U., Koeberl, C. Johnson, S. and McDonald, I., 2000. Early Archaean spherule beds in the Barberton Mountain Land, South Africa: impact or terrestrial origin? In: I. Gilmour and C. Koeberl (Eds.), Impacts and the Early Earth. Lecture Notes in Earth Sci., 92. Springer-Verlag, Berlin, pp. 117-180. Reineck, H.-E., 1979. Rezente und fossile Algenmatten und Wurzelhorizonte. Natur und Museum, 109: 290-296.

882

References

Reineck, H.E. and Singh, I.B., 1980. Depositional Sedimentary Environments, 2nd edn. SpringerVerlag, Berlin, 549 pp. Reineck, H.-E., Gerdes, G., Claes, M., Dunajtschik, K., Riege, H. and Krumbein, W.E., 1990. Microbial modification of sedimentary surface structures. In: D. Heling, P. Rothe, U. F6rstner and P. Stoffers (Eds.), Sediments and Environmental Geochemistry. Springer-Verlag, Berlin, pp. 254-276. Reis, L.A., Martins-Neto, M.A., Gomes, N.S., Endo, I. and Jordt-Evangelista, H., 2002. A bacia de antepafs paleoproteroz6ica Sabar~i, Quadril~itero Ferrifero, Minas Gerais, Brasil. Revista Brasileira de Geoci6ncias, 32: 27-42. Renger, EE., Noce, C.M., Romano, A.W. and Machado, N., 1995. Evoluqao sedimentar do Supergrupo Minas: 500 Ma. de registro geol6gico no Quadril~itero Ferrifero, Minas Gerais, Brasil. Geonomos, 2:1-11. Renner, R., Nisbet, E.G., Cheadle, M.J., Arndt, N.T., Bickle, M.J. and Cameron, W.E., 1994. Komatiite flows from the Reliance, Belingwe belt, Zimbabwe: I. petrography and mineralogy. J. Petrol., 35: 361-400. Retallack, G.J. and German-Heins, J., 1994. Evidence from paleosols for the geological antiquity of rain forest. Science, 265: 499-502. Reynolds, J.H., 1960. Determination of the age of the elements. Phys. Rev. Lett., 4: 8-10. Richards, M.A., Bunge, H.P., Ricard, Y. and Baumgardner, J.R., 1999. Polar wandering in mantle convection models. Geophys. Res. Lett., 26:1777-1780. Richter, F.M., 1985. Models for the Archean thermal regime. Earth Planet. Sci. Lett., 73: 350-360. Riding, R., 1991. Classification of microbial carbonates. In R. Riding (Ed.), Calcareous Algae and Stromatolites. Springer-Verlag, Berlin, pp. 21-51. Riding, R., 2000. Microbial carbonates: the geological record of calcified bacterial-algal mats and biofilms. Sedimentology, 47 (Suppl.1): 179-214. Ridley, J.R., 1992. The thermal causes and effects of voluminous, late Archaean monzogranite plutonism. Geol. Dept. Univ. Extension, Univ. Western Aust, 22: 275-285. Ridley, J.R. and Kramers, J.D., 1990. The evolution and tectonic consequences of a tonalitic magma layer within Archaean continents. Can. J. Earth Sci., 27: 219-228. Righter, K., 2000. A comparison of basaltic volcanism in the Cascades and western Mexico: compositional diversity in continental arcs. Tectonophysics, 318:99-117. Ringwood, A.E., 1966. Chemical evolution of the terrestrial planets. Geochim. Cosmochim. Acta, 30: 40-104. Rivadeneyra, M.A., Delgado, G., Ramos-Cormenzana, A. and Delgado, R., 1998. Biomineralization of carbonates by Halomonas euryhalina in solid and liquid media with different salinities: crystal formation sequence. Res. Microbiol., 149: 277-287. Riverin, G., LaBrie, M., Salmon, B., Cazavant, A., Asselin, R. and Gagnon, M., 1990. The geology of the Ansil deposit, Rouyn-Noranda, Qu6bec. In: M. Rive, P. Verpaelst, Y. Gagnon, J. M. Lulin, G. Riverin and A. Simard (Eds.), La Ceinture polym6tallique du nord-ouest qu6b6cois. Can. Inst. Min. Metall. Spec. Vol., 43, pp. 143-151. Robb, L.J. and Meyer, EM., 1995. The Witwatersrand basin, South Africa: geological framework and mineralization processes. Information Circular 293, Econ. Geol. Res. Unit, Univ. Witwatersrand, South Africa, 37 pp. Robb, L.J., Armstrong, R.A. and Waters, D.J., 1999. The history of granulite-facies metamorphism and crustal growth from single zircon U-Pb geochronology: Namaqualand, South Africa. J. Petrol., 40:1747-1770.

References

883

Robbins, E.J., 1987. Appelella ferrifera, a possible new iron coated microfossil in the Isua iron formation southwestern Greenland. In: P.W. Appel and G. LaBerge (Eds.), Precambrian IronFormations. Theophrastus, Athens, pp. 141-153. Roberts, J.D., 1976. Late Precambrian dolomites, Vendian glaciation and synchroneity of Vendian glaciations. J. Geol., 84: 47-63. Robin, C., Eissen, J.-P. and Monzier, M., 1993. Giant tuff cone and 12-km-wide associated caldera at Ambrym Volcano (Vanuatu, New Hebrides arc). J. Volcanol. Geotherm. Res., 55: 225-238. Roche, O., Druitt, T.H. and Merle, O., 2000. Experimental study of caldera formation. J. Geophys. Res. 105, B 1: 395-416. Roden, G.I., 1963. Sea level variations at Panama. J. Geophys. Res., 68: 5701-5710. RCe, S.-L., 1987. Cross-strata and bedforms of probable dune to upper-stage plane bed origin from a late Precambrian fluvial sandstone, northern Norway. Sedimentology, 34: 89-101. RCe, S.-L. and Hermansen, M., 1993. Processes and products of large, Late Precambrian sandy rivers in northern Norway. In: M. Marzo and C. Puigdef~ibreagas (Eds.), Alluvial Sedimentation. IAS Spec. Publ., 17. Blackwell, Oxford, pp. 151-166. Roedder, E., 1981. Are the 3,800-Myr-old Isua objects microfossils, limonite-stained fluid inclusions, or neither? Nature, 293: 459-462. Roering, C., van Reenen, D.D., Smit, C.A., Barton, J.M., Jr., De Beer, J.H., de Wit, M.J., Stettler, E.H., van Schalkwyk, J.E, Stevens, G. and Pretorius, S., 1992. Tectonic model for the evolution of the Limpopo mobile belt. Precambrian Res., 55: 539-552. Rogers, D.A. and Astin, T.R., 1991. Ephemeral lakes, mud pellet dunes and wind-blown sand and silt: reinterpretations of Devonian lacustrine cycles in north Scotland. In: P. Anadon, L. Cabrera and K. Kelts (Eds.), Lacustrine Facies Analysis. Blackwell, Oxford, IAS Spec. Publ. 13, pp. 199-221. Rogers, J.J.W., 1993. India. India and Ur. J. Geol. Soc., 42:217-222. Rogers, J.J.W., 1996. A history of continents in the past three billion years. J. Geol., 104: 91-107. Rogers, J.J.W. and Santosh, M., 2002. Configuration of Columbia, a Mesoproterozoic supercontinent. Gondwana Res., 5: 5-22. Rogers, N., McNicoll, V., van Staal, C.R. and Tomlinson, K.Y., 2000. Lithogeochemical studies in the Uchi-Confederation greenstone belt, northwestern Ontario: implications for Archean tectonics. Current Research, Part C, Geol. Surv. Canada, pp. 1-11. Rohon, M., Vialette, Y., Clark, T., Roger, G., Ohnenstetter, D. and Vidal, Ph., 1993. Aphebian maficultramafic magmatism in the Labrador Trough (New Quebec): its age and the nature of its mantle source. Can. J. Earth Sci., 30: 1582-1593. Rollinson, H.R., 1993. A terrane interpretation of the Archaean Limpopo belt. Geol. Mag., 130: 755-765. Rollinson, H., 1997. Eclogite xenoliths in west African kimberlites as residues from Archaean granitoid crust formation. Nature, 389: 173-176. Roobol, M.J., Ramsay, C.R., Jackson, N.J. and Darbyshire, D.P.E, 1983. Late Proterozoic lavas of the central Arabian Shield---evolution of an ancient volcanic arc system. J. Geol. Soc. London, 140:185-202. Ropchan, J.R., Luinstra, B., Benn. K., Ayer, J., Berger, B., Dahn, R., Labine, R., Amelin, Y. and Fowler, A.D., 2002. Host rock and structural controls on the nature and timing of gold mineralization at the Holloway mine, Abitibi Subprovince, Ontario. Econ. Geol., 97:291-309. Rosa, M.L.S., Conceicao, H., Martin, H., Conceicao, R.V., Nardi, L.V.S., Oberli, E, Meier, M., Paim, M., Barreto Santos, E., Perreira Santos, C.G., Oliveira, A.E. and Leahy, G.A.S., 1999. Significance of Paleoproterozoic potassic-ultrapotassic syenitic provinces in the Bahia state, Brazil.

884

References

In: B. Barbarin (Ed.), The Origin of Granites and Related RocksnFourth Hutton Symposium. BRGM, Clermont-Ferrand, France, p. 186. Rosa, M.L.S., Conceicao, H., Oberli, E, Meier, M., Martin, H., Macambira, M.J.B., Dos Santos, E.B., Paim, M.M., Leahy, G.A.S. and Leal, L.R.B., 2000. Geochronology (U-Pb/Pb-Pb) and isotopic signature (Rb-Sr/Sm-Nd) of the Paleoproterozoic Guanambi Batholith, southwest Bahia State (NE Brazil). Revista Brasileira de Geosciencias, 30: 62-65. Rose, N.M., Rosing, M.T. and Bridgwater, D., 1996. The origin of metacarbonate rocks in the Archean Isua supracrustal belt, West Greenland. Amer. J. Sci., 296: 1004-1044. Rosen, O.M., Condie, K.C., Natapov, L.M. and Nozhkin, A.D., 1994. Archean and Early Proterozoic evolution of the Siberian craton: a preliminary assessment. In: K.C. Condie (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 411-459. Rosen, O.M., Zhuravlev, D.Z., Sukhanov, M.K., Bibikova, E.V. and Zlobin, V.L., 2000. Early Proterozoic terranes, collision zones and associated anorthosites in the northeast of the Siberian craton: isotope geochemistry and age characteristics. Russian Geology and Geophysics, 41(2): 159-178. Rosen, O.M., Serenko, V.R, Spethius, Z.V., Manakov, A.V. and Zinchuk N.N., 2002. Yakyt kimberlite province: position in the structure of the Siberia craton, compositional features of the upper and lower crust. Russian Geology and Geophysics, 43(1): 3-26. Rosing, M.T., 1999. 13C-depleted carbon microparticles in > 3700-Ma sea-floor sedimentary rocks from West Greenland. Science, 283: 674-676. Rosing, M.T., Rose, N.M., Bridgwater, D. and Thomsen, H.S., 1996. Earliest part of Earth's stratigraphic record: a reappraisal of the > 3.7 Ga Isua (Greenland) supracrustal sequence. Geology, 24: 43-46. Ross, G.M., 2000. Proterozoic stratigraphy of Western Canada: a short review. In: K. Kyser (Ed.), Fluids and Basin Evolution. Short Course Series, Vol. 28. Mineral. Assoc. Can., pp. 211-224. Ross, G.M. and Eaton, D.W., 2002. Proterozoic tectonic accretion and growth of western Laurentia: results from Lithoprobe studies in northern Alberta. Can. J. Earth Sci., 39:313-329. Ross, G.M. and Villeneuve, M.E., 1997. U-Pb geochronology of stranger stones in Neoproterozoic diamictites, Canadian Cordillera: implications for provenance and age of deposition. Geol. Surv. Canada, Current Research, 1997-F: 141-155. Ross, G.M., Parrish, R.S., Villeneuve, M.E. and Bowring, S.A., 1991. Geophysics and geochronology of the crystalline basement of the Alberta basin, Western Canada. Can. J. Earth Sci., 28:512-522. Ross, G.M., Parrish, R.R. and Winston, D., 1992. Provenance and U-Pb geochronology of the Mesoproterozoic Belt Supergroup (northwestern United States): implications for age of deposition and pre-Panthalassa plate reconstructions. Earth Planet. Sci. Lett., 113: 57-76. Ross, G.M., B loch, J.D. and Krouse, H.R., 1995. Neoproterozoic strata of the southern Canadian Cordillera and the isotopic evolution of seawater sulfate. Precambrian Res., 73: 71-99. Ross, G.M., Eaton, D., Boerner, D. and Miles, W., 2000. Tectonic entrapment and its role in the evolution of continental lithosphere: an example from the Precambrian of western Canada. Tectonics, 19:116-134. Rosseto, J.A., Alkmim, EF. and Pereira, M.M., 1987. Litof~icies e modelo deposicional para a Fromaqao Cambotas no maciqo do Caraqa. Simp6sio sobre Sistemas Deposicionais no Precambriano. Ouro Preto, 1987. SBG-Nticelo MG, Roteiro de excurs6es, 25 pp. Rost, R., Dolgov, YU. A., and Vishnevsky, S.A, 1978. Gases in inclusions of impact glasses of the Ries crater (Germany) and the discovery of high pressure polymorphs of carbon. Doklady Akademii Nauk SSSR, 24: 695-698. Rothschild, L.J. and Mancinelli, R.L., 2001. Life in extreme environments. Nature, 409:1092-1101.

References

885

Royden, L.H., 1993. The tectonic expression of slab pull at continental convergent boundaries. Tectonics, 12: 303-325. Rubin, D.M. and Hunter, R.E., 1982. Bedform climbing in theory and nature. Sedimentology, 29: 121-138. Rubin, D.M. and Hunter, R.E., 1987. Bedform alignment in directionally varying flow. Science, 237: 276-278. Rubin, D.M. and Ikeda, H., 1990. Flume experiments on the alignment of transverse, oblique and longitudinal dunes in directly varying flows. Sedimentology, 37: 673-684. Rubincam, D.P., 1995. Has climate changed the Earth's tilt? Paleoceanography, 10: 365-372. Rubincam, D.P., 1999. Mars secular obliquity change due to water ice caps. J. Geophys. Res., 104: 30,765-30,771. Rubincam, D.P., Chao, B.E and Bills, B.G., 1998. The incredible shrinking tropics. Sky and Telescope, 95(6): 36-38. Rudnick, R.L., 1992. XenolithsmSamples of the lower continental crust. In: D.M. Fountain, R. Arculus and R.W. Kay (Eds.), Continental Lower Crust. Elsevier, Amsterdam, pp. 269-316. Rudnick, R.L., Barth, M., Horn, I. and McDonough, W.E, 2000. Rutile-bearing refractory eclogites: missing link between continents and depleted mantle. Science, 287: 278-281. Rumble, D., Bekker, A., Farquhar, J., Holland, H.D., Hu, G., Maruyama, S., Ono, S., Ohmoto, H., Rosing, M., Ueno, Y. and Wing, B.A., 2002. Geologic age distribution of sulfur isotope anomalies: data from a new UV and IR laser microprobe. Abstract, Amer. Geophys. Union, Spring Meeting, Washington, DC, 2002, Vol. 41 A, p. 2. Runcorn, S.K., 1964. Changes in the Earth's moment of inertia. Nature, 204: 823-825. Runcorn, S.K., 1966. Change in the moment of inertia of the Earth as a result of a growing core. In: B.G. Marsden and A.G.W. Cameron (Eds.), The Earth-Moon System. Plenum, New York, pp. 82-92. Runnegar, B., 1992. Evolution of the earliest animals. In: J.W. Schopf (Ed.), Major Events in the History of Life. Jones and Bartlett, Boston, pp. 65-94. Runnegar, B., 2001. Archean sulfates from Western Australia: implications for Earth's early atmosphere and ocean. Abstract, l lth Goldschmidt Conference, Virginia, USA, 2001, abstract no. 3859. Runnegar, B., Coath, C.D., Lyons, J.R. and McKeegan, K.D., 2002. Mass-independent and massdependent sulfur processing throughout the Archean. Geochim. Cosmochim. Acta, 66:A655 (abstract). Russell, M.J. and Hall, A.J., 1997. The emergence of life from iron monosulphide bubbles at a submarine hydrothermal redox and pH front. J. Geol. Soc. London, 154: 377-402. Russell, S.S., Srinivasan, G., Huss, G.R. and Wasserburg, G.J., 1996. Evidence for widespread 26A1 in the solar nebula and constraints for nebula time scales. Science, 273: 757-762. Rutter, E.H. and Brodie, K.H., 1992. Rheology of the lower crust. In: D.M. Fountain, R. Arculus and R.W. Kay (Eds.), Continental Lower Crust. Elsevier, Amsterdam, pp. 201-267. Ruzicka, A., Snyder, G.A. and Taylor, L.A., 2001. Comparative geochemistry of basalts from the Moon, Earth, HED asteroid, and Mars: implications for the origin of the Moon. Geochim. Cosmochim. Acta, 65: 979-997. Ryder, G., 2002. Mass flux in the ancient Earth-Moon system and benign implications for the origin of life on Earth. J. Geophys. Res., 107: 10, 1029/2001JE001583. Ryder, G., 2003. Bombardment of the Hadean Earth: wholesome or deleterious? Astrobiology, 3: 3-6.

886

References

Ryder, G., Koeberl, C. and Mojzsis, S.J., 2000. Heavy bombardment of the Earth at ~ 3.85 Ga: the search for petrographic and geochemical evidence. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 475-492. Rye, R. and Holland, H.D., 1998. Paleosols and the evolution of atmospheric oxygen: a critical review. Amer. J. Sci., 298: 621-672. Rye, R. and Holland, H.D., 2000. Life associated with a 2.76 Ga ephemeral pond? Evidence from Mount Roe #2 paleosol. Geology, 28: 483-486. Rye, R.O., Kuo, P.H. and Holland, H.D., 1995. Atmospheric carbon dioxide concentrations before 2.2 billion years ago. Nature, 378: 603-605. S.A.C.S. (South African Committee for Stratigraphy), 1980. Stratigraphy of South Africa. Part 1 (Comp. L.E. Kent). Lithostratigraphy of the Republic of South Africa, South West Africa/Namibia and the Republics of Boputhatswana, Transkei and Venda. Handbook of the Geological Survey of South Africa, Pretoria, 8, 690 pp. Sagan, C. and Chyba, C., 1997. The early faint Sun paradox: organic shielding of ultraviolet-labile greenhouse gases. Science, 276:1217-1221. Sagan, C. and Mullen, G., 1972. Earth and Mars: evolution of atmospheres and surface temperatures. Science, 177: 52-56. Saha, A.K., Ray, S.L. and Sarkar, S.N., 1988. Early history of the Earth: evidence from the Eastern Indian Shield. In: D. Mukhopadhyay (Ed.), Precambrians of the Eastern Indian Shield. Mem. Geol. Soc. India, 8, pp. 13-38. Sajipal, S., Goswami, J.N., Davis, A.M., Grossman, L. and Lewis, R.S., 1998. A stellar origin for the short-lived nuclides in the early solar system. Nature, 391: 559-561. Sami, T.T. and James, N.P., 1993. Evolution of an early Paleozoic foreland basin carbonate platform, lower Pethei Group, Great Slave Lake, north-west Canada. Sedimentology, 40: 403-430. Sandeman, H., Davis, W.J., Peterson, T.D. and Hanmer, S., 1998. Preliminary petrochemistry of Archean mafic volcanic rocks of the Heninga-to-Tavani corridor of the Rankin-Ennadai greenstone belt, Kivalliq region, N.W.T. 26th Yellowknife Geoscience Forum, INAC Geology Division, Yellowknife, pp. 97-98. Sandeman, H.A., MacLachlan, K. and Relf, C., 1999. Preliminary geochemical and Nd isotopic investigations of Archean volcanic and volcaniclastic rocks of the Yathkyed greenstone belt, Kivalliq region, Northwest Territories. Current Research Paper, 1999-F, Geol. Surv. Can., pp. 43-52. Sandeman, H., Cousens, B.L., Peterson, T., Hemmingway, C., Davis, W.J., Hanmer, S., Tella, S., Ryan, J. and Aspler, L., 2000a. Petrochemistry and Nd isotopic evolution of Proterozoic mafic rocks of the Western Churchill Province, Nunavut, and their implications for the evolution of the lithospheric mantle. Abstract, Geol. Assoc. Can.-Mineral. Assoc. Can. Annu. Meeting, 2000 (CD-ROM). Sandeman, H.A., Lepage, L.D., Ryan, J.J. and Tella, S., 2000b. Field evidence for commingling between ca. 1830 Ma lamprophyric, monzonitic and monzogranitic magmas, MacQuoid-Gibson lakes area, Nunavut. Current Research Paper, 2000-C5, Geol. Surv. Canada, 10 pp. Sandiford, M., 1989a. Secular trends in the thermal evolution of metamorphic terrains. Earth Planet. Sci. Lett., 95: 85-96. Sandiford, M., 1989b. Horizontal structures in granulite terrain: a record of mountain building or mountain collapse? Geology, 17: 449-452. Sanschagrin, Y., 1982. Etude des variations lat6rals et verticales de facies dans des coulees de basalts thol6iitiques du Groupe de Kinoj6vis, Canton d'Aiguebelle, Abitibi. M.Sc. Thesis (unpubl.), Universit6 du Qu6bec a Chicoutimi, 114 pp.

References

887

Santos, J.O.S., Hartmann, L.A., McNaughton, N.J., Easton, R.M., Rea, R.G. and Potter, EE., 2002. Sensitive high resolution ion microprobe (SHRIMP) detrital zircon geochronology provides new evidence for a hidden Neoproterozoic foreland basin to the Grenville Orogen in the eastern Midwest, U.S.A. Can. J. Earth Sci., 39: 1505-1515. Sarkar, S., Chakraborty, S., Banerjee, S. and Bose, P.K., 2002. Facies sequence and occult imprint of sag tectonics in Sirbu Shale. In: W. Altermann and P.L. Corcoran (Ed.), Precambrian Sedimentary Environments: A Modem Approach to Ancient Depositional Systems. Blackwell, Oxford, pp. 369-381. Sarmiento, J., 1992. Biological ocean models. In: K.E. Trenberth (Ed.), Climate System Modeling. Cambridge Univ. Press, Cambridge, U.K., pp. 519-564. Sastry, M.V.A., 1980. Stromatolites and their systematics. Geol. Surv. India, Misc. Publ., 44: 1-10. Saunders, A.D., Tarney, J., Kerr, A.C. and Kent, R.W., 1996. The formation and fate of large oceanic igneous provinces. Lithos, 37: 81-95. Saverikko, M., 1985. The pyroclastic komatiite complex at Sattasvaara in northern Finland. Geol. Soc. Finland Bull., 57: 55-87. Sawyer, E.W. and Barnes, S.J., 1994. Thrusting, magmatic intraplating, and metamorphic core complex development in the Archaean Belleterre-Angliers Greenstone belt, Superior Province, Quebec, Canada. Precambrian Res., 68: 183-200. Sawyer, E.W. and Benn, K., 1993. Structure of the high-grade Opatica belt and adjacent low-grade Abitibi Subprovince, Canada: an Archaean mountain front. J. Struct. Geol., 15: 1443-1458. Saylor, B.Z. and Grotzinger, J.P., 1996. Reconstruction of important Proterozoic-Cambrian boundary exposures through the recognition of thrust deformation in the Nama Group of southern Namibia. Comm. Geol. Surv. Namibia, 11: 1-12. Saylor, B.Z., Kaufman, A.J., Grotzinger, J.P. and Urban, E, 1998. A composite reference section for terminal Proterozoic strata of southern Namibia. J. Sediment. Res., 68: 1223-1235. Schaefer, S.J. and Morton, P., 1991. Two komatiitic pyroclastic units, Superior Province, northwestern Ontario: their geology, petrography, and correlation. Can. J. Earth Sci., 28: 1455-1470. Schaller, M., Steiner, O., Studer, I., Holzer, L., Herwegh, M. and Kramers, J.D., 1999. Exhumation of Limpopo Central Zone granulites and dextral continent-scale transcurrent movement at 2.0 Ga along the Palala Shear Zone, Northern Province, South Africa. Precambrian Res., 96: 263-288. Schandelmeier, H., Abdel Rahman, E.M., Wipfler, E., Ktister, D., Utke, A. and Matheis, G., 1994. Late Proterozoic magmatism in the Nakasib suture, Red Sea Hills, Sudan. J. Geol. Soc. London, 151: 485-497. Schenk, P.E., 1965. Depositional environment of the Gowganda Formation (Precambrian) at the south end of Lake Timagami, Ontario. J. Sediment. Petrol., 35: 309-318. Schenk, P.M., 2002. Impact through the layered icy crusts of Europa, Ganymede, and Callisto. Abstract, Lunar and Planetary Sciences Conference 33, Houston, Texas, 2002, abstract no. 1974. Schermerhorn, L.J.G., 1974. Late Precambrian mixtites: glacial and/or nonglacial? Amer. J. Sci., 274: 673-824. Schidlowski, M.A., 1988. A 3,800-million-year isotopic record of life from carbon in sedimentary rocks. Nature, 333: 313-318. Schidlowski, M., 2001. Carbon isotopes as biogeochemical records of life over 3.8 Ga of Earth history: evolution of a concept. Precambrian Res., 106:117-134. Schidlowski, M. and Aharon, P., 1992. Carbon cycle and carbon isotopic record: geochemical impact on life over 3.8 Ga of Earth history. In: M. Schidlowski (Ed.), Early Organic Evolution: Implications for Mineral and Energy Resources. Springer-Verlag, Heidelberg, pp. 147-175.

888

References

Schidlowski, M., Hayes, J.M. and Kaplan, I.R., 1983. Isotopic inferences of ancient biochemistries: Carbon, sulfur, hydrogen, and nitrogen. In: J.W. Schopf (Ed.), Earth's Earliest Biosphere: Its Origin and Evolution. Princeton Univ. Press, Princeton, pp. 149-186. Schieber, J., 1986. The possible role of benthic microbial mats during the formation of carbonaceous shales in shallow Mid-Proterozoic basins. Sedimentology, 33: 521-536. Schieber, J., 1989. Pyrite mineralization in microbial mats from the Mid-Proterozoic Newland Formation, Belt Supergroup, Montana, U.S.A. Sediment. Geol., 64: 79-90. Schieber, J., 1998. Possible indicators of microbial mat deposits in shales and sandstones: examples from the Mid-Proterozoic Belt Supergroup, Montana, USA. Sediment. Geol., 120: 105-124. Schieber, J., 1999. Microbial mats in terrigenous clastics: the challenge of identification in the rock record. Palaios, 14: 3-12. Schlager, W., 1981. The paradox of drowned reefs and carbonate platforms. Geol. Soc. Amer. Bull., 274: 673-824. Schlager, W., 1991. Depositional bias and environmental change--important factors in sequence stratigraphy. Sediment. Geol., 70: 109-130. Schlager, W., 1992. Sedimentology and sequence stratigraphy of reefs and carbonate platforms. AAPG Course Notes Series (Tulsa, Oklahoma) 34, 71 pp. Schmid, R., 1981. Descriptive nomenclature and classification of pyroclastic deposits and fragments: recommendations of the IUGS subcommission on the systematics of igneous rocks. Geology, 9: 41-43. Schmidt, G., Zhou, L. and Wasson, J.T., 1993. Iridium anomaly associated with the Australasian tektite-producing impact; masses of the impactor and of the Australasian tektites. Geochim. Cosmochim. Acta, 57:4851-4859. Schmidt, P.W., 1999. Proterozoic glaciations: low-latitudes or non-dipole field? Proceedings Nordic Palaeomagnetic Symposium, Aarhus University. Aarhus Geosci., 8:109-112. Schmidt, P.W., 2001. Palaeomagnetic evidence for low-latitude glaciation: the geocentric axial dipole hypothesis for the Proterozoic? Geol. Soc. Aust. Abs., 65: 95-98. Schmidt, P.W. and Williams, G.E., 1995. The Neoproterozoic climatic paradox: equatorial palaeolatitude for Marinoan glaciation near sea level in South Australia. Earth Planet. Sci. Lett., 134: 107-124. Schmidt, P.W. and Williams, G.E., 1999. Paleomagnetism of the Paleoproterozoic hematitic breccia and paleosol at Ville-Marie, Qurbec: further evidence for the low paleolatitude of Huronian glaciation. Earth Planet. Sci. Lett., 172: 273-285. Schmidt, P.W., Williams, G.E. and Embleton, B.J.J., 1991. Low palaeolatitude of Late Proterozoic glaciation: early timing of remanence in haematite of the Elatina Formation, South Australia. Earth Planet. Sci. Lett., 105: 355-367. Schoenberg, R., Kamber, B.S., Collerson, K.D. and Eugster, O., 2002a. New W-isotope evidence for rapid terrestrial accretion and very early core formation. Geochim. Cosmochim. Acta, 66: 3151-3160. Schoenberg, R., Kamber, B.S., Collerson, K.D. and Moorbath, S., 2002b. Tungsten isotope evidence from ~ 3.8-Gyr metamorphosed sediments for early meteorite bombardment of the Earth. Nature, 418: 403-405. Scholle, P.A. and Spearing, D.R. (Eds.), 1982. Sandstone Depositional Environments. AAPG, Tulsa, Oklahoma, Mem., 31,410 pp. Scholle, P.A., Bebout, D.G. and Moore, C.H. (Eds.), 1983. Carbonate Depositional Environments. AAPG, Tulsa, Oklahoma, Mem., 33, 708 pp.

References

889

Schopf, J.W., 1992a. Paleobiology of the Archean. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere, a Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 25-39. Schopf, J.W., 1992b. Proterozoic prokaryotes: affinities, geologic distribution, and evolutionary trends. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere, a Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 195-218. Schopf, J.W., 1992c. Evolution of the Proterozoic biosphere: benchmarks, tempo, and mode. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere, a Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 583-600. Schopf, J.W., 1992d. The oldest fossils and what they mean. In: J.W. Schopf (Ed.), Major Events in the History of Life. Jones and Bartlett, Boston, pp. 29-64. Schopf, J.W., 1993. Microfossils of the early Archean Apex chert: new evidence of the antiquity of life. Science, 260: 640-646. Schopf, J.W., 1999. Cradle of Life: The Discovery of Earth's Earliest Fossils. Princeton Univ. Press, Princeton, New Jersey, 367 pp. Schopf, J.W., 2000. The fossil record: tracing the roots of the cyanobacterial lineage. In: B.A. Whitton and M. Potts (Eds.), The Ecology of Cynaobacteria: Their Diversity in Time and Space. Kluwer, Dordrecht, pp. 13-35. Schopf, J.W. and Klein, C. (Eds.), 1992. The Proterozoic Biosphere--A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, 1348 pp. Schopf, J.W. and Oehler, D.Z., 1976. How old are the eukaryotes? Science, 193: 47-49. Schopf, J.W. and Packer, B.M., 1987. Early Archean (3.3-billion to 3.5-billion-year, old) microfossils from Warrawoona Group, Australia. Science, 237: 70-73. Schopf, J.W. and Sovietov, Yu.K., 1976. Microfossils in Conophyton from the Soviet Union and their bearing on Precambrian Biostratigraphy. Science, 193: 143-146. Schopf, J.W. and Walter, M.R., 1983. Archean microfossils: new evidence of ancient microbes. In: J.W. Schopf (Ed.), Earth's Earliest Biosphere, Its Origin and Evolution. Princeton Univ. Press, Princeton, New Jersey, pp. 214-239. Schopf, J.W., Kudryavtsev, A.B., Agresti, D.G., Wdowiak, T.J. and Czaja, A.D., 2002. Laser-Raman imagery of Earth's earliest fossils. Nature, 416: 73-76. Schopf, J.W., Kudryavtsev, A.B., Agresti, D.G., Czaja, A.D and Wdowiak, T.J. (in preparation). Laser-Raman chemistry and organic metamorphism of Precambrian microscopic fossils. Precambrian Res. Schopf, T.J.M., 1980. Paleoceanography. Harvard Univ. Press, Cambridge, MA, 341 pp. Schrank, A. and Silva, M.G., 1993. Greenstone belts do CrOton do Sao Francisco, Brasil. In: J.M.L. Dominguez and A. Misi (Eds.), O Crfiton do Sao Francisco. SBG/SGM/CNPq, Salvador, pp. 85-118. Schreiber, B.C. and Hsti, K.J., 1980. Evaporites. In: G.D. Hobson (Ed.), Developments in Petroleum Geology, 2. Applied Science Publishers, Barking, Essex, pp. 87-138. Schubert, G., 1988. Physical processes in the growth of the continental crust. In: L.D. Ashwal (Ed.), Workshop on the growth of continental crust. Technical report, Lunar Planetary Inst., Houston, Texas, pp. 131-132. Schumm, S.A., 1968, Speculations concerning paleohydrologic controls of terrestrial sedimentation. Geol. Soc. Amer. Bull., 79: 1573-1588. Schumm, S.A., 1993. River response to baselevel change: implications for sequence stratigraphy. J. Geol., 101: 279-294. Schwab, EL., 1976. Modem and ancient sedimentary basins: comparative accumulation rates. Geology, 4: 723-727.

890

References

Schwerdtner, W.M. and Clark, A.R., 1967. Structural analysis of Mokka Fiord and South Fiord domes, Axel Heiberg Island, Canadian Arctic. Can. J. Earth Sci., 4: 1229-1245. Schwerdtner, W.M. and Van Kranendonk, M.J., 1984. Structure of Stolz diapir--a well-exposed salt dome on Axel Heiberg Island, Canadian Arctic Archipelago. Can. Petrol. Geol. Bull., 32: 237-241. Scoates, R.EJ., 1981. Volcanic rocks of the Fox River Belt, northeastern Manitoba. Manitoba Department of Energy and Mines. Geological Report GR81-1, Mineral Resource Division, 109 pp. Scotese, C.R., Boucout, A.J. and McKerrow, W.S., 1999. Gondwanan palaeogeography and palaeoclimatology. J. Afr. Earth Sci., 28:99-114. Scott, B.P., 1978. The geology of an area east of Thluicho Lake, Saskatchewan. Report No. 167, Saskatchewan Dep. Mineral Resources, pp. 1-51. Scott, C.R., Mueller, W.U. and Pilote, P., 2002. Physical volcanology, stratigraphy and lithogeochemistry of an Archean volcanic arc: evolution from plume-related volcanism to arc rifting of SE Abitibi Greenstone Belt, Val d'Or, Canada. Precambrian Res., 115: 223-260. Scott, D., St.-Onge, M., Lucas, S. and Helmstaedt, H., 1991. Geology and chemistry of the Early Proterozoic Purtuniq ophiolite, Cape Smith belt, northern Quebec, Canada. In: T. Peters, A. Nicolas and R. Coleman (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer Academic Publishers, Amsterdam, pp. 817-849. Scott, D., Helmstaedt, H. and Bickle, M., 1992. Purtuniq ophiolite, Cape Smith belt, northern Quebec, Canada: a reconstructed section of Early Proterozoic oceanic crust. Geology, 20: 173-176. Scott, D.J., 1999. U-Pb geochronology of the eastern Hall Peninsula, Southern Baffin Island, Canada: a northern link between the Archean of West Greenland and the Palaeoproterozoic Torngat Orogen of northern Labrador. Precambrian Res., 93: 5-26. Scridlaite, G. and Motuza, G., 2001. Precambrian domains in Lithuania: evidence of terrane tectonics. Tectonophysics, 339:113-133. Scrivener, C. and Anderson, D.L., 1992. The effect of post Pangaea subduction on global mantle tomography and convection. Geophys. Res. Lett., 19: 1053-1056. Seilacher, A., 1989. Vendozoa: organism construction in the Proterozoic biosphere. Lethaia, 22: 229-239. Seilacher, A., Bose, P.K. and Pfluger, E, 1998. Triploblastic animals more than 1 Billion years ago: trace fossil evidence from India. Science, 282: 80-84. Self, S., Keszthelyi, L. and Thordarson, T., 1998. The importance of p~oehoe. Annu. Rev. Earth Planet. Sci., 26:81-110. Sellers, W.D., 1990. The genesis of energy balance modeling and the cool sun paradox. Palaeogeogr. Palaeoclimatol. Palaeoecol., 82:217-224. Selley, R.C., 1996. Ancient Sedimentary Environments, 4th edn. Chapman and Hall, London, 300 pp. Semikhatov, M.A. and Raaben, M.E., 2000. Proterozoic stromatolite taxonomy and biostratigraphy. In: R.E. Riding and S.M. Awramik (Eds.), Microbial Sediments. Springer-Verlag, Berlin, pp. 295-306. Semikhatov, M.A., Gebelein, C.D., Cloud, P., Awramik, S.M. and Benmore, W.C., 1979. Stromatolite morphogenesis-progress and problems. Can. J. Earth Sci., 16:992-1015. Shackleton, R.M., 1994. Review of late Precambrian sutures, ophiolitic m61anges and tectonics of eastern Egypt and north-east Sudan. Geol. Rundsch., 8: 537-546. Shanley, K.W. and McCabe, P.J., 1994. Perspectives on the sequence stratigraphy of continental strata. Amer. Assoc. Petrol. Geol. Bull., 78: 544-568. Shanley, K.W., McCabe, P.J. and Hettinger, R.D., 1992. Significance of tidal influence in fluvial deposits for interpreting sequence stratigraphy. Sedimentology, 39: 905-930.

References

891

Shanmugam, G., 1988. Origin, recognition and importance of erosional unconformities in sedimentary basins. In: K.L. Kleinspehn and C. Paola (Eds.), New Perspectives in Basin Analysis. Springer-Verlag, New York, pp. 83-108. Sharkov, E.V. and Smolkin, V.E, 1997. The Early Proterozoic Pechenga-Varzuga belt: a case of Precambrian back-arc spreading. Precambrian Res., 82: 133-151. Shaw, A.B., 1964. Time in Stratigraphy. McGraw-Hill, New York, 365 pp. Shaw, H.R., 1994. Craters, Cosmos, and Chronicles. Stanford Univ. Press, California, 688 pp. Shaw, R.D., Wellmam, P., Gunn, P., Whitaker, A.J., Tarlowski, C. and Morse, M., 1996. Guide to using the Australian crustal elements map. Aust. Geol. Surv. Org. Record, 1996/30, 49 pp. Shearer, C.K. and Floss, C., 2000. Evolution of the Moon's mantle and crust as reflected in traceelement microbeam studies of lunar magmatism. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 339-359. Shegelski, R.J., 1980. Archean cratonization, emergence and red bed development, Lake Shebandowan area, Canada. Precambrian Res., 12: 331-347. Shelkov, D.A., Verchovsky, A.B., Milledge, H.J., Kaminsky, EV. and Pillinger, C.T., 1998. Carbon, nitrogen, argon and helium study of impact diamonds from Ebeliakh alluvial deposits and Popigai Crater. Meteoritics Planet. Sci., 33: 985-992. Shen, J.J., Papanastassiou, D.A. and Wasserburg, G.J., 1998. Re-Os systematics in pallasite and mesosiderite metal. Geochim. Cosmochim. Acta, 62: 2715-2723. Shen, Y., Buick, R. and Canfield, D.E., 2001. Isotopic evidence for microbial sulphate reduction in the early Archaean era. Nature, 410:77-81. Shen, Y., Canfield, D.E. and Knoll, A.H., 2002. Middle Proterozoic ocean chemistry: evidence from the McArthur Basin, northern Australia. Amer. J. Sci., 302:81-109. Shen, Y., Knoll, A.H. and Walter, M.R., 2003. Evidence for low sulphate and anoxia in a midProterozoic marine basin. Nature, 423: 632-635. Shields, G. and Veizer, J., 2002. Precambrian marine carbonate isotope database: version 1.1. Geochem. Geophys. Geosyst., 3:10.1029/2001 GC000266. Shiraishi, K., Ellis, D.J., Hiroi, Y., Fanning, C.M., Motoyoshi, Y. and Nakai, Y., 1994. Cambrian orogenic belt in East Antarctica and Sri Lanka: implications for Gondwana assembly. J. Geol., 102: 47-65. Shirey, S.B., Carlson, R.W., Richardson, S.H., Menzies, A., Gurney, J.J., Pearson, D.G., Harris, J.W. and Wiechert, U., 2001. Archaean emplacement of eclogitic components into the lithospheric mantle during formation of the Kaapvaal Craton. Geophys. Res. Lett., 28:2509-2512. Shock, E.L., 1992. Chemical environments in submarine hydrothermal systems. In: N. Holm (Ed.), Marine Hydrothermal Systems and the Origins of Life. Origin of Life and the Evolution of Biosphere, 22 (special issue): 67-107. Shock, E.L., 1997. High-temperature life without photosynthesis as a model for Mars. J. Geophys. Res., 102: 23,687-23,694. Shoemaker, E.M., 1983. Asteroid and comet bombardment of the Earth. Ann. Rev. Earth Planet. Sci., 11 : 461-494. Shoemaker, E.M., 1998. Long-term variations in the impact cratering rate on Earth. In: M.M. Grady, R. Hutchison, G.J.H. McCall and D.A. Rothery (Eds.), Meteorites: Flux with Time and Impact Effects. Geol. Soc. London, Spec. Publ. 140, pp. 7-10. Shoemaker, E.M. and Shoemaker, C.S., 1996. The Proterozoic impact record of Australia. J. Aust. Geol. Geophys., 16: 379-398. Shoemaker, E.M., Wolfe, R.S. and Shoemaker, C.S., 1990. Asteroid and comet flux in the neighborhood of earth. In: V.L. Sharpton and P.D. Ward (Eds.), Global Catastrophes in Earth History:

892

References

An Interdisciplinary Conference on Impacts, Volcanism, and Mass Mortality. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 247, pp. 155-170. Shore, M., 1996. Cooling and Crustallization of Komatiite Flows. Ph.D. Thesis (unpubl.), Univ. Ottawa, Ottawa, 211 pp. Shore, M. and Fowler, A.D., 1996. Oscillatory zoning in minerals: a common phenomenon. Can. Mineral., 34:1111-1126. Shore, M. and Fowler, A.D., 1999. The origin of spinifex texture in komatiites. Nature, 397:691-693. Short, N.M. and Bunch, T.E., 1966. A worldwide inventory of features characteristic of rocks associated with presumed meteorite impact structures. In: B.M. French and M.N. Short (Eds.), Shock Metamorphism of Natural Materials. Proceedings, First Conference on Shock Metamorphism of Natural Materials, Greenbelt, Maryland, Mono. Book Corp., pp. 255-266. Showman, A.P. and Malhotra, R., 1999. The Galilean satellites. Science, 286: 77. Shride, A.E, 1967. Younger Precambrian Geology in southern Arizona. U.S. Geol. Surv. Prof. Paper, 566, 89 pp. Shu, EH., Shang, H. and Lee, T., 1996. Towards an astrophysical theory of chondrites. Science, 271 : 1545-1552. Shukolyukov, A. and Lugmair, G.W., 1998. Isotopic evidence for the Cretaceous-Tertiary impactor and its type. Science, 282: 927-929. Shukolyukov, A. and Lugmair, G.W., 1999. The 53Mn-53Cr isotope systematics of the enstatite chondrites. Abstracts, 30th Lunar and Planetary Science Conference. Lunar and Planetary Sci. Instit., Houston, Texas, abstract no. 1093 (CD-ROM). Shukolyukov, A., Kyte, ET., Lugmair, G.W., Lowe, D.R. and Byerly, G.R., 2000. The oldest impact deposits on earth--first confirmation of an extraterrestrial component. In: I. Gilmour and C. Koeberl (Eds.), Impacts and the Early Earth. Lecture Notes Earth Sci., 92. Springer-Verlag, Berlin, pp. 99-115. Shukolyukov, A., Castillo, P., Simonson, B.M. and Lugmair, G.W., 2002. Chromium in late Archaean spherule layers from Hamersley basin, Western Australia: isotopic evidence for extraterrestrial component. Abstracts, 33rd Lunar and Planetary Science Conference, Houston, Texas, 2002, abstract no. 1369. Siedlecka, A. and Edwards, M.B., 1980. Lithostratigraphy and sedimentation of the Riphean B~snaering Formation, Veranger Peninsula, North Norway. Nor. Geol. Unders., 355: 27-47. Siedlecka, A., Pickering, K.T. and Edwards, M.B., 1989. Upper Proterozoic passive margin deltaic complex, Finnmark, N. Norway. In: M.K.G. Whateley and K.T. Pickering (Eds.), Deltas: Sites and Traps for Fossil Fuels. Geol. Soc. Spec. Publ., 41. Blackwell, Oxford. pp. 205-219. Siever, R., 1992. The silica cycle in the Precambrian. Geochim. Cosmochim. Acta, 56: 3265-3272. Silver, L.T., 1978. Precambrian formations and Precambrian history in Cochise County, southeastern Arizona. New Mexico Geol. Soc. Field Conf. Guidebook 29, pp. 157-163. Simkin, T. and Fiske, R.S., 1983. Krakatau, 1883--the Volcanic Eruption and its Effects. Smithsonian Inst. Press, Washington, 464 pp. Simonson, B.M., 1992. Geological evidence for a strewn field of impact spherules in the early Precambrian Hamersley Basin of Western Australia. Geol. Soc. Amer. Bull., 104: 829-839. Simonson, B.M., 2003. Petrographic criteria for recognizing certain types of impact spherules in well-preserved Precambrian successions. Astrobiology, 3: 49-65. Simonson, B.M. and Carney, K.E., 1999. Roll-up structures: evidence of in situ microbial mats in Late Archean deep shelf environments. Palaios, 14: 13-24. Simonson, B.M. and Harnik, P., 2000. Have distal impact ejecta changed through geologic time? Geology, 28: 975-978.

References

893

Simonson, B.M. and Hassler, S.W., 1996. Was the deposition of large Precambrian iron formations linked to major marine transgressions? J. Geol., 104: 665-676. Simonson, B.M. and Hassler, S.W., 1997. Revised correlations in the Early Precambrian Hamersley Basin based on a horizon of resedimented impact melt. Aust. J. Earth Sci., 44: 37-48. Simonson, B.M., Hassler, S.W. and Schubel, K.A., 1993a. Lithology and proposed revision in stratigraphic nomenclature of the Wittenoom Formation (Dolomite) and overlying formations, Hamersley Group, Western Australia. Report 34, Geol. Surv. Western Aust., pp. 65-79. Simonson, B.M., Schubel, K.A. and Hassler, S.W., 1993b. Carbonate sedimentology of the Early Precambrian Hamersley Group of Western Australia. Precambrian Res., 60: 287-335. Simonson, B.M., Davies, D., Wallace, M., Reeves, S. and Hassler, S.W., 1998. Iridium anomaly but no shocked quartz from Late Archaean microkrystite layer: oceanic impact ejecta? Geology, 26: 195-198. Simonson, B.M., Hassler, S.W. and Beukes, N., 1999. Late Archaean impact spherule layer in South Africa that may correlate with a Western Australian layer. In: B.O. Dressler-Burkhard and V.L. Sharpton (Eds.), Large Meteorite Impacts and Planetary Evolution II. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 339, pp. 249-261. Simonson, B.M., Hornstein, M. and Hassler, S.W., 2000a. Particles in late Archaean Carawine Dolomite (Western Australia) resemble Muong Nong-type tektites. In: I. Gilmour and C. Koeberl (Eds.), Impacts and the Early Earth. Lecture Notes in Earth Sci., 92. Springer-Verlag, Berlin, pp. 181-214. Simonson, B.M., Davies, D. and Hassler, S.W. 2000b. Discovery of a layer of probable impact melt spherules in the late Archaean Jeerinah Formation (Fortescue Group, Western Australia). Aust. J. Earth Sci., 47:315-325. Simonson, B.M., Koeberl, C., McDonald, I. and Reimold, W.U., 2000c. Geochemical evidence for an impact origin for a Late Archaean spherule layer, Transvaal Supergroup, South Africa. Geology, 28:1103-1106. Simonson, B.M., Cardiff, M. and Schubel, K.A., 2001. New evidence that a spherule layer in the Late Archaean Jeerinah Formation of Western Australia was produced by a major impact. Abstracts, 32nd Lunar and Planetary Science Conference, Houston, Texas, 2001, abstract no. 1141. Simonson, B.M., Hassler, S.W., Smit, J. and Sumner, D., 2002. How many Late Archaean impacts are recorded in the Hamersley basin of Western Australia? Abstracts, 33rd Lunar and Planetary Science Conference, Houston, Texas, 2002, abstract no. 1772. Simpson, E.L. and Eriksson, K.A., 1991. Depositional facies and controls on parasequence development in siliciclastic tidal deposits from the early Proterozoic, upper Mount Guide Quartzite, Mount Isa Inlier, Australia. In: D.G. Smith, G.E. Reinson, B.A. Zaitlin and R.A. Rahmani (Eds.), Clastic Tidal Sedimentology. Can. Soc. Petrol. Geologists, Calgary, Mem., 16, pp. 371-387. Simpson, E.L. and Eriksson, K.A., 1993. Thin eolianites interbedded within a fluvial and marine succession: early Proterozoic Whitworth Formation, Mount Isa Inlier, Australia. Sediment. Geol., 87: 39-62. Simpson, E.L., Kuklis, C.A., Eriksson, K.A., Eriksson, P.G. and Bumby, A.J., 1999. Interbedded eolian dune, interdune and playa deposits, ~ 1.8 Ga Makgabeng Formation, Waterberg Group, South Africa: climate implications. Geol. Soc. Amer. Abstr. Progr., 31 (7): 284-285. Simpson, E.L., Eriksson, K.A., Eriksson, P.G. and Bumby, A.J., 2002. Eolian dune degradation and generation of massive sandstones in the Paleoproterozoic Makgabeng Formation, Waterberg Group, South Africa. J. Sediment. Res., 72: 40--45.

894

References

Simpson, E.L., Eriksson, K.A., Kuklis, C.A., Eriksson, EG., Bumby, A.J. and van Jaarsveld, C.E (in review). A ~ 1.8 Ga Playa Deposit in the Makgabeng Formation, Waterberg Group, South Africa. Sediment. Geol. Simpson, P.R. and Bowles, J.EW., 1981. Detrital uraninite and pyrite: are they evidence for a reducing atmosphere? In: Armstrong, EC. (Ed.), Genesis of uranium- and gold-bearing Precambrian quartz-pebble conglomerates. U.S. Geol. Surv. Prof. Paper, 1161-A-BB, pp. S 1-S 12. Sims, P.K. and Peterman, Z.E., 1986. Early Proterozoic Central Plains Orogen. Geology, 14: 488--491. Sims, P.K., van Schmus, W.R., Schulz, K.J. and Peterman, Z.E., 1989. Tectono-stratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean orogen. Can. J. Earth Sci., 26: 2145-2158. Singh, J. and Johannes, W., 1996. Dehydration melting of tonalites. Part I. Beginning of melting. Contrib. Mineral. Petrol., 125:16-25. Singh, S.P., 1998. Precambrian stratigraphy of Biharuan overview. In: B.S. Paliwal (Ed.), The Indian Precambrian. Scientific Publishers, Jodhpur, India, pp. 1-45. Skilling, I.P., 1993. Incremental caldera collapse of Suswa volcano, Gregory Rift Valley, Kenya. J. Geol. Soc. London, 150: 885-896. Skulski, T., Wares, R. and Smith, A., 1993. Early Proterozoic (1.88-1.87 Ga) tholeiitic magmatism in the new Quebec orogen. Can. J. Earth Sci., 30: 1505-1520. Skulski, T., Sandeman, H., Sanborn-Barrie, M., MacHattie, T., Hyde, D., Johnstone, S., Panagapko, D. and Byrne, D., 2002. Contrasting crustal domains in the Committee Bay belt, Walker Lake--Arrowsmith River area, central Nunavut. Current Research 2002-C 11. Geol. Surv. Can., 11 pp. (CD-ROM). Sleep, N.H., 1992. Archean plate tectonics: what can be learned from continental geology? Can. J. Earth Sci., 29:2066-2071. Sleep, N.H., 1994. Martian plate tectonics. J. Geophys. Res., 99: 5639-5655. Sleep, N.H., 2002. Oxygenating the atmosphere. Nature, 410:317-319. Sleep, N.H. and Windley, B.E, 1982. Archean plate tectonics; constraints and inferences. J. Geol., 90: 363-379. Sleep, N.H., Zahnle, K.J., Kasting, J.E and Morowitz, H.J., 1989. Annihilation of ecosystems by large asteroid impacts on the early Earth. Nature, 342: 139-142. Sloss, L.L., 1963. Sequence in the cratonic interior of North America. Geol. Soc. Amer. Bull., 74: 93-114. Sloss, L.L., Krumbein, W.C. and Dapples, E.C., 1949. Integrated facies analysis. In: C.R. Longwell (Ed.), Sedimentary facies in geologic history. Geol. Soc. Amer. Mem., 39, pp. 91-124. Smit, C.A., 1984. Petrochemistry and tectonic significance of a chromite-bearing ultramafic suite of rocks in the Limpopo Metamorphic Complex, South Africa. Trans. Geol. Soc. S. Afr., 87: 303-314. Smit, C.A., Roering, C. and van Reenen, D.D., 1992. The structural framework of the southern margin of the Limpopo Belt, South Africa. Precambrian Res., 55:51-67. Smit, J., 1999. The global stratigraphy of the Cretaceous-Tertiary boundary impact ejecta. Annu. Rev. Earth Planet. Sci., 27:75-113. Smit, J., Alvarez, W., Montanari, A., Swinburne, N., Van Kempen, T.M., Klaver, G.T. and Lustenhouwer W.J., 1992. "Tektites" and microkrystites at the Cretaceous Tertiary boundary: two strewn fields, one crater? Proc. Lunar Plan. Sci., 22: 87-100. Smit, J., Rocchia, R. and Robin, E., 1997. Preliminary iridium analysis from a graded spherule layer at the K/T boundary and late Eocene ejecta from ODP sites 1049, 1052, 1053, Blake Nose,

References

895

Florida. Abstracts with Programs, Annu. Meeting Geol. Soc. Amer., Salt-Lake City, Utah, 29, p. 141. Smith, D.M., 2001. A petrological and geochemical study of volcanic flows from the Paleoproterozoic Christopher Island Formation, Thirty Mile Lake, Nunavut Territory, Canada. B.Sc. Thesis (unpubl.), Carleton Univ., Ottawa, 66 pp. Smith, H.S. and Erlank, A.J., 1982. Geochemistry and petrogenesis of komatiites from the Barberton greenstone belt, South Africa. In: N.T. Arndt, and E.G. Nisbet (Eds.). Komatiites. George Allen and Unwin, London, pp. 347-398. Smith, H.S., Erlank, A.J. and Duncan, A.R., 1980. Geochemistry of some ultramafic komatiite lava flows from the Barberton Mountain Land, South Africa. Precambrian Res., 11: 399-415. Smith, R.D., Cameron, K.L., McDowell, F.W., Niemeyer, S. and Sampson, D.E., 1996. Generation of voluminous silicic magmas and formation of mid-Cenozoic crust beneath north-central Mexico: evidence from ignimbrites, associated lavas, deep crustal granulites, and mantle pyroxenites. Contrib. Miner. Petrol., 123: 375-389. Smith, R.K., Tremallo, R.L. and Lofgren, G.E., 2001. Growth of megaspherulites in a rhyolitic vitrophyre. Amer. Mineral., 86: 589--600. Smith, R.L. and Bailey, R.A., 1968. Resurgent cauldrons. In: R.R. Coats, R.L. Hay and C.A. Anderson (Eds.), Studies in Volcanology--A Memoir in Honor of Howel Williams. Geol. Soc. Amer. Mem., 116, pp. 613-662. Smith, T. and Batiza, R., 1989. New field and laboratory evidence for the origin of hyaloclastite flows on seamount summits. Volcanol. Bull., 51:96-114. Smithies, R.H., 2000. The Archaean tonalite-trondhjemite-granodiorite (TTG) series is not an analogue of Cenozoic adakite. Earth Planet. Sci. Lett., 182:115-125. Smithies, R.H. and Champion, D.C., 2000. The Archaean high-Mg diorite suite: links to tonalitetrondhjemite-granodiorite magmatism and implications for Early Archaean crustal growth. J. Petrol., 41: 1653-1671. Smithies, R.H., Hickman, A.H. and Nelson, D.R., 1999. New constrains on the evolution of the Mallina Basin, and their bearing on relationships between the contrasting eastern and western granite-greenstone terranes of the Archaean Pilbara Craton, Western Australia. Precambrian Res., 94:11-28. Smithies, R.H., Nelson, D.R. and Pike, G., 2001. Development of the Archaean Mallina Basin, Pilbara Craton, northwestern Australia: a study of detrital and inherited zircon ages. Sediment. Geol., 141/142: 79-94. Smithies, R.H., Champion, D.C. and Cassidy, K.E, in press. Formation of Earth's early continental crust. Precambrian Res. Smithson, S.B., Wenzel, F., Ganchin, Y.V. and Morozov, I.V., 2000. Seismic results at Kola and KTB deep scientific boreholes: velocities, reflections, fluids and crustal composition. Tectonophysics, 329:301-317. Smrekar, S.E. and Stofan, E.R., 1997. Corona formation and heat loss on Venus by coupled upwelling and delamination. Science, 277: 1289-1294. Smrekar, S.E. and Stofan, E.R., 1999. Origin of corona-dominated topographic rises on Venus. Icarus, 139:100-115. Sneh, A., 1983. Desert stream sequences in the Sinai Peninsula. J. Sediment. Petrol., 53: 1271-1279. Snow, J.E. and Schmidt, G., 1998. Constraints on Earth accretion deduced from noble metals in the oceanic mantle. Nature, 391: 166-169.

896

References

Snyder, G.A., Borg, L.E., Nyquist, L.E. and Taylor, L.A., 2000. Chronology and isotopic constraints on lunar evolution. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 361-395. Soegaard, K. and Eriksson, K.A., 1985. Evidence of tide, storm, and wave interaction on a Precambrian shelf: the 1.7 Ga Ortega Group, New Mexico. J. Sediment. Petrol., 55: 672--684. Sohl, L.E., Christie-Blick, N. and Kent, D.V., 1999. Paleomagnetic polarity reversals in Marinoan (ca. 600 Ma) glacial deposits of Australia: implications for the duration of low-latitude glaciation in Neoproterozoic time. Geol. Soc. Amer. Bull., 111 : 1120-1139. Solomon S.C., Smrekar, S.E., Bindschadler, L., Grimm, R.E., Kaula, W.M., McGill, G.E., Phillips, R.J., Saunders, R.S., Schubert, G., Squyres, S.W. and Stofan, E.R., 1992. Venus tectonics: an overview of Magellan observations. J. Geophys. Res., 97: 13,199-13,256. Solomotov, V.S., 2000. Fluid dynamics of the terrestrial magma ocean. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona Press, Tucson, Arizona, pp. 323-338. Somers, G.E and Brown, M., 1978. The affinity of trichomes of blue-green algae for calcium ions. Estuaries, 1: 17-28. SCnderholm, M. and Tirsgaard, H., 1998. Proterozoic fluvial styles: response to base-level changes and climate (Riveradal sandstones, eastern North Greenland). Sediment. Geol., 120: 257-274. Sonett, C.P. and Chan, M.A., 1998. Neoproterozoic Earth-Moon dynamics: Rework of the 900 Ma Big Cottonwood Canyon tidal rhythmites. Geophys. Res. Lett., 25: 539-542. Sonett, C.P., Kvale, E.P., Zakharian, A., Chan, M.A. and Demko, T.M., 1996a. Late Proterozoic and Palaeozoic tides, retreat of the Moon, and rotation of the Earth. Science, 273:100-104. Sonett, C.P., Zakharian, A. and Kvale, E.P., 1996b. Ancient tides and length of day: correction. Science, 274: 1068-1069. Sorjonen-Ward, P., Claue-Long, J. and Huhma, H., 1994. SHRIMP isotope studies of granulite zircons and their relevance to Early Proterozoic tectonics in northern Fennoscandia. U.S. Geol. Surv. Circ., 1107, pp. 299. Spada, G., Ricard, Y. and Sabadini, R., 1992. Excitation of true polar wander by subduction. Nature, 360: 425--454. Spencer, A.M., 1971. Late Pre-Cambrian glaciation in Scotland. Geol. Soc. London Mem., 6, 99 pp. Sproule, R.A., Lesher, C.M., Ayer, J.A., Thurston, P.C. and Herzberg, C.T., 2002. Secular variation in the geochemistry of komatiitic rocks from the Abitibi greenstone belt, Canada. Precambrian Res., 115: 153-186. Squire, R.J., Cas, R.A.E, Clout, J.M.F. and Behets, R., 1998. Volcanology of the Archaean Lunnon Basalt and its relevance to nickel sulfide-bearing trough structures at Kambalda, Western Australia. Aust. J. Earth Sci., 45: 695-715. Squyres, S. and Kasting, J.E, 1994. Early Mars: how warm and how wet? Science, 265: 744-749. Squyers, S.W., Reynolds, R.T., Cassen, P.M. and Peale, S.J., 1983. Liquid water and active resurfacing on Europa. Nature, 301: 225-226. Srinivasan, G., Sahijpal, S., Ulyanov, A.A. and Goswami, J.N., 1996. Ion microprobe studies of Efremovka CAIs" II. Potassium isotope composition and 41 Ca in the early solar system. Geochim. Cosmochim. Acta, 60:1823-1835. Srinivasan, G., Goswami, J.N. and Bhandari, N., 1999. 26A1 in eucrite Pipla Kalan: plausible heat source and formation chronology. Science, 284:1348-1350. Srinivasan, R. and Ojakangas, R.W., 1986. Sedimentology of quartz-pebble conglomerates and quartzites of the Archean Bababudan Group, Dharwar Craton, South India: evidence for early crustal stability. J. Geol., 94:199-214.

References

897

Stacey, J.S. and Kramers, J.D., 1975. Approximation of terrestrial lead isotope evolution by a twostage model. Earth Planet. Sci. Lett., 26:207-221. Stal, L.J., 2000. Cyanobacterial mats and stromatolites. In: B.A. Whitton and M. Potts (Eds.), The Ecology of Cyanobacteria. Kluver Academic Publishers, Dordrecht, pp. 61-120. Stanistreet, I.G. and McCarthy, T.S., 1991. Changing tectono-sedimentary scenarios relevant to the development of the Late Archaean Witwatersrand Basin. J. Afr. Earth Sci., 13(1): 65-81. Staudacher, T., 1987. Upper mantle origin for Harding Well noble gases. Nature, 325: 605-7. Staudigel, H. and Schmincke, H.-U., 1984. The Pliocene seamount series of La Palma/Canary Islands. J. Geophys. Res., 89:11,195-11,215. Staudt, W.J. and Schoonen, M.A.A., 1995. Sulfate incorporation into sedimentary carbonates. In: M.A. Vairavamurthy and M.A.A. Schoonen (Eds.), Geochemical Transformations of Sedimentary Sulfur, Symposium Series 612. Amer. Chem. Soc., Washington, DC, pp. 332-345. Stauffer, M.R., 1984. Manikewan: an Early Proterozoic ocean in central Canada, its igneous history and orogenic closure. Precambrian Res., 25:257-281. Steiger, R.H. and Wasserburg, G.J., 1969. Comparative U-Th-Pb systematics in 2.7 • 10 yr. plutons of different geological histories. Geochim. Cosmochim. Acta, 33: 1213-1232. Stein, M. and Goldstein, S.L., 1996. From plume head to continental lithosphere in the ArabianNubian Shield. Nature, 382: 773-778. Stein, M. and Hofmann, A.W., 1994. Mantle plumes and episodic crustal growth. Nature, 372: 63-68. Steinmann, G., 1906. Geologische Beobachtungen in den Alpen, II. Die schardtsche Uberfaltungstheorie und die geologische Bedeutung der Tiefseeabs~itze und der ophiolitische Massenngesteine. Bericht Naturforschung Gesellschaft Freiburg, i., band 16, pp. 44-65. Steinmann, G., 1927. Die ophiolitischen Zonen in den Mediterranean Kettengebirgen. 14th International Geological Congress, Madrid, Spain, Compte Rendu, 2: 638-667. Stephansson, O., 1977. Granite diapirism in Archean rocks. J. Geol. Soc. London, 133:357-361. Stern, C.R., Hanson, G.N. and Shirey, S.B., 1989. Petrogenesis of mantle-derived, LILE enriched Archaean monzodiorites and trachyandesites (sanukitoids) in southwestern Superior Province. Can. J. Earth. Sci., 26: 1688-1712. Stern, R. and Lucas, S., 1994. U-Pb zircon age constraints on the early tectonic history of the Flin Flon accretionary collage, Saskatchewan. Report 1994-F, Geol. Surv. Can., pp. 75-86. Stern, R.J., 1994. Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Annu. Rev. Earth Planet. Sci., 22:319-351. Stern, R.J. and Abdelsalam, M.G., 1998. Formation of juvenile continental crust in the ArabianNubian Shield: evidence from granitic rocks of the Nakasib suture, NE Sudan. Geol. Rundsch., 8: 150-160. Stern, R.A. and Bleeker, W., 1998. Age of the world's oldest rocks refined using Canada's SHRIMP: the Acasta gneiss complex, Northwest Territories, Canada. Geosci. Can., 25:27-31. Stem, R.J. and Kr6ner, A., 1993. Late Precambrian crustal evolution in NE Sudan: isotopic and geochronologic constraints. J. Geol., 10: 555-574. Stern R.A., Syme, E.C., Bailes, A.H. and Lucas, S.B., 1995a. Paleoproterozoic (1.90-1.86 Ga) arc volcanism in the Flin Flon belt, Trans-Hudson Orogen, Canada. Contrib. Miner. Petrol., 11: 117-141. Stem, R.A., Syme. S.C. and Lucas, S.B., 1995b. Geochemistry of 1.9 Ga MORB- and OIB-like basalts from the Amisk collage, Flin Flon belt, Canada: evidence for an intra-oceanic origin. Geochim. Cosmochim. Acta, 59:3131-3154.

898

References

Stetter, K.O., 1998. Hyperthermophiles and their possible role as ancestors of modem life. In: A. Brack (Ed.), The Molecular Origins of Life. Cambridge Univ. Press, Cambridge, pp. 315-335. Stettler, E.H., Van Schalkwyk, J.F., Stevens, G. and Pretorius, S., 1992. Tectonic model for the evolution of the Limpopo Mobile belt. Precambrian Res., 55: 539-552. Stettler. E.H., de Beer, J.H., Eberle, D., Ludden, J. and Mareschal, M., 1997. Geophysics and Deep Structures. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Science Publications, Oxford, pp. 339-375. Stevens, T.O. and McKinley, J.P., 1995. Lithotrophic microbial ecosystems in deep basalt aquifers. Science, 270: 450-454. Stevenson, D.J., 1987. Origin of the M o o n I t h e collision hypothesis. Annu. Rev. Earth Planet. Sci., 15: 271-315. Stewart, G.R., 2000. Outstanding questions for the Giant Impact Hypothesis. In: R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon. Univ. Arizona, Tucson, Arizona, pp. 217-223. Stewart, J.H., 1972. Initial deposits in the Cordilleran geosyncline: evidence of a late Precambrian (< 850 m.y.) continental separation. Geol. Soc. Amer. Bull., 83: 1345-1360. Stewart, J.H., Gehrels, G.E., Bart, A.P., Link, P.K., Christie-Blick, N. and Wrucke, C.T., 2001. Detrital zircon provenance of Mesoproterozoic to Cambrian arenites in the western United States and northwestern Mexico. Geol. Soc. Amer. Bull., 113: 1343-1356. Stix, J., 1991, Subaqueous, intermediate to silicic-composition explosive volcanism; a review. Earth Sci. Rev., 31:21-53. Stoesser, D.B. and Camp, V.E., 1985. Pan-African microplate accretion of the Arabian Shield. Geol. Soc. Amer. Bull., 96:817-826. Stoffers, P., Wright, I.C., De Ronde, C., Hannington, M., Villinger, H., Herzig, P. and The Shipboard Scientific Party, 1999. Little-studied arc-backarc system in the spotlight. Eos, 80(32): 353-354. St.-Onge, M.R., Lucas, S.B. and Scott, D.J., 1997. The Ungava Orogen and the Cape Smith thrust belt. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 772-780. St.-Onge, M.R., Lukas, S.B., Scott, D.J. and Wodichka, N., 1999. Upper and lower plate juxtaposition, deformation and metamorphism during crustal convergence, Trans-Hudson Orogen (Quebeck-Baffin segment), Canada. Precambrian Res., 93: 27-49. Storey, B.C., 1995. The role of mantle plumes in continental breakup: case histories from Gondwanaland. Nature, 377:301-308. Storey, M., Mahoney, J.J., Kroenke, L.W. and Saunders, A.D., 1991. Are oceanic plateaus sites of komatiite formation. Geology, 19: 376-379. Stothers, R.B. and Rampino, M.R., 1990. Periodicity in flood basalts, mass extinctions and impacts: a statistical view and a model. In: V.L. Sharpton and P.D. Ward (Eds.), Global Catastrophes in Earth History, An Interdisciplinary Conference on Impacts, Volcanism, and Mass Mortality. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper., 247, pp. 9-18. Stott, G.M., 1997. The Superior Province, Canada. In: M.J. de Wit and L.D. Ashwal. (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 480-507. Stowe, C.W., 1974. Alpine-type structure in the Rhodesian basement complex at Selukwe. J. Geol. Soc. London, 130:411-425. Stowe, C.W., 1984. The early Archaean Selukwe nappe, Zimbabwe. In A. Kr6ner and R. Greiling (Eds.), Precambrian Tectonics Illustrated. E. Schweizerbart'sche-Verlag, Stuttgart, pp. 41-56. Strauss, H., 1993. The sulfur isotopic record of Precambrian sulfates: new data and a critical evaluation of the existing record. Precambrian Res., 63: 225-246.

References

899

Strauss, H., 1997. The isotopic composition of sedimentary sulfur through time. Palaeogeog., Palaeoclimat., Palaeoecol., 132:97-118. Strauss, H., 1999. Geological evolution from isotope proxy signals--sulfur. Chem. Geol., 161: 89-101. Strauss, H., 2002. The isotopic composition of Precambrian sulphides--seawater chemistry and biological evolution. In: W. Altermann and P.L. Corocoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. IAS Spec. Publ., 33. Blackwell, Oxford, pp. 67-105. Strauss, H. and Moore, T.B., 1992. Abundances and isotopic compositions of carbon and sulfur species in whole rock and kerogen samples. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, pp. 709-798. Strauss, H. and Schieber, J., 1989. A sulfur isotope study of pyrite genesis: the Mid-Proterozoic Newland formation, Belt Supergroup, Montana. Geochim. Cosmochim. Acta, 54: 197-204. Strauss, H. and Schieber, J., 1990. A sulfur isotope study of pyrite genesis: the Mid-Proterozoic Newland Formation, Belt Supergroup, Montana. Geochim. Cosmochim. Acta, 54: 197-204. Strauss, H., Des Marais, D.J., Summons, R.E. and Hayes, J.M., 1992a. The carbon-isotopic record. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 117-128. Strauss, H., Des Marais, D.J., Hayes, J.M. and Summons, R.E., 1992b. Concentrations of organic carbon and maturities, and elemental compositions of kerogens. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere, A Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 95-99. Strom, S.E., Edwards, S. and Skrutskie, M.E, 1993. Evolutionary time scales for circumstellar disks associated with intermediate- and solar-type stars. In: E.H. Levy and J.I. Lunine (Eds.), Protostars and Protoplanets III. Univ. Arizona Press, Tucson, Arizona, pp. 837-866. Stumm, W. and Morgan, J.J., 1996. Aquatic Chemistry, Chemical Equilibria and Rates in Natural Waters, 3rd edn. John Wiley and Sons, New York, 1022 pp. Sturm, M. and Matter, A., 1978. Turbidites and varves in Lake Brienz (Switzerland): deposition of clastic detritus by density currents. In: A. Matter and M.E. Tucker (Eds.), Modern and Ancient Lake Sediments. IAS Spec. Publ., 2. Blackwell, Oxford, pp. 145-166. Suarez, M., Herve, M. and Puig, A., 1987. Cretaceous diapiric plutonism in the southern Cordillera, Chile. Geol. Mag., 124: 569-575. Sugitani, K., 1992. Geochemical characteristics of Archaean cherts and other sedimentary rocks in the Pilbara Block, W. Australia: evidence for Archaean seawater enriched in hydrothermallyderived iron and silica. Precambrian Res., 57:21-47. Sugitani, K., Horiuchi, Y., Adachi, M. and Sugisaki, R., 1996. Anomalously low A1203/TiO2 values for Archaean cherts from the Pilbara Block, Western Australia--possible evidence for extensive chemical weathering on the early Earth. Precambrian Res., 80: 49-76. Sugitani, K., Yamamoto, K., Wada, H., Binu-Lal, S.S. and Yoneshige, M., 2002. Geochemistry of Archean carbonaceous cherts deposited at an immature island-arc setting in the Pilbara block, Western Australia. Sediment. Geol., 151: 45-66. Sukhanov, M.K. and Zhuravlev, D.Z., 1989. Sm-Nd dating of the Dzhugdzhur Precambrian anorthosites. Trans. USSR Acad. Sci., 304(2): 964-968. Summons, R.E., 1992. Abundance and composition of extractable organic matter. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere, A Multidisciplinary Study. Cambridge Univ. Press, New York, pp. 101-115.

900

References

Summons, R.E., Jahnke, L.L., Hope, J.M. and Logan, G.A., 1999.2-Methylhopanoids as biomarkers for cyanobacterial oxygenic photosynthesis. Nature, 400: 554-557. Sumner, D.Y., 1997. Carbonate precipitation and oxygen stratification in the late Archean seawater as deduced from facies and stratigraphy of the Gamohaan and Frisco Formations, Transvaal Supergroup, South Africa. Amer. J. Sci., 297: 455-487. Sumner, D.Y., 2000. Microbial vs. environmental influences on the morphology of late Archean fenestrate microbialites. In: R.E. Riding and S.M. Awramik (Eds.), Microbial Sediments. SpringerVerlag, Berlin, pp. 307-314. Sumner, D.Y., 2002. Decimetre-thick encrustations of calcite and aragonite on the sea-floor and implications for Neoarchean and Neoproterozoic ocean chemistry. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. IAS Spec. Publ., 33. Blackwell, Oxford, pp. 107-122. Sumner, D.Y. and Grotzinger, J.P., 1996a. Herringbone Calcite: petrography and environmental significance. J. Sediment. Res., A 66(3): 419-429. Sumner, D.Y. and Grotzinger, J.P., 1996b. Were kinetics of Archean calcium carbonate precipitation related to oxygen concentration? Geology, 24:119-122. Sumner, D.Y., Kirschvink, J.L. and Runnegar, B.N., 1987. Soft-sediment paleomagnetic field tests of late Precambrian glaciogenic sediments (abstract). Eos, 68:1251. Sun, S.-S., 1984. Chemical characteristics of Archean ultramafic and mafic volcanic rocks: impli- : cations for mantle compositions and evolution. In: A. Kroner, G.N. Hanson and A.M. Goodwin (Eds.), Archean Geochemistry. Springer-Verlag, Berlin, pp. 25-46. Sun, S.-S. and McDonough, W.E 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: A.D. Saunders and M.J. Norry (Eds.), Magmatism in the Ocean Basins. Geol. Soc. London, Spec. Publ., 42, 313-345. Sun. S-S. and Nesbitt, R.W., 1978. Geochemical regularities and genetic significance of ophiolitic basalts. Geology, 6: 689-693. Stindermann, J., 1982. The resonance behaviour of the world ocean. In: P. Brosche and J. Stindermann (Eds.), Tidal Friction and the Earth's Rotation II. Springer-Verlag, Berlin, pp. 165-174. Sundquist, E., 1985. Geological perspectives on Carbon Dioxide and the Carbon cycle. In: E. Sundquist and W. Broecker (Eds.), The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present. Amer. Geophys. Union, Washington, pp. 5-59. Sutton, A.N., Blake, S., Wilson, C.J.N. and Charlier, B.L.A., 2000. Late Quaternary evolution of a hyperactive rhyolite magmatic system: Taupo volcanic center, New Zealand. J. Geol. Soc., London, 157: 537-552. Svensson, U., 1992. Die L6sungs- und Abscheidungskinetik nattirlicher Calcitminerale in w~issrigen CO2-L6sungen nahe dem Gleichgewicht. Dissertation, Univ. Bremen, Fachbereich Geowissenschaften, Selbstverlag, Bremen, 113 pp. Swager, C., Witt, W.K., Griffin, T.J., Ahmat, A.L., Hunter, W.M., McGoldrick, EJ. and Wyche, S., 1990. A regional overview of the Late Archaean granite-greenstones of the Kalgoorlie terrane. In: S.E. Ho, J.E. Glover, J.S. Myers and J.R. Muhling (Eds.), Third International Archaean Symposium, Excur. Guidebook. Dept. Geol., Univ. Western Aust., Spec. Publ., 21, pp. 205-303. Swager, C.E, 1997. Tectono-stratigraphy of late Archaean greenstone terranes in the southern Eastern Goldfields, Western Australia. Precambrian Res., 83:11-42. Sweet, I.E, 1988. Early Proterozoic stream deposits: braided or meandering---evidence from central Australia. Sediment. Geol., 58: 277-293.

References

901

Sweetapple, M.T. and Collins, L.E, 2002. Genetic framework for the classification and distribution of Archean rare metal pegmatites in the North Pilbara Craton, Western Australia. Econ. Geol., 97: 873-895. Swezey, C., Kocurek, G., Lancaster, N., Deynoux, M., Blum, M., Price, D. and Pion, J.C., 1999. Response of aeolian systems to Holocene climatic and hydrologic changes on the northern margin of the Sahara; a high-resolution record from the Chott Rharsa Basin, Tunisia. The Holocene, 9: 141-147. Swift, D.J.P., 1975. Barrier-island genesis: evidence from the central Atlantic shelf, eastern U.S.A. Sediment. Geol., 14: 1-43. Swift, D.J.P. and Thorne, J.A., 1991. Sedimentation on continental margins I: a general model for shelf sedimentation. In: D.J.P. Swift, G.E Oertel, R.W. Tillman and J.A. Thorne (Eds.), Shelf Sand and Sandstone Bodies--Geometry, Facies and Sequence Stratigraphy. IAS Spec. Publ., 14. B lackwell, Oxford, pp. 3-31. Swift, D.J.P., Kofoed, J.W., Saulsbury, EB. and Sears, P.C., 1972. Holocene evolution of the shelf surface, central and southern Atlantic shelf of North America. In: D.J.P. Swift, D.B. Duane and O.H. Pilkey (Eds.), Shelf Sediment Transport: Process and Pattern. Dowden, Hutchinson and Ross, Stroudsburg, Pennsylvania, pp. 499-574. Swindle, T.D., Caffee, M.W., Hohenberg, C.M. and Taylor, S.R., 1986. I-Pu-Xe dating and relative ages of the Earth and moon. In: W.K. Hartman, R.J. Phillips and G.J. Taylor (Eds.), Origin of the Moon. Lunar and Planetary Sci. Inst., Houston, Texas, pp. 331-337. Sylvester, P.J. and Attoh, K., 1992. Lithostratigraphy and composition of 2.1 Ga greenstone belts of the West African craton and their bearing on crustal evolution and the Archean-Proterozoic boundary. J. Geol., 100: 377-393. Sylvester, P.J., Harper, G.D., Byerly, G.R. and Thurston, P.C., 1997. Volcanic aspects. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Oxford Science Publications, Oxford, pp. 55-90. Syme, E.C., 1988. Athapapuskow Lake project. Report of Field Activities, 1988, Manitoba Energy and Mines, Min. Div., pp. 20-34. Syme, E.C., Lucas, S.B., Bailes, A.H. and Stem, R.A., 1999. Contrasting arc and MORB-like assemblages in the Paleoproterozoic Flin Flon belt, Manitoba, and the role of intra-arc extension in localizing volcanic-hosted massive sulphide deposits. Can. J. Earth Sci., 36:1767-1788. Symons, D.T.A., 1991. Paleomagnetism of the Proterozoic Wathaman batholith and the suturing of the Trans-Hudson orogen in Saskatchewan. Can. J. Earth Sci., 28:1931-1938. Tack, L., Fernandez-Alonso, M., Tahon, A., Wingate, M. and Baritt, S., 2002. The "Northeastern Kibaran Belt" (NKB) and its mineralisation reconsidered: new constraints from a revised lithostratigraphy, a GIS compilation of existing geological maps and a review of recently published as well as unpublished igneous emplacement ages in Burundi. Abstracts, 1 l th Quadrennial IAGOD Symposium and Geocongress 2002, Windhoek, 22-26 July 2002, Geol. Surv. Namibia (CD-ROM). Tackley, P.J., 1997. Effects of phase transitions on three-dimensional mantle convection. In: D.J. Crossley (Ed.), Earth's Deep Interior. Gordon and Breach, Amsterdam, pp. 273-336. Tackley, P.J., 2000. Mantle convection and plate tectonics: towards an integrated physical and chemical theory. Science, 288: 2002-2007. Tackley, P.J., Stevenson, D.J., Glatzmaier, G.A. and Schubert, G., 1994. Effects of an endothermic phase transition at 670 km depth in a spherical model of convection in the Earth's mantle. J. Geophys. Res., 99: 15,877-15,901. Tai, C.Y. and Chen, E-B., 1998. Polymorphism of CaCO3 precipitated in a constant composition environment. AIChE J., 44:1790-1798.

902

References

Tajika, E. and Matsui, T., 1993. Degassing history and carbon cycle of the Earth: from an impactinduced steam atmosphere to the present atmosphere. Lithos, 30: 267-280. Takayama, H., Tada, R., Matsui, T., Iturralde Vinent, M.A., Oji, T., Tajika, E., Kiyokawa, S., Garcia, D., Okada, H., Hasegawa, T. and Toyoda, K., 2000. Origin of the Penalver Formation in northwestern Cuba and its relation to KfF boundary impact event. Sediment. Geol., 135: 1-4. Tanner, W.E, 1971. Numerical estimates of ancient waves, water depths and fetch. Sedimentology, 16: 71-88. Taran, L.N. and Bogdanova, S.V., 2001. The Fennoscandia-Sarmatia junction in Belarus: new inferences from a PT-study. Tectonophysics, 339: 193-214. Tarney, J., Dalziel, I.W.D. and de Wit, M.J., 1976. Marginal basin "Rocas Verdas" complex from S. Chile: a model for Archean greenstone belt formation. In: B.E Windley (Ed.), The Early History of the Earth. Wiley, London, pp. 405-418. TassE, N., Lajoie, J. and Dimroth, E., 1978. The anatomy and interpretation of an Archean volcaniclastic sequence, Noranda region, QuEbec. Can. J. Earth Sci., 15: 874-888. Taylor, P.N. and Kalsbeek, F., 1990. Dating the metamorphism of Precambrian marbles: examples from Proterozoic mobile belts in Greenland. Chem. Geol., 86:21-28. Taylor, R.N., Nesbitt, R.W., Vidal, P., Harmon, R., Auvray, B. and Croudace, I.W., 1994. Mineralogy, chemistry, and genesis of the boninite series volcanics, Chichijima, Bonin Islands, Japan. J. Petrol., 35:577-617. Taylor, S., Lever, J.H. and Harvey, R.P., 2000. Numbers, types and compositions of an unbiased collection of cosmic spherules. Meteoritics Planet. Sci., 35:651-666. Taylor, S.R., 1987. The origin of the Moon. Amer. Scientist, 75: 469-477. Taylor, S.R. and Levinson, A.A., 1970. Geochemistry of tektites. Geochim. Cosmochim. Acta, 33: 1013-1147. Taylor, S.R. and McLennan, S.M., 1985. The Continental Crust: Its Composition, and Evolution. Blackwell, Oxford, 312 pp. Taylor, S.R. and McLennan, S.M., 1995. The geochemical evolution of the continental crust. Rev. Geophysics, 33: 241-265. Teixeira, W., SabatE, P., Barbosa, J., Noce, C.M. and Carneiro, M.A., 2000. Archaean and Palaeoproterozoic tectonic evolution of the Sao Francisco craton, Brazil. In: U.G. Cordani, E.J. Milani, A.T. Filho and D.A. Campos (Eds.), Tectonic Evolution of the South America. Proceedings, 31st International Geology Congress, Rio de Janeiro, pp. 101-137. Teixeira, J.B.G., Vasconcelos, P.M. and Misi, A., 2002. Geodynamic setting of orogenic gold deposits in the Atlantica paleocontinent. Abstracts, 11th Quadrennial IAGOD Symposium and Geocongress 2002, Windhoek, 22-26 July 2002, Geol. Surv. Namibia (CD-ROM). Terwindt, J.H.J., 1981. Origin and sequences of sedimentary structures in inshore mesotidal deposits of the North Sea. In: S.-D. Nio, R.T.E. Shuttenhelm and T.C.E. van Weering (Eds.), Holocene Marine Sedimentation in the North Sea Basin. IAS Spec. Publ., 5. Blackwell, Oxford, pp. 4-26. Tessier, B., 1993. Upper intertidal rhythmites in the Mont-Saint-Michel Bay (NW France): perspectives for paleoconstruction. Marine Geol., 110: 355-367. Tessier, B., Archer, A.W., Lanier, W.P. and Feldman, H.R., 1995. Comparison of ancient tidal rhythmites (Carboniferous of Kansas and Indiana, U.S.A.) with modern analogues (the Bay of MontSaint-Michel, France). In: B.W. Flemming and A. Bartholoma (Eds.), Tidal Signatures in Modern and Ancient Sediments. IAS Spec. Publ., 24. Blackwell, Oxford, pp. 259-271. Theriault, R.J. and Tella, S., 1997. Sm-Nd isotopic study on mafic volcanic rocks from the Rankin Inlet and Tavani regions, District of Keewatin, Northwest Territories. Current Research Paper 1997-F, Geol. Surv. Can., pp. 61-66.

References

903

Thiemens, M.H., 1999. Mass-independent isotope effects in planetary atmospheres and the early solar system. Science, 283:341-345. Thomas, R.J., 1989. A tale of two tectonic terranes. S. Afr. J. Geol., 92: 306-321. Thomas, R.J., Von Veh, M.W. and McCourt, S., 1993. The tectonic evolution of southern Africa: an overview. J. Afr. Earth Sci., 16: 5-24. Thomas, R.J., Cornell, D.H. and Armstrong, R.A., 1999. Provenance age and metamorphic history of the Quha Formation, Natal Metamorphic Province: a U-Th-Pb zircon SHRIMP study. S. Afr. J. Geol., 102: 83-88. Thompson, A.B., 1999. Melting of the continental crust: some thermal and petrological constraints on anatexis in continental collision zones and other tectonic settings. In: A. Castro, C. Fernandez and J.L. Vigneresse (Eds.), Understanding Granites: Integrating New and Classical Techniques. Geol. Soc. London Spec. Publ., 158, pp. 7-25. Thompson, A.B. and Ridley, J.R., 1987. Pressure-temperature-time ( P - T - t ) histories of orogenic belts. Phil. Trans. R. Soc. Lond., A321: 27-45. Thompson, EH., 1992. Proterozoic evolution of the Northern Thelon Tectonic Zone. Ph.D. Thesis (unpubl.), Princeton Univ., Princeton, New Jersey, 311 pp. Thomson, B., 1989. Petrology and stratigraphy of some texturally well preserved thin komatiites from Kambalda, Western Australia. Geol. Mag., 126: 249-261. Thorarinsson, S., 1967. Surtsey. The New Island in the North Atlantic. The Viking Press, New York, 47 pp. Thorkelson, D.J., Mortensen, J.K., Creaser, R.A., Davidson, G.J. and Abbott, J.G., 2001. Early Proterozoic magmatism in Yukon, Canada: constraints on the evolution of northwestern Laurentia. Can. J. Earth Sci., 38: 1479-1494. Thorne, A.M. and Seymour, D.D., 1991. Geology of the Ashburton Basin. Western Aust. Geol. Surv. Bull., 139, 141 pp. Thorpe, R.A., Hickman, A.H., Davis, D.W., Mortensen, J.K. and Trendall, A.E, 1992a. U-Pb zircon geochronology of Archaean felsic units in the Marble Bar region, Pilbara Craton, Western Australia. Precambrian Res., 56: 169-189. Thorpe, R.I., Hickman, A.H., Davis, D.W., Mortensen, J.K. and Trendall, A.E, 1992b. Constraints to models for lead evolution from precise zircon U/Pb geochronology for the Marble Bar region, Pilbara Craton, Western Australia. In: J.E. Glover and S.E. Ho (Eds.), The Archaean: Terrains, Processes and Metallogeny. Geol. Dept. Univ. Ext., Univ. Western Aust., Publ., 22, pp. 395-406. Throop, H.B., Bally, J., Esposito, L.W. and McCaughrean, M.J., 2001. Evidence for dust growth in young circumstellar disks. Science, 292:1686-1689. Thurston, EC., 1980. Subaerial volcanism in the Archean Uchi-Confederation Volcanic Belt. Precambrian Res., 12: 79-98. Thurston, EC., 1986. Volcanic cyclicity in mineral exploration; the caldera cycle and zoned magma chambers. In: J. Wood and H. Wallace (Eds.), Volcanology and Mineral Deposits. Ontario Geol. Surv., Misc. Paper, 129, pp. 104-123. Thurston, EC., 1994. Archean volcanic patterns. In: K.C. Condie (Ed.), Archean Crustal Evolution. Developments in Precambrian Geology, 11. Elsevier, Amsterdam, pp. 45-84. Thurston, EC., 2002. Autochthonous development of Superior Province greenstone belts? Precambrian Res., 115:11-36. Thurston, EC. and Chivers, K.M., 1990. Secular variation in greenstone sequence development emphasizing Superior Province, Canada. Precambrian Res., 46:21-58.

904

References

Thurston, RC., Ayres, L.D., Edwards, G.R., Gelinas, L., Ludden, J.N. and Verpaelst, R, 1985. Archean bimodal volcanism. In: L.D. Ayres, P.C. Thurston, K.D. Card and W. Weber (Eds.), Evolution of Archean Supracrustal Successions. Geol. Assoc. Can. Spec. Paper, 28, pp. 7-21. Thurston, P.C., Osmani, I.A. and Stone, D., 1991. Northwestern Superior Province: review and terrane analysis. In: P.C. Thurston, H.R. Williams, R.H. Sutcliffe and G.M. Stott (Eds.), Geology of Ontario. Ontario Geol. Surv., Spec. Vol. 4, pp. 81-142. Tilling, R.I. and Dvorak, J.J., 1993. Anatomy of a basalt volcano. Nature, 363: 125-133. Tilton, G.R., 1988. Age of the Solar System. In: J.F. Kerridge and M.S. Matthews (Eds.), Meteorites and the Early Solar System. Univ. Arizona Press, Tucson, pp. 259-275. Tirsgaard, H., 1993. The architecture of Precambrian high energy tidal channel deposits: an example from the Lyell Land Group (Eleonore Bay Supergroup), northeast Greenland. Sediment. Geol., 88: 137-152. Tirsgaard, H., 1996. Cyclic sedimentation of carbonate and siliciclastic deposits on a late Precambrian ramp: The Elisabeth Bjerg Formation (Eleonore Bay Supergroup), East Greenland. J. Sediment. Res., B66(4): 698-712. Tirsgaard, H. and Oxnevad, I.E.I., 1998. Preservation of pre-vegetational fluvio-aeolian deposits in a humid climatic setting: an example from the Middle Proterozoic Eriksfjord Formation, southwest Greenland. Sediment. Geol., 120:295-317. Tirsgaard, H. and SCnderholm, M., 1997. The sedimentary evolution and lithostratigraphy of the Late Precambrian Lyell Land Group (Eleonore Bay Supergroup), East and Northeast Greenland. Bull. Geol. Surv. Denmark and Greenland, 178, 60 pp. Tohver, E., van der Pluijm, B.A., Van der Voo, R., Rizzoto, G. and Scandolara, J.E., 2002. Paleogeography of the Amazon craton at 1.2 Ga: early Grenvillian collision with the Llano segment of Laurentia. Earth Planet. Sci. Lett., 199:185-200. Tomlinson, K.Y. and Condie, K.C., 2001. Archean mantle plumes: evidence from greenstone belt geochemistry. In: R.E. Ernst and K.L. Buchan (Eds.), Mantle Plumes: Their Identification Through Time. Geol. Soc. Amer. Spec. Paper, 352, pp. 341-357. Tomlinson, K.Y., Hughes, D.J., Thurston, P.C. and Hall, R.P., 1999. Plume magmatism and crustal growth at 2.9 to 3.0 Ga in the Steep Rock and Lumby Lake area, Western Superior Province. Lithos, 46: 103-136. Toomre, A., 1974. On the 'nearly diurnal wobble' of the Earth. Geophys. J. Roy. Astr. Soc., 38: 335-348. Tooth, S., 2000. Process, form and change in drytand rivers: a review of recent research. Earth Sci. Rev., 51: 67-107. Torsvik, T.H. and Rehnstr6m, E.F., 2001. Cambrian paleomagnetic data from Baltica: implications for true polar wander and Cambrian paleogeography. J. Geol. Soc. London, 158: 321-329. Torsvik, T.H. and Van der Voo, R., 2002. Refining Gondwana and Pangea palaeogeography: estimates of Phanerozoic non-dipole (octupole) fields. Geophys. J. Int., 151: 1-24. Torsvik, T.H., Meert, J.G. and Smethurst, M.A., 1998. Polar wander and the Cambrian. Science, 279: 9a. Touret, J.L.R. and Hartel, T.H.D., 1990. Synmetamorphic fluid inclusions in granulites. In: D. Vielzeuf and Ph. Vidal (Eds.), Granulites and Crustal Evolution. Kluwer Academic Publishers, Dordrecht, pp. 397-417. Towe, K.M., 1978. Early Precambrian oxygen: a case against photosynthesis. Nature, 217:657-661. Towe, K.M., 1990. Aerobic respiration in the Archean? Nature, 348: 54-56.

References

905

Towe, K.M., 1994. Earth's early atmosphere: constraints and opportunities for early evolution. In: S. Bengston (Ed.), Early Life on Earth, Nobel Symposium No. 84. Columbia Univ. Press, New York, pp. 36--47. Treloar, P.J. and Blenkinsop, T.G., 1995. Archaean deformation patterns in Zimbabwe: true indicators of Tibetan-style crustal extrusion or not? In: M.P. Coward and A.C. Reis (Eds.), Early Archaean Processes. Geol. Soc. London Spec. Publ., 95, pp. 78-108. Treloar, P.J., Coward, M.P. and Harris, N.B.W., 1992. Himalayan-Tibetan analogies for the evolution of the Zimbabwe craton and Limpopo Belt. Precambrian Res., 55: 571-587. Tremblay, L.P., 1972. Geology of the Beaverlodge Mining Area, Saskatchewan. Geol. Surv. Can. Mem., 367, pp. 1-265. Trendall, A.E, 1965. Discussion of "Origin of Precambrian Iron Formations". Econ. Geol., 60: 1065-1070. Trendall, A.E, 1968. Three great basins of Precambrian banded iron-formation deposition: a systematic comparison. Geol. Soc. Amer. Bull., 79: 1527-1544. Trendall, A.E, 1973a. Iron-formations of the Hamersley Group of Western Australia: type examples of varved Precambrian evaporites. Proceedings, Kiev Symposium: Genesis of Precambrian Iron and Manganese Deposits, 1970. UNESCO, Paris, Earth Sciences, 9, pp. 257-267. Trendall, A.F., 1973b. Varve cycles in the Weeli Wolli Formation of the Precambrian Hamersley Group, Western Australia. Econ. Geol., 68: 1089-1097. Trendall, A.E, 1976. Striated and faceted boulders from the Turee Creek Formation--evidence for a possible Huronian glaciation on the Australian continent. Western Australian Geological Surv., Annu. Report, 1975: 88-92. Trendall, A.E, 1981. The lower Proterozoic meteorite bore member, Hemersley Basin, Western Australia. In: M.J. Hambrey and W.B. Harland (Eds.), Earth's Pre-Pleistocene Glacial Record. Cambridge Univ. Press, Cambridge, pp. 555-557. Trendall, A.E, 1983a. Introduction. In: A.E Trendall and R.C. Morris (Eds.), Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 1-11. Trendall, A.E, 1983b. The Hamersley Basin. In: A.E Trendall and R.C. Morris (Eds.), IronFormation: Facts and Problems. Elsevier, Amsterdam, pp. 69-129. Trendall, A.E, 1990. Hamersley Basin. In: Geology and Mineral Resources of Western Australia. Western Australian Geol. Surv. Mem., 3, pp. 163-191. Trendall, A.E, 1995. Paradigms for the Pilbara. In: M.P. Coward and A.C. Reis (Eds.), Early Archaean Processes. Geol. Soc. London, Spec. Publ. 95, pp. 127-142. Trendall, A.E, 1998. Precise measurements of the depositional rate of a Precambrian banded iron formation (B IF) by two independent methods. Abstracts, Geosci. 98, Biennial Meeting Geol. Soc. London, Keele, pp. 48. Trendall, A.E, 2002. The significance of iron-formation in the Precambrian stratigraphic record. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Depositional Systems. IAS Spec. Publ., 44. Blackwell, Oxford, pp. 33-66. Trendall, A.F. and Blockley, J.G., 1970. The iron formations of the Precambrian Hamersley Group, Western Australia, with special reference to the associated crocidolite. Western Australia Geol. Surv., Bull. 119, 365 pp. Trendall, A.F. and Morris, R.C. (Eds.), 1983. Iron-Formation: Facts and Problems. Elsevier, Amsterdam, 558 pp. Trendall, A.E, Nelson, D.R., de Laeter, J.R. and Hassler, S.W., 1998. Precise zircon U-Pb ages from the Marra Mamba Iron Formation and Wittenoom Formation, Hamersley Group, Western Australia. Aust. J. Earth Sci., 45: 137-142.

906

References

Trieloff, M., Kunz, J., Clague, D.A., Harrison, D. and All~gre, C.J., 2000. The nature of pristine noble gases in mantle plumes. Science, 288:1036-1038. Trompette, R., de Alvarenga, C.J.S. and Walde, D., 1998. Geological evolution of the Neoproterozoic Corumb~i graben system (Brazil). Depositional context of the stratified Fe and Mn ores of the Jacadigo Group. J. South Amer. Earth Sci., 11: 587-597. Trotet, F., Jolivet, L. and Vidal, O., 2001. Tectono-metamorphic evolution of Syros and Sifnos islands (Cyclades, Greece). Tectonophysics, 338: 179-206. Trusheim, F., 1936. Wattenpapier. Natur und Volk, 66:103-106. Tsoar, H., 1982. Internal structure and surface geometry of longitudinal (seif) dunes. J. Sediment. Petrol., 52:823-831. Tsoar, H., 1983. Dynamic processes acting on a longitudinal (seif) sand dune. Sedimentology, 30: 567-578. Tucker, M.E., 1982. The Field Description of Sedimentary Rocks. Halstead, New York, 112 pp. Tucker, M.E., 1991. Sedimentary Petrology. Blackwell, Oxford, 260 pp. Tuisku, P. and Huhma, H., 1999. Geochronology of Lapland granulite belt and implications for the Fennoscandian assembly. The architecture and evolution of the Palaeoproterozoic Lapland-Kola orogen. Abstracts, SVEKALAPKO, a Europrobe project, Oulu, Finland, p. 65. Turner, C.C. and Walker, R.G., 1973. Sedimentology, stratigraphy, and crustal evolution of the Archean greenstone belt near Sioux Lookout, Ontario. Can. J. Earth Sci., 10: 817-845. Turner, E.C., Narbonne, G.M. and James, N.P., 1993. Neoproterozoic reef microstructures from the Little Dal Group, northwestern Canada. Geology, 21: 259-262. Turner, E.C., James, N.P. and Narbonne, G.M., 2000. Taphonomic control on microstructure in early Neoproterozoic reefal stromatolites and thrombolites. Palaios, 15:87-111. Turner, J.S., Huppert, H.E. and Sparks, R.S.J., 1986. Komatiites II: experimental and theoretical investigations of post-emplacement cooling and crystallization. J. Petrol., 27: 397-437. Turner, R.J.W., 1992. Formation of Phanerozoic stratiform sediment-hosted zinc-lead deposits: evidence for the critical role of ocean anoxia. Chem. Geol., 99: 165-188. Tyler, I.M. and Thorne, A.M., 1990. The northern margin of the Capricorn Orogen, Western Australia--an example of an early Proterozoic collision zone. J. Struc. Geol., 12: 685-701. Tyler, S.A. and Barghoorn, E.S., 1954. Occurrence of structurally preserved plants in pre-Cambrian rocks of the Canadian shield. Science, 119: 606-608. Tyler, S.A. and Barghoorn, E.S., 1963. Ambient pyrite grains in Precambrian cherts. Amer. J. Sci., 261: 424--432. Ueno, Y., Isozaki, Y., Yurimoto, H. and Maruyama, S., 2001. Carbon isotopic signatures of individual Archean microfossils(?) from Western Australia. Internat. Geol. Rev., 40: 196-212. Ullrich, L. and Van der Voo, R., 1981. Minimum continental velocities with respect to the pole since the Archean. Tectonophysics, 74:17-27. Umino, S., Lipman, P.W. and Obata, S., 2000. Subaqueous lava flow lobes, observed on ROV KAIKO dives off Hawaii. Geology, 28: 503-506. UNESCO, 1973. Genesis of Precambrian Iron and Manganese Deposits. Proceedings of the Kiev Symposium, August 1970. UNESCO, Paris, Earth Sci. 9, 382 pp. Unrug, R., 1984. The mid-Proterozoic Mporokoso Group of northern Zambia: stratigraphy, sedimentation and regional position. Precambrian Res., 24:99-121. Vail, J.R., 1985. Pan-African (late Precambrian) tectonic terrains and the reconstruction of the Arabian-Nubian Shield. Geology, 13: 839-842.

References

907

Vail, ER., Mitchum, R.M., Jr. and Thompson, S., III., 1977. Seismic stratigraphy and global changes of sea level. In: C.E. Payton (Ed.), Seismic StratigraphymApplications to Hydrocarbon Exploration. Amer. Assoc. Petrol. Geol. Mem., 26, pp. 49-212. Vail, ER., Audemard, E, Bowman, S.A., Eisner, EN. and Perez-Cruz, C., 1991. The stratigraphic signatures of tectonics, eustasy and sedimentology--an overview. In: G. Einsele, W: Ricken and A. Seilacher (Eds.), Cycles and Events in Stratigraphy. Springer-Verlag, Berlin, pp. 617-659. Valley, J.W., Peck, W.H., King, E.M. and Wilde, S.A., 2002. A cool early Earth. Geology, 30: 351-354. Van Berkel, J,T., Schwerdtner, W.M. and Torrance, J.G., 1984. Wall-and-basin structure: an intriguing tectonic prototype in the central Sverdrup Basin, Canadian Arctic Archipelago. Can. Petrol. Geol. Bull., 32: 343-358. Van Biljon, W.J., 1977. An introduction to the Limpopo Mobile Belt. Bull. Geol. Surv. Botswana, 12: 3-8. Van der Neut, M. and Eriksson, EG., 1999. Palaeohydrological parameters of a Proterozoic braided fluvial system (Wilgerivier Formation, Waterberg Group, South Africa) compared with a Phanerozoic example. In: N.D. Smith (Ed.), Proceedings 6th International Fluvial Conference. IAS Spec. Publ., 28. Blackwell, Oxford, pp. 381-392. Van der Voo, R., 1994. True polar wander during the middle Palaeozoic? Earth Planet. Sci. Lett., 122: 239-243. Van der Voo, R. and Torsvik, T.H., 2001. Evidence for late Paleozoic and Mesozoic non-dipole fields provides an explanation for the Pangea reconstruction problems. Earth Planet. Sci. Lett., 187: 71-81. Van der Westhuizen, W.A. and de Bruiyn, H., 2000. High temperature ash-flow wet sediment interaction in the Makwassie Formation, Ventersdorp Supergroup, South Africa. Precambrian Res., 101: 341-351. Van Haaften, W.M. and White, S.H., 1998. Evidence for multiphase deformation in the Archean basal Warrawoona Group in the Marble Bar area, East Pilbara, Western Australia. Precambrian Res., 88: 53-66. Van Hise, C.R. and Leith, C.K., 1911. The Geology of the Lake Superior Region. U.S. Geol. Surv., Monograph 52, 641 pp. Van Houten, EB., 1974. Northern Alpine molasse and similar Cenozoic sequences of Southern Europe. In: R.H. Dott, Jr. and R.H. Shaver (Eds.), Modem and Ancient Geosynclinal Sedimentation. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 19, pp. 260-273. Van Kooten, G.K., 1981. Pb and Sr systematics of ultrapotassic and basaltic rocks from the central Sierra Nevada, California. Contr. Mineral. Petrol., 76: 378-385. Van Kranendonk, M.J., 1996. Tectonic evolution of the Palaeoproterozoic Torngat orogen" evidence from pressure-temperature-time-deformation paths in the North River map area, Labrador. Tectonics, 15(4): 843-869. Van Kranendonk, M.J., 1998. Litho-tectonic and structural components of the North Shaw 1 : 100, 000 sheet, Archaean Pilbara Craton. Western Aust. Geol. Surv., Annual. Rev., 1997-1998: 63-70. Van Kranendonk, M.J., 2000. Geology of the North Shaw 1:100, 000 Sheet. Western Aust. Geol. Surv., 1 : 100000 Geol. Ser. Expl. Notes, 86 pp. Van Kranendonk, M.J., in press. Volcanic degassing, hydrothermal circulation and the flourishing of early life on Earth: new evidence from the c. 3.46 Ga Warrawoona Group, Pilbara Craton, Western Australia. Precambrian Res.

908

References

Van Kranendonk, M.J. and Collins, W.J., 1998. Timing and tectonic significance of Late Archaean, sinistral strike-slip deformation of the Central Pilbara Corridor, Pilbara Craton, Western Australia. Precambrian Res., 88: 207-232. Van Kranendonk, M.J. and Collins, W.J., 2001. A review of the evidence for vertical tectonics in the Archaean Pilbara Craton, Western Australia. In: K.F. Cassidy, J.M. Dunphy and M.J. Van Kranendonk (Eds.), Extended Abstract volume, 4th International Archaean Symposium, Perth, 2001. Aust. Geol. Surv. Org., Geoscience Australia, Record 2001/37, pp. 365-367. Van Kranendonk, M.J., Hickman, A.H., Williams, I.R. and Nijman, W., 2001. Archaean geology of the East Pilbara granite-greenstone terrane, Western Australia--a field guide. Western Aust. Geol. Surv., Record 2001/9, 134 pp. Van Kranendonk, M.J., Hickman, A.H., Smithies, R.H. and Nelson, D.R., 2002. Geology and tectonic evolution of the Archaean North Pilbara terrain, Pilbara craton, Western Australia. Econ. Geol., 97: 695-732. Van Kranendonk, M.J., Collins, W.J., Hickman, A.H. and Pawley, M.J., in press. Critical tests of vertical vs horizontal tectonic models for the Archaean East Pilbara granite-greenstone terrane, Pilbara craton, Western Australia. Precambrian Res. Van Reenen, D.D., Barton, J.M., Jr., Roering, C., Smit, C.A. and Van Schalkwyk, J.F., 1987. Deep crustal response to continental collision: the Limpopo belt of southern Africa. Geology, 15: 11-14. Van Reenen, D.D., Roering, C., Ashwal, L.D. and de Wit, M.J., 1992. Regional geological setting of the Limpopo belt. Precambrian Res., 55: 1-5. Van Schmus, W.R., Bickford, M.E., Anderson, J.L., Bender, E.E., Anderson, R.R., Bauer, P.W., Robertson, J.M., Bowring, S.A., Condie, K.C., Denison, R.E., Gilbert, M.C., Grambling, J.A., Mawer, C.K., Shearer, C.K., Hinze, W.J., Karlstrom, K.E., Kisvarsanyi, E.B., Lidiak, E.G., Reed, J.C., Sims, P.K., Tweto, O., Treves, S.B., Williams, M.L. and Wooden, J.L., 1993. Transcontinental Proterozoic provinces. In: J.C. Reed, M.E. Bickford, R.S. Houston, P.K. Link, D.W. Rankin, P.K. Sims and W.R. Van Schmus (Eds.), Precambrian: Conterminous U.S. Geol. Soc. Amer., Geol. North America C-2, pp. 171-334. Van Thienen, P., van den Berg, A.P., de Smet, J.H., van Hunen, J. and Drury, M.R., in press. The interaction between small scale mantle diapirs and a continental root. Geochem. Geophys. Geosyst. Van Wagoner, J.C., Mitchum, R.M., Jr., Campion, K.M. and Rahmanian, V.D., 1990. Siliciclastic Sequence Stratigraphy in Well Logs, Core, and Outcrops: Concepts for High-Resolution Correlation of Time and Facies. Methods in Exploration Series, 7. Amer. Assoc. Petrol. Geol., 55 pp. Van Wagoner, J.C., Posamentier, H.W., Mitchum, R.M., Vail, P.R., Sarg, J.E, Loutit, T.S. and Hardenbol, J., 1988. An overview of sequence stratigraphy and key definitions. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Eds.), Sea Level Changes--An Integrated Approach. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ., 42, pp. 39-45. Van Zuilen, M.A., Lepland, A. and Arrhenius, G., 2002. Reassessing the evidence for the earliest traces of life. Nature, 418: 627-630. Vearncombe, J.R., 1991. A possible Archaean island arc in the Murchison belt, Kaapvaal craton, South Africa. J. Afr. Earth Sci., 13: 299-304. Veevers, J.J., 1988. Gondwana facies started when Gondwanaland merged in Pangea. Geology, 16: 732-734. Veevers, J.J., 1989. Middle/Late Triassic (230 4- 5 Ma) singularity in the stratigraphic and magmatic history of the Pangean heat anomaly. Geology, 17: 784-787.

References

909

Veevers, J.J., 1990. Tectonic-climatic supercycle in the billion-year plate-tectonic eon: Permian Pangean icehouse alternates with Cretaceous dispersed-continents greenhouse. Sediment. Geol., 68:1-16. Veevers, J.J., 1995. Emergent, long-lived Gondwanaland vs. submergent, short-lived Laurasia: Supercontinental and Pan-African heat imparts long-term buoyancy by mafic underplating. Geology, 23:1131-1134. Veevers, J.J., Walter, M.R. and Scheibner, E., 1997. Neoproterozoic tectonics of Australia-Antarctica and Laurentia and the 560 Ma birth of the Pacific Ocean reflect the 400 m.y. Pangean supercycle. J. Geol., 105: 225-242. Veizer, J., 1983a. Geologic evolution of the Archean--early Proterozoic earth. In: J.W. Schopf (Ed.), Earth's Ealiest Biosphere: Its Origin and Evolution. Princeton Univ. Press, Princeton, New Jersey, pp. 240-259. Veizer, J., 1983b. Trace elements and isotopes in sedimentary carbonates. In: R.J. Reeder (Ed.), Carbonates: Mineralogy and Chemistry. Mineral. Soc. Amer., pp. 265-393. Veizer, J., 1988. The evolving exogenic cycle. In: C.B. Gregor, R.M. Garrels, ET. Mackenzie and J.B. Maynard (Eds.), Chemical Cycles in the Evolution of the Earth. John Wiley and Sons, New York, pp. 175-219. Veizer, J. and Jansen, S.L., 1985a. Basement and sedimentary recycling and continental evolution. J. Geol., 87: 341-370. Veizer, J. and Jansen, S.L., 1985b. Basement and sedimentary recyclingm2: time dimensions to global tectonics. J. Geol., 93: 625-643. Veizer, J., Hoefs, J., Ridler, R.H., Jensen, L.S. and Lowe, D.R., 1989. Geochemistry of Precambrian carbonates: I. Archean hydrothermal systems. Geochim. Cosmochim. Acta, 53: 845-857. Veizer, J., Clayton, R.N., Hinton, R.W., von Brunn, V., Mason, T.R., Buck. G. and Hoefs, J., 1990. Geochemistry of Precambrian carbonates: 3-shelf seas and non-marine environments of the Archaean. Geochim. Cosmochim. Acta, 54: 2717-2729. Veizer, J., Plumb, K.A., Clayton, R.N., Hinton, R.W. and Grotzinger, J.P., 1992. Geochemistry of Precambrian carbonates: V. Late Proterozoic seawater. Geochim. Cosmochim. Acta, 56: 2487-2501. Venkatachala, B.S., Shukla, M., Sharma, M., Naqvi, S.M., Srinivasan, R. and Udairaj, B., 1990. Archaean microbiota from the Donimalai Formation, Dharwar Supergroup, India. Precambrian Res., 47: 27-34. Vennemann, T.W. and Smith, H.S., 1999. Geochemistry of mafic and ultramafic rocks in the type section of the Kromberg Formation, Barberton Greenstone Belt, South Africa. In: D.R. Lowe and G.R. Byerly (Eds.), Geologic Evolution of the Barberton Greenstone Belt, South Africa. Geol. Soc. Amer., Spec. Paper, 329, pp. 133-149. Verni~res, J., Godard, M. and Bodinier, J.-L., 1997. A plate model for the simulation of trace element fractionation during partial melting and magma transport in the Earth's upper mantle. J. Geophys. Res., 102:24,771-24,784. Vetrin, V.P., 1998. The Belomorian megablock lower crust: its age, structure and conditions of formation (by result of deep seated xenolith studying). Proccedings of the Murmansk State Technical University, 1(3): 7-18 (in Russian). Vigneresse, J.L., Tikoff, B. and Am6glio, L., 1999. Modification of the regional stress field by magma intrusion and formation of tabular granitic plutons. Tectonophysics, 302: 203-244. Viljoen, M.J. and Viljoen, R.P., 1969a. The geology and geochemistry of the lower ultramafic unit of the Onverwacht Group and a proposed new class of igneous rocks. Geol. Soc. S. Afr., Johannesburg, Spec. Publ., 2, pp. 55-86.

910

References

Viljoen, M.J. and Viljoen, R.E, 1969b. Evidence for the existence of a mobile extrusive peridotitic magma from the Komati Formation of the Onverwacht Group. Geol. Soc. S. Afr., Johannesburg, Spec. Publ., 2, pp. 87-112. Viljoen, R.P. and Viljoen, M.J., 1969c. The geological and geochemical significance of the upper formations of the Onverwacht Group. Geol. Soc. S. Afr. Spec. Publ., 2, pp. 113-151. Villeneuve, M.E., Ross, G.M., Theriault, R.J., Miles, W., Parrish, R.R. and Broome, J., 1993. Tectonic subdivision and U-Pb geochronology of the crystalline basement of the Alberta Basin, Western Canada. Geol. Surv. Can. Bull., 447: 1-86. Vishnevsky, S.A., 1975. Some opaque minerals in impact glasses from the Popigai structure: methodology and methods of geological and geophysical investigations in Siberia. Inst. Geol. Geophys., Novosibirsk, pp. 105-115 (in Russian). Visknevsky, S.A., Kovaleva, L.T. and Pal'chik, N.A., 1975. Stishovite in rocks of the Popigai structure (in Russian). Dokl. Akad. Nauk SSSR, 221 : 1167-1169. Visser, J.N.J., 1981. The Mid-Precambrian tillite in the Griqualand West and Transvaal Basins, South Africa. In: M.J. Hambrey and W.B. Harland (Eds.), Earth's Pre-Pleistocene Glacial Record. Cambridge Univ. Press, Cambridge, pp. 180-184. Visser, M.J., 1980. Neap-spring cycles reflected in Holocene large-scale bedform deposits: a preliminary note. Geology, 8: 543-546. Vissher, P.T., Reid, P.R. and Bebout, B.M., 2000. Microscale observations of sulfate reduction: correlation of microbial activity with lithified micritic laminae in modern marine stromatolites. Geology, 28: 919-922. Vlaar, N.J., Van Keken, P.E. and Van den Berg, A.P., 1994. Cooling of the Earth in the Archaean: consequences of pressure-release melting in a hotter mantle. Earth Planet. Sci. Lett., 121:1-18. Vogel, D.C., Vuollo, J.l., Alapieti, T.T. and James, R.S., 1998. Tectonic, stratigraphic and geochemical comparisons between ca. 2500-2440 Ma mafic igneous events in the Canadian and Fennoscandian Shields. Precambrian Res., 92:89-116. Von Brunn, V. and Gold, D.J.D., 1993. Diamictite in the Archean Pongola sequence of southern Africa. J. Afr. Earth Sci., 16: 367-374. Vos, R.G. and Eriksson, K.A., 1977. An embayment model for tidal and wave swash deposits occurring within a fluvially dominated middle Proterozoic sequence in South Africa. Sediment. Geol., 18: 161-173. Vuollo, J., 1994. Palaeoproterozoic basic igneous events in Eastern Fennoscandian Shield between 2.45 and 1.97 Ga. Acta Univ. Oulu, A250, 32 pp. W~ichtersh~iuser, G., 1988. An all-purine precursor of nucleic acids. Proc. Natl. Acad. Sci. U.S.A., 85:1134-1135. W~ichtersh~iuser, G., 1990. Evolution of the first metabolic cycles. Proc. Natl. Acad. Sci. U.S.A., 87: 200-204. W~ichtersh~iuser, G., 1994. Vitalysts and virulysts: a theory of self-expanding reproduction. In: S. Bengtson (Ed.), Early life on Earth. Nobel Symposium No. 84. Columbia Univ. Press, New York, pp. 124-132. Wadge, G., Archer, D.J. and Millington, A.C., 1994. Monitoring playa sedimentation using sequential radar images. Terra Nova, 6:391-396. Wadsworth, M.E., 1880. Notes on the Geology of the Iron and Copper Districts of Lake Superior. Bull. Museum Comparative Zoology, Harvard, 7: 1-157. Wahlstrom, E.E., 1948. Pre-Fountain and recent weathering on Flagstaff Mountain near Boulder, Colorado. Geol. Soc. Amer. Bull., 59:1173-1190.

References

911

Waitt, R.B.J., 1985. Case for periodic, colossal joekulhlaups from Pleistocene glacial Lake Missoula. Geol. Soc. Amer. Bull., 96: 1271-1286. Walcott, C.D., 1914. Cambrian geology and paleontology III, no. 2, Pre-Cambrian Algonkian algal flora. Smithsonian Misc. Collections, 64(2): 77-156. Walde, D.H.G., Gierth, E. and Leonardos, O.H., 1981. Stratigraphy and mineralogy of the manganese ores of Urucum, Matto Grosso, Brazil. Geol. Rundschau, 70: 1077-1085. Walkden, G., Parker, J. and Kelley, S., 2002. A late Triassic impact ejecta layer in southwestern Britain. Science, 298:2185-2188. Walker, J.C.G., 1977. Evolution of the Atmosphere. Macmillian, New York, 318 pp. Walker, J.C.G. and Zahnle, K.J., 1986. Lunar nodal tide and distance to the Moon during the Precambrian. Nature, 320: 600-602. Walker, J.C.G., Hays, P.B. and Kasting, J.E, 1981. A negative feedback mechanism for the long-term stabilization of Earth's surface temperature. J. Geophys. Res., 86: 9776-9782. Walker, R.G. and Pettijohn, EJ., 1971. Archaean sedimentation: analysis of the Minnitaki basin, northwestern Ontario, Canada. Geol. Soc. Amer. Bull., 82:2099-2130. Walker, R.G. and Plint, A.G., 1992. Wave- and storm-dominated shallow marine systems. In: R.G. Walker and N.P. James (Eds.), Facies Models, Response to Sea Level Change. Geol. Assoc. Can., Toronto, pp. 219-238. Walraven, E and Martini, J., 1995. Zircon Pb-evaporation age determinations of the Oak Tree Formation, Chuniespoort Group, Transvaal Sequence: implications for Transvaal-Griqualand West basin correlations. S. Afr. J. Geol., 98: 58-67. Walsh, M.M., 1992. Microfossils and possible microfossils from the Early Archean Onverwacht Group, Barberton Mountain Land, South Africa. Precambrian Res., 54: 271-293. Walsh, M.M. and Lowe, D.R., 1985. Filamentous microfossils from the 3,500-Myr-old Onverwacht Group, Barberton Mountain Land, South Africa. Nature, 314: 530-532. Walter, M.J., 1998. Melting of garnet peridotite and the origin of komatiite and depleted lithosphere. J. Petrol., 39: 29-60. Walter, M.R., 1972. Stromatolites and the biostratigraphy of the Australian Precambrian and Cambrian. Paleont. Assoc. London, Spec. Paper, 11,190 pp. Walter, M.R., 1976. Introduction. In: M.R. Walter (Ed.), Stromatolites. Developments in Sedimentology, Vol. 20, Elsevier, Amsterdam, pp. 1-3. Walter, M.R., 1983. Archean stromatolites: evidence of the Earth's earliest benthos. In: J.W. Schopf (Ed.), Earth's Earliest Biosphere, Its Origin and Evolution. Princeton Univ. Press, Princeton, New Jersey, pp. 187-213. Walter, M.R., 1994. Stromatolites: the main information on the evolution of early benthos. In: S. Bengston (Ed.), Early Life on Earth. Nobel Symposium, 84. Columbia Univ. Press, New York, pp. 270-286. Walter, M.R. and Bauld, J., 1983. The association of sulphate evaporites, stromatolitic carbonates and glacial sediments: examples from the Proterozoic of Australia and the Cainozoic of Antarctica. Precambrian Res., 21: 129-148. Walter, M.R., Buick, R. and Dunlop, J.S.R. 1980. Stromatolites, 3.400-3.500 Myr old from the North Pole area, Western Australia. Nature, 284: 443-445. Walter, M.R., Rulin, D. and Horodyski, R.J., 1990. Coiled carbonaceous megafossils from the Middle Proterozoic of Jixian (Tianjin) and Montana. Amer. J. Sci., 290-A: 133-148. Walter, M.R., Grotzinger, J.P. and Schopf, J.W., 1992. Proterozoic stromatolites. In: J.W. Schopf and C. Klein (Eds.), The Proterozoic Biosphere. Cambridge Univ. Press, New York, pp. 253-260.

912

References

Walter, M.R., Veevers, J.J., Calver, C.R. and Grey, K., 1995. Neoproterozoic stratigraphy of the Centralian Superbasin, Australia. Precambrian Res., 73: 173-195. Walter, M.R., Veevers, J.J., Calver, C.R., Gorjan, P. and Hill, A.C., 2000. Dating the 840-544 Ma Neoproterozoic interval by isotopes of strontium, carbon, and sulfur in seawater, and some interpretative models. Precambrian Res., 100: 371-433. Wannamaker, P.E., Hulen, J.B. and Heizler, M.T., 2000. Early Miocene lamproite from the Colorado Plateau tectonic province, southeastern Utah, USA. J. Volcanol. Geotherm. Res., 96: 175-190. Ward, P.D. and Brownlee, D., 2000. Rare Earth. Copernicus, Springer-Verlag, New York, 333 pp. Warme, J.E., Morgan, M. and Kuehner, H.-C., 2002. Impact-generated carbonate accretionary lapilli in the Late Devonian Alamo Breccia. In: C. Koeberl and K.G. McLeod (Eds.), Catastrophic Events and Mass Extinctions: Impacts and Beyond. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 356, pp. 489-504. Washburn, A.L., 1980. Geocryology. A Survey of Periglacial Processes and Environments. John Wiley and Sons, New York, 406 pp. Wasserburg, G.J. and Arnould, M., 1987. A possible relationship between extinct 26A1 and 53Mn in meteorites and early solar activity. In: W. Hillebrandt, R. Kuhfuss, E. MOiler and J.W. Turan (Eds.), Nuclear Astrophysics. Springer-Verlag, Berlin, pp. 262-276. Wasserburg, G.J., Busso, M. and Gallino, R., 1996. Abundances of actinides and short-lived nonactinides in the interstellar medium: diverse supernova sources for the r-processes. Astrophys. J., 466:L109-113. Wasson, R.J. and Hyde, R., 1983. Factors determining desert dune type. Nature, 304: 337-339. Watanabe, Y., Condie, K.C., Ohmoto, H., Naraoka, H. and Wronkiewicz, D.J., 1997. Carbon, nitrogen, and sulfur geochemistry of Archean and Proterozoic shales from the Kaapvaal Craton, South Africa. Geochim. Cosmochim. Acta, 61:3441-3459. Watanabe, Y., Martini, J.E.J. and Ohmoto, H., 2000. Geochemical evidence for terrestrial ecosystems 2.6 billion years ago. Nature, 408: 574-578. Waters, D.J. 1989., Metamorphic evidence for the heating and cooling of Namaqualand granulites. In: J.S. Daly, R.A. Cliff and B.W.D. Yardley (Eds.), Evolution of Metamorphic Belts. Geol. Soc. London Spec. Publ., 43, pp. 357-363. Watkeys, M., 1984. The Precambrian geology of the Limpopo Belt north and west of Messina. Ph.D. Thesis (unpubl.), Univ. Witwatersrand, Johannesburg, South Africa, 349 pp. Watson, E.B. and Cherniak, D.J., 1997. Oxygen diffusion in zircon. Earth Planet. Sci. Lett., 148: 527-544. Wayne, W.J., 1990. Periglacial sand-wedge polygons in eastern Nebraska. Geol. Soc. Amer. Abs. Progr., 22(7): A85. Webb, D.J., 1982. On the reduction in tidal dissipation produced by increases in the Earth's rotation rate and its effect on the long-term history of the Moon's orbit. In: P. Brosche and J. Stindermann (Eds.), Tidal Friction and the Earth's Rotation II. Springer-Verlag, Berlin, pp. 210-221. Wedepohl, K.H., 1969. Composition and abundance of common igneous rocks. In: K.H. Wedepohl (Ed.), The Handbook of Geochemistry, Vol. 1. Springer-Verlag, Berlin, pp. 247-248. Wei, J., 2002. The late Palaeoproterozoic orogeny in the North China Craton. Gondwana Res., 5(1): 95-99. Weichert, U., Halliday, A.N., Lee, D.-C., Snyder, G.A., Taylor, L.A. and Rumble, D., 2001. Oxygen isotopes and the Moon-forming giant impact. Science, 294: 345-348. Weidenschilling, S.J., 1977. Aerodynamics of solid bodies in the solar nebula. Monthly Notices Roy. Astronom. Soc., 180: 57-70.

References

913

Weinberg, R.E and Podladchikov, Y.Y., 1995. The rise of solid-state diapirs. J. Struct. Geol., 17: 1183-1195. Weinberg, R.F. and Sandiford, M., 2001. Diapirism in the Western Australian Archaean. In: K.E Cassidy, J.M. Dunphy and M.J. Van Kranendonk (Eds.), Extended Abstract Volume, 4th International Archaean Symposium, Perth, 2001, Aust. Geol. Surv. Org., Geoscience Australia, Record 2001/37, p. 371. Weinberg, R.E and Schmelling, H., 1992. Polydiapirs: multiwavelength gravity structures. J. Struct. Geol., 14: 425-436. Weiss, G.C. and Middleton, L.T., 1986. Depositional analyses of the arkose member of the Troy Quartzite (younger Precambrian) in central Arizona: an ancient eolian-fluvial system. Geol. Soc. Amer. Abstr. Progr., 18, p. 421. Wellman, P., 2000. Upper crust of the Pilbara Craton, Australia: 3D geometry of a granite-greenstone terrain. Precambrian Res., 104:187-206. Wells, G., Bryan, W.B. and Pearce, T.H., 1979. Comparative morphology of ancient and modem pillow lavas. J. Geology, 87: 427--440. West, G.E and Mareschal, J.-C., 1979. A model for Archaean tectonism. Part 1. The thermal conditions. Can. J. Earth Sci., 16: 1942-1950. Westall, E, 1999. The nature of fossil bacteria. J. Geophys. Res., 104: 16,437-16,451. Westall, E, 2002a. Le contexte grologique de l'origine de la vie et les signatures min6rales de la vie fossile. In: M. Gargaud (Ed.), Exobiologie, les traces du vivant et l'origine de la vie. Presses Universitaires de Bordeaux (in press). Westall, E, 2002b. Why potential life on Mars was likely to have been spatially limited as well as heterogeneous. Abstract, Lunar and Planetary Sciences Conference 33, Houston, Texas, 2002, abstract no. 1071. Westall, E, de Wit, M.J., Dann, J., van der Gaast, S., de Ronde, C. and Gerneke, D., 2001a. Early Archaean fossil bacteria and biofilms in hydrothermally-influenced, shallow water sediments, Barberton greenstone belt, South Africa. Precambrian Res., 106:93-116. Westall, E and Folk, R.L., 2003. Endogenous and exogenous carbonaceous microstructures in Early Archaean cherts and B IFs from the Isua greenstone belt: implications for the search for life in ancient rocks. Precambrian Res., 126:313-330. Westall, F., Nijman, W., Brack, A., Steele, A. and Toporski, J., 200lb. The oldest fossil life on Earth, its geological context and life on Mars. Proceedings, First European Workshop on Exo/Astrobiology, Frascati. ESA Spec. Publ., 496: 81-90. Westall, E, Walsh, M., De Vries, S. and Nijman, W., 2001c. Fossil microbial biofilms and their sedimentological context in the Early Archaean cherts of the Barberton and Pilbara greenstone belts. In: 4th International Archaean Symposium, Perth, Australia, Sept. 2001, pp. 266-268. Wetherill, G.W., 1994. Provenance of the terrestrial planets. Geochim. Cosmochim. Acta, 58: 4513-4520. Wetherill, G.W. and Stewart, G.R., 1993. Formation of planetary embryos: effects of fragmentation, low relative velocity, and independent variation of eccentricity and inclination. Icarus, 106: 190-209. Whalen, J.B., Syme, E.C. and Stern, R.A., 1999. Geochemical and Nd isotopic evolution of Paleoproterozoic arc-type granitoid magmatism in the Flin-Flon belt, Trans-Hudson Orogen, Canada. Can. J. Earth Sci., 36: 227-250. Wheeler, H.E., 1958. Time stratigraphy. Amer. Assoc. Petrol. Geol. Bull., 42: 1047-1063. Wheeler, H.E., 1959. Unconformity bounded units in stratigraphy. Amer. Assoc. Petrol. Geol. Bull., 43: 1975-1977.

914

References

Wheeler, H.E., 1964. Baselevel, lithosphere surface, and time-stratigraphy. Geol. Soc. Amer. Bull., 75: 599-610. Wheeler, H.E. and Murray, H.H., 1957. Baselevel control patterns in cyclothemic sedimentation. Amer. Assoc. Petrol. Geol. Bull., 41:1985-2011. Whitby, J.A., Burgess, R., Turner, G., Gilmour, J. and Bridges, J., 2000. Extinct 1291 in halite from a primitive meteorite: evidence for evaporite formation in the early solar system. Science, 288: 1819-1821. White, D.J. and Lucas, S.B., 1994. A closer look at the Superior boundary zone. In: Z. Hajnal and J. Lewry (Eds.), Trans-Hudson Orogen transect. Report, Fourth Transect Workshop, Lithoprobe Report 38, pp. 35-41. White, J.D.L., 1994. Review of: McPhie J., Doyle M., Allen R.L. (1993) Volcanic textures: a guide to the interpretation of textures in volcanic rocks. Bull. Volcanol., 56:412-413. White, J.D.L., 1996. Pre-emergent construction of a lacustrine basaltic volcano, Pahvant Butte, Utah (USA). Bull. Volcanol., 58: 249-262. White, J.D.L., 2000. Subaqueous eruption-fed density currents and their deposits. Precambrian Res., 101: 87-109. Whitehead, J., Papanastassiou, D.A., Spray, J.G., Grieve, R.A.F. and Wasserburg, G.J., 2000. Late Eocene impact ejecta; geochemical and isotopic connections with the Popigai impact structure. Earth Planet. Sci. Lett., 181: 473-487. Whitham, A.G. and Sparks, R.S.J., 1986. Pumice. Bull. Volcanol., 48: 209-223. Whitten, G.F., 1970. The investigation and exploitation of the Razorback Ridge iron deposit. S. Aust. Geol. Surv., Report of Investigations 33, 165 pp. Widdell, F., Schnell, S., Heising, S., Ehrenreich, A., Assmuss, B. and Schink, B., 1993. Ferrous iron oxidation by anoxygenic phototrophic bacteria. Nature, 362: 834-836. Wiesner, M.G., Wang, Y. and Zheng, L., 1995. Fallout of volcanic ash to the deep South China Sea induced by the 1991 eruption of Mount Pinatubo (PhiUipines). Geology, 23: 885-888. Wiggering, H. and Beukes, N.J., 1990. Petrography and geochemistry of a 2000-2200 Ma-old hematitic paleo-alteration on Ongeluk basalt of the Transvaal Supergroup, Griqualand West, South Africa. Precambrian Res., 46:241-258. Wilde, S.A., Valley, J.W., Peck, W.H. and Graham, C.M., 2001. Evidence from detrital zircons for the existence of continental crust and oceans on the Earth 4.4 Gyr ago. Nature, 409:175-178. Wilgus, C.K., Hastings, B.S., Kendall, C.G.St.C., Posamentier, H.W., Ross, C.A. and Van Wagoner, J.C., 1988. Sea Level ChangesIAn Integrated Approach. Soc. Econ. Paleontol. Mineral., Tulsa, Oklahoma, Spec. Publ. 42, 407 pp. Wilkins, C., 1997. Regional and contact metamorphism. In: M.J. de Wit and L.D. Ashwal. (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 126-163. Wilks, M.E. and Nisbet, E.G., 1988. Stratigraphy of the Steep Rock Group, northwest Ontario: a major Archaean unconformity and Archaean stromatolites. Can. J. Earth Sci., 25:370-391. Williams, B.A.P., Hirt, R.P., Lucocq, J.M. and Embley, T.M., 2002. A mitrochondrial remnant in the microsporidian Trachipleistophoroa hominis. Nature, 418: 865-869. Williams, D.A., Kerr, R.C. and Lesher, C.M., 1998. Emplacement and erosion by Archean komatiite lava flows at Kambalda: revisited. J. Geophys. Res., 103: 27,533-27,549. Williams, D.A.C. and Furnell, R.G., 1979. Reassessment of part of the Barberton type area, South Africa. Precambrian Res., 9: 325-347. Williams, D.M., Kasting, J.F. and Frakes, L.A., 1998. Low-latitude glaciation and rapid changes in the Earth's obliquity explained by obliquity-oblateness feedback. Nature, 396: 453-455.

References

915

Williams, G.E., 1972. Geological evidence relating to the origin and secular rotation of the solar system. Modem Geol., 3: 165-181. Williams, G.E., 1975. Late Precambrian glacial climate and the Earth's obliquity. Geol. Mag., 112: 441-465. Williams, G.E., 1979. Sedimentology, stable-isotope geochemistry and palaeoenvironment of dolostones capping late Precambrian glacial sequences in Australia. J. Geol. Soc. Aust., 26: 377-386. Williams, G.E., 1986. Precambrian permafrost horizons as indicators of palaeoclimate. Precambrian Res., 32: 233-242. Williams, G.E., 1989a. Late Precambrian tidal rhythmites in South Australia and the history of the Earth's rotation. J. Geol. Soc., London, 146:97-111. Williams, G.E., 1989b. Precambrian tidal sedimentary cycles and Earth's paleorotation. Eos, Trans. Amer. Geophys. Union, 70: 33; 40--41. Williams, G.E., 1989c. Tidal rhythmites: geochronometers for the ancient Earth-Moon system. Episodes, 12:162-171. Williams, G.E., 1990. Tidal rhythmites: key to the history of the Earth's rotation and the lunar orbit. J. Phys. Earth, 38: 475-491. Williams, G.E., 1991. Upper Proterozoic tidal rhythmites, South Australia: sedimentary features, deposition, and implications for the Earth's paleorotation. Can. Soc. Petrol. Geol. Mem., 16: 161-177. Williams, G.E., 1993. History of the Earth's obliquity. Earth Sci. Rev., 34: 1-45. Williams, G.E., 1994a. The enigmatic late Proterozoic glacial climate: an Australian perspective. In: M. Deynoux, J.M.G. Miller, E.W. Donack, N. Eyles, I.J. Fairchild and G.M. Young (Eds.), Earth's Glacial Record. Cambridge Univ. Press, Cambridge, pp. 146-164. Williams, G.E., 1994b. Resonances of the fluid core for a tidally decelerating Earth: cause of increased plume activity and tectonothermal reworking events? Earth Planet. Sci. Lett., 128: 155-167. Williams, G.E., 1994c. History of the Earth's rotation and the moon's orbit: a key datum from Precambrian strata in South Australia. Aust. J. Astron., 5: 135-147. Williams, G.E., 1996. Soft-sediment deformation structures from the Marinoan glacial succession, Adelaide foldbelt: implications for the palaeolatitude of late Neoproterozoic glaciation. Sediment. Geol., 106: 165-175. Williams, G.E., 1997. Precambrian length of day and the validity of tidal rhythmite paleotidal values. Geophys. Res. Lett., 24: 421-424. Williams, G.E., 1998a. Late Neoproterozoic periglacial aeolian sand sheet, Stuart Shelf, South Australia. Aust. J. Earth Sci., 45: 733-741. Williams, G.E., 1998b. Precambrian tidal and glacial clastic deposits: implications for Precambrian Earth-Moon dynamics and palaeoclimate. Sediment. Geol., 120: 55-74. Williams, G.E., 2000. Geological constraints on the Precambrian history of Earth's rotation and the Moon's orbit. Rev. Geophys., 38: 37-59. Williams, G.E., 2001. Neoproterozoic (Torridonian) alluvial fan succession, northwest Scotland, and its tectonic setting and provenance. Geol. Mag., 138/2: 161-184. Williams, G.E. and Schmidt, P.W., 1997. Paleomagnetism of the paleoproterozoic Gowganda and Lorrain Formations, Ontario: low paleolatitude for Huronian glaciation. Earth Planet. Sci. Lett., 153: 157-169. Williams, G.E. and Schmidt, P.W., 2000. Proterozoic equatorial glaciation: has 'snowball Earth' a snowball's chance? Aust. Geologist, 117: 21-25.

916

References

Williams, G.E. and Tonkin, D.G., 1985. Periglacial structures and palaeoclimatic significance of a late Precambrian block field in the Cattle Grid copper mine, Mount Gunson, South Australia. Aust. J. Earth Sci., 32: 297-300. Williams, H.R., Stott, G.M. and Thurston, P.C., 1992. Tectonic evolution of Ontario: summary and synthesis. Part 1: Revolution in the Superior Province. Ontario Geol. Surv. Spec. Vol. (Geol. Ontario), 4, pp. 1255-1294. Williams, I.R., 1999. Geology of the Muccan 1:100000 Sheet. Western Aust. Geol. Survey, 1 : 100000 Geol. Ser. Explanatory Notes, 39 pp. Williams, I.S. and Collins, W.J., 1990. Granite-greenstone terranes in the Pilbara Block, Australia, as coeval volcano-plutonic complexes; evidence from U-Pb zircon dating of the Mount Edgar Batholith. Earth Planet. Sci. Lett., 97: 41-53. Williams, L.A. and Reimers, C., 1983. Role of bacterial mats in oxygen-deficient marine basins and coastal upwelling regimes: preliminary report. Geology, 11: 267-269. Williams, S.N., 1983a. Geology and eruptive mechanics of Maya caldera complex, Nicaragua. Ph.D. Thesis (unpubl.), Dartmouth College, Hanover, New Hampshire, 169 pp. Williams, S.N., 1983b. Plinian airfall deposits of basaltic composition. Geology, 11: 211-214. Willner, A.P., 1993. Grundlagen zur Dynamik der mittleren Erdkruste: PT-Pfad, Fluidentwicklung und Fluorverteilung in einer Modellregion im Namaqua Mobile Belt (Siidafrika). Ferdinand Enke Verlag, Stuttgart, 191 pp. Wilson, C.J.N., 1993. Stratigraphy, chronology, styles, and dynamics of Late Quaternary eruptions from Taupo volcano, New Zealand. Philos. Trans. Roy. Soc. London Ser. A, 314: 205-306. Wilson, C.J.N., Houghton, B.E, McWilliams, M.O., Lanphere, M.A., Weaver, S.D. and Briggs, R.M., 1995. Volcanic and structural evolution of the Taupo Volcanic Zone, New Zealand: a review. J. Volcanol. Geotherm. Res., 68: 1-28. Wilson, J.A., 1985. Geology of the Athabasca Group in Alberta. Alberta Research Council, Geol. Surv. Dep. Bull., 49: 1-78. Wilson, J.E, Nesbitt, R.W. and Fanning, C.M., 1995. Zircon geochronology of Archaean felsic sequences in the Zimbabwe cratonma revision of greenstone stratigraphy and a model for crustal growth. In: M.P. Coward and A.C. Reis (Eds.), Early Archaean Processes. Geol. Soc. London Spec. Publ., 95, pp. 109-126. Wilson, N.S., Stasiuk, L.D., Fowler, M.G., Jiricka, D., Sopuck, V., Zaluski, G., Morrison, D. and Wheatley, K., 2002. Organic petrology and geochemistry of organic matter within Proterozoic Athabasca Group shales and sandstones distal to U deposits, northern Saskatchewan and Alberta. Abstract, Geol. Assoc. Can., Mineral. Assoc. Can., Joint Annual Meeting, Vol. 27, p. 124. Winchell, N.H., 1900. The geology of Minnesota, structural and petrographic geology of the Taconic and Archaean. Minnesota Geology and Natural History Survey, 5: 1-80. Windley, B., 1992. Proterozoic collisional and accretionary orogens. In: K.C. Condie (Ed.), Proterozoic Crustal Evolution. Elsevier, Amsterdam, pp. 419--446. Windley, B.E, 1995. The Evolving Continents, 3rd edn. John Wiley and Sons, Chichester, 526 pp. Windley, B.E and Davies, EB., 1978. Volcano spacings and lithospheric/crustal thickness in the Archaean. Earth Planet. Sci. Lett., 38:291-297. Windley, B.M., 1993. Uniformitarianism today: plate tectonics is the key to the past. J. Geol. Soc. London, 150:7-19. Wingate, M.T.D., 1998. A palaeomagnetic test of the Kaapvaal-Pilbara (Vaalbara) connection at 2.78 Ga. S. Afr. J. Geol., 101: 257-274.

References

917

Wingate, M.T.D. and Giddings, J.W., 2000. Age and palaeomagnetism of the Mundine Well dyke swarm, Western Australia: implications for an Australia-Laurentia connection at 750 Ma. Precambrian Res., 100: 335-357. Wingate, M.T.D., Campbell, I.H., Compston, W. and Gibson, G.M., 1998. Ion microprobe U-Pb ages for Neoproterozoic basaltic magmatism in south-central Australia and implications for the breakup of Rodinia. Precambrian Res., 87: 135-159. Winston, D., 1986. Sedimentology of the Ravalli Group, Middle Belt Carbonate and Missoula Group, Middle Proterozoic Belt Supergroup, Montana, Idaho, and Washington. Montana Bureau Min. Geol. Spec. Publ., 94, pp. 85-124. Winston, D., 1990. Evidence for intracratonic, fluvial, and lacustrine settings of Middle to Late Proterozoic basins of western U.S.A. In: C.E Gower, T. Rivers and B. Ryan (Eds.), Mid-Proterozoic Laurentia-Baltica. Geol. Assoc. Canada, Spec. Paper, 38, pp. 535-564. Winston, D. and Link, P.K., 1993. Middle Proterozoic rocks of Montana, Idaho and eastern Washington. In: J.C. Reed, Jr., M.E. Bickford, R.S. Houston, P.K. Link, D.W. Rankin, P.K. Sims and R.E. Van Schmus (Eds.), Precambrian: Conterminous U.S., The Geology of North America, v. C-2. Geol. Soc. Amer., Boulder, Colorado, pp. 487-517. Wise, D.U., 1972. Freeboard of continents through time. Geol. Soc. Amer. Mem., 132: 87-100. Wise, D.U., 1974. Continental margins, freeboard and volumes of continents and oceans through time. In: C.A. Burke and C.L. Drake (Eds.), The Geology of Continental Margins. SpringerVerlag, Berlin, pp. 45-58. Wizevich, M.C., 1997. Fluvial-eolian deposits in the Devonian New Mountain Sandstone, Table Mountain, Southern Victoria Land, Antarctica: sedimentary architecture, genesis and stratigraphic evolution. In: The Antarctic Region: Geological Evolution and Processes, 1997, pp. 933-944. Wohletz, K.H., 2003. Water/magma interaction: physical considerations for the deep submarine environment. In: J. White, D. Clague and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph, (in press). Wolde, B., Asres, Z., Desta, Z. and Gonzalez, J.J., 1996a. Neoproterozoic zirconium-depleted boninite and tholeiitic series rocks from Adola, southern Ethiopia. Precambrian Res., 80: 261-279. Wolde, B. and Gore-Gambella Geotraverse Team, 1996b. Tonalite-trondhjemite-granite genesis by partial melting of newly underplated basaltic crust: an example from the Neoproterozoic Birbir magmatic arc, western Ethiopia. Precambrian Res., 76: 3-14. Wolf, M.B. and Wyllie, P.J., 1993. Garnet growth during amphibolite anatexis: implications of a garnetiferous restite. J. Geol., 101: 357-373. Wolmarans, L.G. and Kent, L.E., 1982. Geological investigations in western Dronning Maud Land, Antarctica--a synthesis. S. Afr. J. Antarctic Res., Supplement 2: 1-93. Wood, J., 1980. Epiclastic sedimentation and stratigraphy in the North Spirit Lake and Rainy Lake areas: a comparison. Precambrian Res., 12: 227-255. Woodhead, J.D., Hergt, J.M. and Simonson, B.M., 1998. Isotopic dating of an Archaean bolide impact horizon, Hamersley Basin, Western Australia. Geology, 26: 47-50. Wopenka, B. and Pasteris, J.D., 1993. Structural characterization of kerogens to granulite-facies graphite: applicability of Raman microprobe spectroscopy. Amer. Mineral., 78: 533-557. Worm, H.U. and Banerjee, S.K., 1987. Rock magnetic signature of the Cretaceous-Tertiary boundary. Geophys. Res. Lett., 14: 1083-1086. Worsley, T.R., Nance, D. and Moody, J.B., 1984. Global tectonics and eustasy for the past 2 billion years. Marine Geol., 58: 373-400.

918

References

Worsley, T.R., Nance, R.D. and Moody, J.B., 1986. Tectonic cycles and the history of the Earth's biogeochemical and paleoceanographic record. Paleoceanogr., 1: 233-263. Wortmann, U.G., Bernasconi, S.M. and B6ttcher, M.E., 2001. Hypersulfidic deep biosphere indicates extreme sulfur isotope fractionation during single-step microbial sulfate reduction. Geology, 29: 647-650. Wright, D.T. and Altermann, W., 2000. Microfacies development in Late Archaean stromatolites and oolites of the Campbellrand Subgroup, South Africa. In: E. lnsalco, P.W. Skelton and T.J. Palmer (Eds.), Carbonate Platform Systems. Components and Interactions. Geol. Soc. London Spec. Publ., 178, pp. 51-70. Wright, I.C. and Gamble, J.A., 1999. Southern Kermadec submarine caldera arc volcanoes (SW Pacific); caldera formation by effusive and pyroclastic eruption. Mar. Geol., 161: 209-229. Wright, I.C., de Ronde, C.E.J., Faure, K. and Gamble, J.A., 1998. Discovery of hydrothermal sulfide mineralization from southern Kermadec arc volcanoes (SW Pacific). Earth Planet. Sci. Lett., 164: 335-343. Wright, V.P. and Marriott, S.B., 1993. The sequence stratigraphy of fluvial depositional systems: the role of floodplain sediment storage. Sediment. Geol., 86:203-210. Wronkiewicz, D.J. and Condie, K.C., 1990. Geochemistry and mineralogy of sediments from the Ventersdorp and Transvaal Supergroups, South Africa: craton evolution during the early Proterozoic. Geochim. Cosmochim. Acta, 54: 343-354. Wrucke, C.T., 1989. The Middle Proterozoic Apache Group, Troy Quartzite, and associated diabase of Arizona. In: J.P. Jenny and S.J. Reynolds (Eds.), Geologic Evolution of Arizona. Arizona. Geol. Soc. Digest, 17: 239-258. Wunsch, C., 1972. Bermuda sea level in relation to tides, weather, and baroclinic fluctuations. Rev. Geophys. Space Phys., 10: 1-49. Wuttke, M. and Radtke, G., 1993. Agglutinierende Mikrobenmatten im Profundal des mitteleozaenen Eckelfelder Maar-Sees bei Manderscheid/Eifel (Bundesrepublik Deutschland); Fossilfundstaette "Eckelfelder Maar" (Mittel-Eozaen)--Nr. 23. Mainzer Naturwissenschaftliches Archiv, 31:115-126. Wyborn, L.A.I., 1988. Petrology, geochemistry and origin of major Australian 1880-1840 Ma felsic volcanic suite: a model for intracontinental felsic magma generation. Precambrian Res., 40/41: 37-60. Wyche, S., Nelson, D.R. and Riganti, A., in press. The significance of 4.35-3.13 Ga detrital zircons in quartzite from the central Yilgarn craton, Western Australia. Aust. J. Earth Sci. Wyllie, P.J., Wolf, M.B. and van der Laan, S.R., 1997. Conditions for formation of tonalites and trondhjemites: magmatic sources and products. In: M.J. de Wit and L.D. Ashwal (Eds.), Greenstone Belts. Clarendon Press, Oxford, pp. 256-266. Wyman, D.A., 1999a. A 2.7 Ga depleted tholeiitic suite: evidence of plume-arc interaction in the Abitibi greenstone belt, Canada. Precambrian Res., 97: 27-42. Wyman, D.A., 1999b. Paleoproterozoic boninites in an ophiolite-like setting, Trans-Hudson orogen, Canada. Geology, 27: 455--458. Wyman, D.A., Kerrich, R. and Groves, D.I., 1999a. Lode gold deposits and Archean mantle-island arc interaction, Abitibi Subprovince, Canada. J. Geol., 107: 715-725. Wyman, D.A., Bleeker, W. and Kerrich, R., 1999b. A 2.7 Ga komatiite, low Ti tholeiite, arc tholeiite transition, and inferred proto-arc geodynamic setting of the Kidd Creek deposit: evidence from precise trace element data. In: M.D. Hannington and C.T. Barrie (Eds.), The Giant Kidd Creek Volcanogenic Massive Sulfide Deposit, Western Abitibi Subprovince, Canada. Econ. Geol., Monograph, 10, pp. 511-528.

References

919

Wyman, D.A., Kerrich, R. and Polat, A., 2002a. Assembly of Archean cratonic mantle lithosphere and crust: plume-arc interaction in the Abitibi-Wawa subduction-accretion complex. Precambrian Res., 115: 37-62. Wyman, D.A., Ayer, J.A. and Devaney J.R., 2002b. Niobium-enriched basalts fom the Wabigoon subprovince, Canada: evidence for adakitic metasomatism above an Archean subduction zone. Earth Planet. Sci. Lett., 179:21-30. Xiao, S., Zhang, Y. and Knoll, A.H., 1998. Three-dimensional preservation of algae and animal embryos in a Neoproterozoic phosphorite. Nature, 395: 553-558. Xie, Q., Kerrich, R. and Fan, J., 1993. HFSE/REE fractionations recorded in three komatiite-basalt sequences, Archean Abitibi greenstone belt: implications for multiple plume sources and depths. Geochim. Cosmochim. Acta, 57:4111-4118. Yakovlev, O.I. and Basilevsky, A.T., 1994. Experimental studies of geochemical aspects of impact cratering. In: B.O. Dressier, R.A.E Grieve and V.L. Sharpton (Edditors), Large Meteorite Impacts and Planetary Evolution. Geol. Soc. Amer., Boulder, Colorado, Spec. Paper, 293, pp. 73-80. Yamagishi, H. and Dimroth, E., 1985. A comparison of Miocene and Archean rhyolite hyaloclastites: evidence for a hot and fluid rhyolite lava. J. Volcanol. Geotherm. Res., 23: 337-355. Yamaguchi, K., 2002. Geochemistry of Archean-Paleoproterozoic black shales: the early evolution of the atmosphere, oceans, and biosphere. Ph.D. Thesis (unpubl.), Pennsylvania State Univ., 465 pp. Yarmolyuk, V.V. and Kovalenko, V.I., 1991. Riftogenic Magmatism of Active Continental Margins and Its Ore Content. Nauka, Moscow, 263 pp. (in Russian). Ye, L. and Kerr, D., 2000. Sequence stratigraphy of the Middle Pennsylvanian Bartlesville Sandstone, northeastern Oklahoma: a case of an underfilled incised valley. Amer. Assoc. Petrol. Geol. Bull., 84(8): 1185-1204. Yeo, G.M., 1981. The Late Proterozoic Rapitan glaciation in the northern Cordillera. In: EH.A. Campbell (Ed.), Proterozoic Basins of Canada. Geol. Surv. Can., Ottawa, pp. 25-46. Yeo, G.M., 1986. Iron-formation in the late proterozoic Rapitan Group, Yukon and Northwest Territories. In: J.A. Morin (Ed.), Mineral Deposits of the Northern Cordillera. Can. Inst. Mining and Metallurgy, pp. 137-158. Yin, Q., Jacobsen, S.B., Yamashita, K., Blichert-Toft, J., Trlouk, P. and Albarrde, F., 2002. A short timescale for terrestrial planet formation from Hf--W chronometry of meteorites. Nature, 418: 949-952. Young, E.D. and Russell, S.S., 1998. Oxygen reservoirs in the early solar nebula inferred from an Allende CAl. Science, 282: 452-455. Young, E.D., Ash, R.D., England, P. and Rumble, D., III., 1999. Fluid flow in chondrite parent bodies: deciphering the compositions of planetesimals. Science, 286:1331-1335. Young, G.M., 1970. An extensive early Proterozoic glaciation in North America? Palaeogeog. Palaeoclimatol. Palaeoecol., 7: 85-101. Young, G.M., 1973. Tillites and aluminous quartzites as possible time markers for Middle Precambrian (Aphebian) rocks of North America. In: G.M. Young (Ed.), Huronian Stratigraphy and Sedimentation. Geol. Assoc. Can., pp. 97-127. Young, G.M., 1976. Iron-formation and glaciogenic rocks of the Rapitan Group, Northwest Territories, Canada. Precambrian Res., 3: 137-158. Young, G.M., 1981. The Early Proterozoic Gowganda Formation, Ontario, Canada. In: M.J. Hambrey and W.B. Harland (Eds.), Earth's Pre-Pleistocene Glacial Record. Cambridge Univ. Press, Cambridge, pp. 807-812.

920

References

Young, G.M., 1983. Tectono-sedimentary history of Early Proterozoic rocks of the northern Great Lakes region. Geol. Soc. Amer. Mem., 160:15-32. Young, G.M., 1988. Proterozoic plate tectonics, glaciation and iron-formations. Sediment. Geol., 58: 127-144. Young, G.M., 1991a. The geologic record of glaciation: relevance to the climatic history of the Earth. Geosci. Canada, 18:100-108. Young, G.M., 199 lb. Stratigraphy, sedimentology and tectonic setting of the Huronian Supergroup. Geol. Assoc. Canada, Toronto'91 Field Trip B5, Guidebook, 34 pp. Young, G.M., 1992. Neoproterozoic glaciation in the Broken Hill area, New South Wales, Australia. Geol. Soc. Amer. Bull., 104: 840-850. Young, G.M. and Gostin, V.A., 1989. An exceptionally thick upper Proterozoic (Sturtian) glacial succession in the Mount Painter area, South Australia. Geol. Soc. Amer. Bull., 101: 834-845. Young, G.M. and Gostin, V.A., 1991. Late Proterozoic (Sturtian) succession of the North Flinders Basin, South Australia; an example of temperate glaciation in an active rift setting. In: J.B. Anderson and G.M. Ashley (Eds.), Glacial Marine Sedimentation; Paleoclimatic Significance. Geol. Soc. Amer., Boulder, Colorado, pp. 207-222. Young, G.M. and Long, D.G.E, 1976. Ice-wedge casts from the Huronian Ramsay Lake Formation (> 2,300 m.y. old) near Espanola, Ontario, Canada. Palaeogeogr. Palaeoclimatol. Palaeoecol., 19: 191-200. Young, G.M. and Nesbitt, H.W., 1985. The Gowganda Formation in the southern part of the Huronian outcrop belt, Ontario, Canada: stratigraphy, depositional environments and regional tectonic significance. Precambrian Res., 29:265-301. Young, G.M. and Nesbitt, H.W., 1999. Paleoclimatology and provenance of the glaciogenic Gowganda Formation (Paleoproterozoic), Ontario, Canada: a chemostratigraphic approach. Geol. Soc. Amer. Bull., 111: 264-274. Young, G.M., von Brunn, V., Gold, D.J.C. and Minter, W.E.L., 1998. Earth's oldest reported glaciation: physical and chemical evidence from the Archean Mozaan Group (~ 2.9 Ga) of South Africa. J. Geol., 106: 523-538. Young, G.M., Long, D.G.E, Fedo, C.M. and Nesbitt, H.W., 2001. Paleoproterozoic Huronian basin: product of a Wilson cycle punctuated by glaciations and a meteorite impact. Sediment. Geol., 141/142: 233-254. Young, R.B., 1928. Pressure phenomena in the dolomitic limestones of the Campbell Rand Series in Griqualand West. Trans. Geol. Soc. S. Afr., 31: 157-165. Yuasa, M., 2003. Submarine silicic calderas on the northern Shichito-lwojima Ridge, Izu-Ogasawara (Bonin) Arc, western Pacific. In: J. White, D. Clague and J. Smellie (Eds.), Explosive Subaqueous Volcanism. Amer. Geophys. Union, Monograph (in press). Yuen, D.A., Hansen, U., Zhao, W., Vincent, A.P. and Malevsky, A.V., 1993. Hard turbulent thermal convection and thermal evolution of the mantle. J. Geophys. Res., 98: 5355-5373. Yurimoto, H., Ito, M. and Nagasawa, H., 1998. Oxygen isotope exchange between refractory inclusion in Allende and solar nebula gas. Science, 282:1874-1877. Zaback, D.A., Pratt, L.M. and Hayes, J.M., 1993. Transport and reduction of sulfate and immobilization of sulfide in marine black shales. Geology, 21: 141-144. Zaitlin, B.A., Potocki, D., Warren, M.J., Rosenthal, L. and Boyd, R., 2000. Sequence stratigraphy in low accommodation foreland basins: an example from the lower Cretaceous Basal Quartz Formation of southern Alberta. Abstracts, GeoCanada 2000 conference, Can. Soc. Petrol. Geol. (CD-ROM).

References

921

Zajac, I.S., 1974. The stratigraphy and mineralogy of the Sokoman Formation in the King Lake area, Quebec and Newfoundland. Geol. Surv. Can. Bull., 220, 159 pp. Zeebe, R.E. and Wolf-Gladrow, D., 2001. CO2 in Seawater: Equilibrium, Kinetics, Isotopes. Elsevier, Amsterdam, 346 pp. Zegers, T.E., 1996. Structural, kinematic and metallogenic evolution of selected domains of the Pilbara granite-greenstone terrain; implications for mid-Archean tectonic regimes. Ph.D. Thesis (unpubl.), Utrecht Univ., Geologica Ultraiectina, 146:208 pp. Zegers, T.E. and Van Keken, P.E., 2001. Middle Archean continent formation by crustal delamination. Geology, 29: 1083-1086. Zegers, T.E., White, S.H., de Keijzer, M. and Dirks, P., 1996. Extensional structures during deposition of the 3460 Ma Warrawoona Group in the eastern Pilbara craton, Western Australia. Precambrian Res., 80:89-105. Zegers, T.E., de Keijzer, M., Passchier, C.W. and White, S.H., 1998a. The Mulgandinnah Shear Zone; an Archaean crustal scale shear zone in the eastern Pilbara, Western Australia. Precambrian Res., 88: 233-247. Zegers, T.E., de Wit, M.J., Dann, J. and White, S.H., 1998b. Vaalbara, Earth's oldest supercontinent: a combined structural, geochronologic, and paleomagnetic test. Terra Nova, 10: 250-259. Zegers, T.E., Nelson, D.R., Wijbrans, J.R. and White, S.H., 2001. SHRIMP U-Pb zircon dating of Archaean core-complex formation and pan-cratonic strike-slip deformation in the Pilbara craton. Tectonics, 20: 883-909. Zhao, J.X., Mcculloch, M.T. and Korsch, R.J., 1994. Characterization of a plume-related ~ 800 Ma magmatic event and its implications for basin formation in central-southern Australia. Earth Planet. Sci. Lett., 121 : 349-367. Zhao, W.L., Yuen, D.A. and Honda, S., 1992. Multiple phase transitions and the style of mantle convection. Phys. Earth Planet. Interiors, 72:185-210. Zimmer, M., Kr6ner, A., Jochum, K.P., Reischmann, T. and Todt, W., 1995. The Gabal Gerf Complex: a Precambrian N-MORB ophiolite in the Nubian Shield, NE Africa. Chem. Geol., 123: 29-51. Zinner, E., 1997. Presolar material in meteorites: an overview. In: T.J. Bernatowicz and E. Zinner (Eds.) Astrophysical Implications of the Laboratory Study of Presolar Materials. Proceedings, AlP conference, St. Louis, Missouri, 1997, pp. 27-58. Zinner, E. and G6pel, C., 1992. Evidence for A1-26 in feldspars from ordinary chondrites. Meteoritics, 27:311-312. Zlobin, V.L., Rosen, O.M. and Abbyasov, A.A., 2002. Two metasedimentary basins in the Early Precambrian granulites of the Anabar Shield (polar Siberia): normative mineral compositions calculated by the MINLITH program and basin facies interpretations. In: W. Altermann and P.L. Corcoran (Eds.), Precambrian Sedimentary Environments: A Modern Approach to Ancient Depositional Systems. Blackwell, Oxford, pp. 275-291. Zwanzig, H.V., Syme, E.C. and Gilbert, H.P., 1999. Updated trace element geochemistry of ca. 1.9 Ga metavolcanic rocks in the Paleoproterozoic Lynn Lake belt. Manitoba Industry, Trade and Mines, Geol. Ser., OF 99-13, 48 pp.

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923

SUBJECT INDEX

A A or B horizon 490 Abbreviated palaeotidal record 474 Abitibi greenstone belt 66, 88, 174, 294 Abyssal impact layers 47 Acasta Gneiss Complex 25 Acceleration--deceleration cycles 638 Accommodation 685 Accommodation space creation 675 Accretion of island arcs 680 Accretionary and armoured lapilli 38 Accretionary orogens 223 Accretionary prism 229 tectonics 205 Acraman impact structure 38 Acritarchs 541 Active margins 65 Actualism 593 Adakites 108 Adakitic melts 87 Adamastor ocean 471 Adhesion laminae 646, 656 ripples 654 structures 642, 651 Aeolian erg 593 transport 601 Aeolianites 601 AFM 525 Aggressive weathering 678 Air-fall deposit 49 Akia terrane 77 Akitkan belt 233 Akulleq terrane 77 Al-depleted flows 291 Al-undepleted flows 291 AI203-CaO + N a 2 0 - K 2 0 plot (A-CN-K triangle) 484 AI203-CaO + Na20 + K20-Fe203(T) + MgO (A--CNK-FM triangle) 484

AI203/TiO2 ratio 297 Alamo impact breccia 48 Alaska 451 Aldan Province 234 Algae 361 Algoma and Superior BIF 379 Algoma type IF 404 ALH84001 517 Alkaline lake deposit 601 Alkalinity 546 Alkalinity pump 549 Allende carbonaceous chondrite 11 Allochthonous 89 Allostratigraphy 694 Alluvial facies 597 fan deposits 600 fans 139, 661,678 systems 599 Alpine-style thrusting 125 Altered (CV3)chondrite Mokoia 13 Aluminous clays 502 Amazon craton 405 Amino acids 575 Ammonia 14 Amphibolite facies metamorphism 80 Amphibolites 78 Amygdule 328 Anaerobic microbial communities 660 photosynthesis 671 Andean-type model 233 Angrite parent body 12 Anhydrous mantle melting 297 Annual (or seasonal) oscillation of sea level 450 oscillation of sea level 476 Anorogenic magmatism 232 Anorthositic 19 Antarctica 450 Antidunes 49

924

Apex basalt 120 chert 532 Apparent polar wander 257 polar wander path 223 Appelella ferrifera 73 Arabian-Nubian Shield 333 Aragonite 560 Aragonite compensation depth 660 Arc complex accretion 162 Archaean and modern sedimentation rates 677 Archaean arcs 315 atmospheric-hydrospheric evolution 359 calderas 345 climate 457 Earth-Moon dynamics 623 fossil record 516 gneiss complex 66 greenstone belts 161,597 metasediments 625 oceanic crust 321 plate tectonics 201 successor basins 318 Archaean-Proterozoic boundary 679 Architectural element 663 Arctic 450 Areachap arc 243 Ash-flow calderas 345 deposits 238 Asteroids 3,575 Asthenosphere 196 Asymptotic Giant Branch (AGB) star 7 Athabasca basin 705 Atmospheres 14 Atmospheric CO2 452, 547 CO2 drawdown 680 02 (oxygen) 359, 495 Atomic force microscopy 525 Australian Precambrian carbon isotopic record 388 Autochthonous terranes 89 B

Bababudan Group 614 Back-arc 213 Back-arc extension 237

Subject Index

Back-scattered electron images 21 Bacteria 362, 560 Bacterial sulphate reduction 422 Baker Lake basin 183 Group 650 Baltic shield 405, 680 Banded iron-formation (BIF) 172, 346, 378, 403,551,599, 613, 660, 678 Barberton 205,582 Barberton greenstone belt 35, 111,143, 286, 599 greenstones 598 Barberton-type 294 Barchanoid dunes 645, 649 Barite 153 Barrier island-lagoon-washover systems 597 islands 598 Basal conglomerate 630 heat flux 237 surface of forced regression 691 Basaltic fountain 342 glasses 17 Base level 681 Base-of-slope submarine clastic systems 150 Basin architecture 139 Basin-margin faults 619 Basnaering delta complex 599 Bed-load dominated, sheet-like, braided alluvial deposits 661 Bedload 599 Belt basin 429, 660 Beniah and Bell Lake Formations 614 Benioff plane 108 Big Cottonwood Formation 478 Bimodal sequences 324 volcanism 211,273 Bimodal-bipolar palaeocurrents and patterns 632 Bio-calcification 543, 546 Biofilms 602 Biogenic methane 363 methane production 371 pyrite 380

Subject Index

925

silica precipitation 544 stromatolites 570 Biogenicity 520 Bioherm 566 Biolamination 666 Biological factors 567 Biomarkers 523 Biomineralisation 544, 547 Biophoric stromatolite 570 Biostratigraphy 542 Biostrome 566 Bioturbated 46 Bioturbation 593 Birimian belt 680 Bishunpur chondrites 12 Black shale 165 Blake River segment 97 Bolide impacts 593 Bomb sags 330 Boninites 84, 109, 325 Borgmassivet Intrusive Suite 245 Boron 600 Bouger gravity 129 Braid-deltaic 658 Braided and ephemeral systems 661 Braided fluvial, braid-delta and tidal flat depositional systems 598 Braidplain systems 599 Brazilian Shield I 18 Brittle-ductile transition 116 Bubble cavities 34 Buck Ridge (BR-) chert 143 volcano-sedimentary complex 143 Budjan Creek Formation 120 Buffalo Head terrane 230 Buhwa belt 614 Buoyancy forces 104 Burial flux of organic carbon 363, 386 Bushmanland terrane 241 Bushveld Complex 221 Butler Hill palaeosol 492 C C isotopic record 472 13C-depleted carbon in graphite 72 13C/12C ratios 460 Ca- and Al-rich inclusions (CAIs) 8 CaCO 3 547 Cadillac-Larder Lake Fault Zone (CLLFZ) 97

89,

Calcification 562 Calcium carbonate precipitation 547 Caldera lakes 152 Caldera-fill sequence 330 Calderas 325 Callisto 576 Campbellrand Subgroup 584 Cap carbonates 360, 441,453 Capricorn orogen 390 Carawine layer 33 Carbon cycle 165, 388 isotope excursions 401 isotopes 360 isotopic fractionation 370 isotopic record 393, 402 sink 548 Carbonaceous chondrites 13 (CI) chondrites Orgueil and Ivuna 13 matter 518, 523 Carbonate 545 Carbonate platform 660, 678 Carbonate-associated sulphate 431 Carbonate-BIF platformal sequences 679 Carlindi granitoid complex 128 Cartwright Hills 91 Caste flysch basins 102 Catastrophic close approach 473 Cathodoluminescence imaging 20 Ce anomalies 80 Cellular fossils 519 Central Pilbara tectonic zone 129 Chainpur 12 Chambers Bluff tillite 450 Channel bank stability 599 "Chaos terranes" 154 Chapais syncline 93 Chemical index of alteration (CIA) 486 "rain-out" sediment 599 varve 410 weathering 483,494 weathering of igneous rocks 483 Chemostratigraphic profiles 465 Cherts 139 Chesapeake Bay impact event 46 Chibougamau pluton 93 Chicobi sedimentary sequence 91

Subject Index

926

unit 93 Chondritic meteorite Zag 13 meteorites 12 Chondrules 12 Chromite deposits 222 Chuniespoort-Ghaap Groups 679 Churchill Province 183 Circum-Karelian belts 228 Clast size measurements 628 Clay minerals smectite, illite, kaolinite, gibbsite and chlorite 484 "Climate friction" 457 Climatic zonation 456 Close approach at c. 1.9 Ga '481 of the Moon 473 CO2 602 Coagulation cumulates 309 Coarsening-upward sequences 614 Coccoidal cyanobacteria 535, 554 fossils 521 Collapse structures 140, 345 Collapsed continental rifts 223 Columbia supercontinent 392 Comets 27, 575 Composition of weathering profiles, palaeoprofiles (palaeosols) 483 Compositional trends on the A-CN-K and A-CNK-FM triangles 486 Compound ripples such as flat-topped, washed-out, double-crested and ladder-back 633 Confining pressures 334 Conglomerate 617 Congo craton 240 Constrained crystal growth 308 Contamination 520 Continent-continent collision 215 Continental arc 618 crust 19 crustal growth rates 594, 658 flood basalts 173,678 growth rates 679 margins 618 platforms 205 Convective centres of descent 212 Coonterunah Formation 125

Coppin Gap greenstone belt (CGB) 140 Cordilleran ophiolites 319 Core-collapse supernova events 4 Core-complex formation 117 Coronae 140 Correlative conformity 681,682 Corundum 8 Corunna Downs granitoid complex 125 Crater size distribution 48 Cratonic keels or roots 162 nuclei 161 Cratonisation 118 Cratons, Pilbara and Kaapvaal 679 Cretaceous-Tertiary (K/T) boundary layer Crust 18 Crust recycling 19 Crustal contamination 228, 294 growth 161 growth rates 598 heat productivity 237 plateaus 18 thickening 116 Cryptodomes 330 Crystal morphology 300 Crystallisation 308 Cyanobacteria 361,671 Cyanobacterial mats 554 Cyclic rhythmites of tidal origin 360 D Dales Gorge spherule layer 39 Damara belt 467 Supergroup 460 Dark nebula 4 Daspoort Formation 724 De Grey Group 125 Debris-flows 661,678 Decompression melting 115,219 Deep mantle 173 offshore shelf 419 Deformation-enhanced segregation Deformed conglomerates 628 Delamination 108 Delta deposits 598 plains 599

116

36

927

Subject Index

Deltaic successions 599 systems 598 Dendritic 305 Denudation rates 597, 658 Depleted harzburgite 112 mantle 195 Depositional environments 626 rate of BIF 415 systems 139 Despinning Earth 480 Destor-Porcupine Manneville Fault Zone (DPMFZ) 89, 97 Devitrification 34 Dhalbhum Formation 645 Dharwar craton 89, 174 Diagenetic-metasomatic alteration 493 Diamond 11 Diapiric 104 Diapiric buoyancy 206 model 119 triple points 130 Diffusion rate 308 Dilute density current deposition 337 Direct precipitation of calcium carbonate 678 Discharge rates 599 "Diurnal inequality" 474 Diurnal laminae 474 temperature range 452 temperature variations 454 Dolomitisation 544 Domal geometry 117 "Dome-and-basin" patterns 118 Dongwanzi 320 Doubly-graded sequences 336 Draa 646 Draa complexes 654 deposits 656 Dresser Formation 143, 532 Dronning Maud Land 245 DSDP site 612 40 Dubawnt Supergroup 650 Duffer Formation 120, 618 Dune and interdune deposits 651

Dune complex 651 complex and interdune deposits deposits 649 field 657 Dunes 649 Duparquet basin 98, 102 Dyke swarm 174 Dynamic loading 210

653

E Earth expansion 479 Earth-Moon dynamics 631 Earth' s forced nutations 480 heat loss 677 moment of inertia 479 obliquity 452 palaeorotation 360 past LOD 480 Precambrian rotation 473 Eastern Malartic segment 97 Ebb and flood currents 632 Ebb-tidal deltas 598 Eburnean supercontinent 213 Eburnian suture 251 Eclogite 61,108 Eclogite-delamination model 116 Ediacara 541 Ediacara fossils 469 Edifice instability 331 Efremovka carbonaceous chondrite Elatina Formation 448, 474 Electron microscopy 520 Enderbite-charnockite magmas 237 Endogenic 275 Endolithic environments 587 Enstatite chondrites 12 Entophysalis granulosa 554 Environmental factors 567 indicators 570 Epeiric embayments 657 marine sediments 657 sea model 657 sea palaeoenvironments 598 seas 598 seas sensu stricto 657 seaways 657 transgression 658

928

Ephemeral braid-delta systems 658 rivers 600 Erg 602, 650 Erg deposits 647, 654 margin/sand sheet/dune systems 642 Erg-margin sedimentation 653 Erosion rate 599, 675 Eruption-fed deposits 334 turbidity currents 334 Eucrite and augite meteorite 12 Eucrite parent body 12 Eukaryotes 366, 540 Euro basalt 125 Europa 549 Eustasy 594 Eustatic and relative sea level changes 594 Evaporite minerals 600 Evidence of life 518 Exobiology 575 Experimental conditions 305 Extension 613 Extensional collapse 116 growth faults 140 normal faults 140 structures 145 thinning 115 Extraterrestrial 575 F Facies changes 325 model 628 "Faint young Sun" 359 Falling stage 682, 697 stage systems tract 695 Fancamp deformation zone 95 Faribault faults 95 Fe-stratified Archaean ocean model 369 Features of Neoproterozoic glaciations 457 Fennoscandia 224 Fennoscandian shield 679 Field geology 606 mapping 625 studies 606 Filaments 522

Subject Index

Flash-flood events 662 Flat-subduction model 116 Flavrian pluton 97 Flin Flon greenstone belt 327 Flood-basalt resurfacing 20 Flood-tidal and washover facies 598 Flooding surfaces 694 Flow architecture 286 fields 290 Fluvial and braid-delta sheet sandstones Fluvial braidplain deposits 657 braidplains 600 Forced regressions 685 regressive 682, 691 Forearc 213 Foreland 230 Foreland basin 210 Foreset bundle 637 Foreshore deposits 597 Fortescue 174 Fortescue Group 541 Fossils 518 Fountaining eruptions 338 Fractional crystallisation 297 Fragmentation processes 279 Fred's Flow 309 Free nutation of the fluid core 480 Freeboard 658 Freeboard concept 594 Fumarolic volcanicity 419 Fungi 560 Fusion reactions 3

598

G Gabbro-anorthosite-rapakivi granite magmatism 232 Gariep and Damara orogenic belts 461 Gariep Supergroup 462 Garnet-beating source 297 Gas-supported 337 Geocentric axial dipole (GAD) 455 Geochemical cycle of carbon 369 of carbon and oxygen 385 of sulphur 369, 380 Geodynamic models 255 Geoid 257 Geoid relief 594

Subject Index

Geostrophic currents 657 Geotherms 107 Ghaap-Chuniespoort succession 659 Ghanzi--Chobe rifts 254 Giant Impact Hypothesis 14 Glacial cycles 453 deposits 602, 661 systems 602 Glaciations 440 Glaciogenic 227 Glaciogenic and periglacial processes 602 Glaciogenic deposits 602 setting 593 Glacioisostacy 594 Glaciomarine deposition 448 deposits 602 Global glaciation 658, 680 oceanic sulphate reservoir 422 plume events 480 sea level 658 superplume 678 warming 165 Globular structures 298 Glycocalyx 554 GodthSbsfjord 77 Gondwana 256, 459 Gorge Creek Group 125 Gowganda Formation 452 Graded bedding 627 Grain size 274 Granada basin 98 Granite diapirism 206 extraction 116 Granite-greenstone 25 Granite-greenstone crust 180 terranes 65 Granitic 180 Granitic palaeosol 493 Granitoid plutons 197 Granodiorite-granite-monzogranite (GGM) suites 104 Granodioritic 180 Granodorite 20 Granular iron-formation (GIF) 404

929

Granulite facies 204 protoliths 238 terranes 223 Granulite-gneiss belts 223 orogenic belts 162 Graphite 11 Gravel-dominated 661 Gravitational collapse folds 142 instability 106, 116 Gravity-driven deformation 119 Great Oxidation Event 379 Greenhouse 680 Greenhouse conditions 680 gas 359, 660 gases methane and ammonia 366 warming 169 Greenland 199 Greenschist 80, 204 Greenstone belts 311 synclines 119 Grenville age 240 belt 245 Grenvillian orogeny 240 Griqualand West basins 44 Growth-fault arrays 140 patterns 147 Grunehogna craton 251 Province 245 Grypania 540, 541 Gypsum deposits 678 H H-burning 8 H2S 549 Hadaean 161 Hadaean-Archaean transition 677 Haematite-rich palaeosols 680 Halite 551 Hamersley and Transvaal-Griqualand West basins 413 Hamersley basin 35, 128, 390

930

Group 679 belt 234 Harzburgites 55 Hawaiian and Icelandic plume 18 Heam domain 680 Heat flux 203 production 106 Hekpoort basaltic palaeosol 492 Helena Formation 660 Herringbone cross-bedding 632 Hibonite 8 Hierarchy 700 High accommodation systems tract 719 energy coastal sand belt 658 energy peritidal flats 658 velocity zone 112 High-discharge braidplain systems 599 High-HFSE association 84 High-Mg basaltic 19 High-Mg diorite 106 High-Nb basalts 109 High-relief, tectonically active settings 502 Highstand 695 Highstand systems tracts 682 Hinterland 230 History of the Moon's orbit 481 Hoggar belt 163 Hooggenoeg Formation 35 Hotspot 174 Hotspot chains 173 Howardite-eucrite-diogenite parent body 12 Humid climatic settings 504 Hunter Mine 346 Hunter Mine caldera 346 caldera complex 91 Huronian sedimentary basin 227 rocks 448 Huronian Supergroup 490, 679 Huroniospora-type (cyanobacteria) 73 Hurwitz Group 680 Hydrological cycle 452, 453 Hydrologically open lakes 601 Hydrostatic confining pressures 334 Hydrothermal activity 478 circulation 140 fluids 346

Subject Index

Hydrous mantle 203 Hyperconcentrated flood flow 617 flood flow and sheetflood deposits Hyperconcentrated flow 661 Hypsometric curves 594

600

I

Ice wedges 450 Icehouse conditions 680 Iceland 17, 112 Impact 548 Impact craters 18 diamond 58 structures 152 Impact-related sedimentary rocks 2 In situ crustal differentiation 108 In situ differentiation and delamination model 111 In weathering profiles 487 Indian cratons 66 Inductively coupled plasma mass spectrometry (MC-ICP-MS) 4 Inflation 288 Inner and outer shelf deposits 597 Inner shelf and fluvial overbank facies 594 Inorganic carbon 546 Integrated sedimentation rates 461 Interdune areas 649 deposits 645, 649 Intertidal to supratidal facies 660 Intra-oceanic accretionary complex 69 island arcs 677 obduction 211 Intracaldera deposits 354 sequence 354 Intracontinental basins 460 sutures 216 Intracratonic 719 Intracratonic rift 210 Intracrustal heating 115 Intraoceanic island arcs 215 Inversely graded beds 627 Inverted density profile 120

Subject Index

931

metamorphic zoning 230 Ion microprobe 4 Iridium anomalies 38 Iron and enstatite meteorites 12 Iron meteorites 12 Iron-formation (IF) 360, 403,443, 449, 453 Iron-stratified ocean model 419 Isotopic signatures 518 Isua 539 Isua greenstone belt 66, 593 Isuasphaera isua 73 Isukasia 77 Itsaq gneiss complex 77 terrane 25 J Jack Hills 21 Jadeite 61 Jeerinah Formation 39 layer 37 Joutel volcanic complex Juvenile crust 166

91

K K-metasomatism 491,493 K-metasomatism of kaolinite and feldspars 492 K20 syenite-granite (SG) suites 104 Kaapvaal and West African cratons 405 Kaapvaal craton 25, 113, 162, 657, 660, 724 Kalahari craton 217, 471 Kapunapotagen 95 Karelian craton 228 Karelian Supergroup 679 Kenorland 162, 679 Kepler's third law 478 Kerogen 524 Ketilidian orogen of South Greenland 39 Kewagama 97 Kheis belt 240 Kibaran orogeny 251 Kinga Formation 601 Kinneyia style ripples 668 Kolvitsa massif 226 Komati Formation 296 Komatiite-spinifex textures 55 Komatiites 66, 174, 271 Komatiitic 19 Komatiitic, tholeiitic and felsic volcanic and volcaniclastic rocks 678

Kromberg Formation

532

L L-tectonite fabrics 130 Labrador 26 Lac Caste Formation 97 sedimentary rocks 102 Laccoliths 104 Lacustrine Systems 600 Lagoonal deposits 598 Lake Abitibi 91 deposits 600 Motitoi 552 Superior region 680 Superior type IF 404 Van 552 Vostok 579 Lakes 600 Lalla Rookh-Western Shaw Structural Corridor 127 Lamproites 187 Lamprophyres 187 Lapland granulite belt 229 Large igneous provinces 161 impact events 45 non-dipole components 455 Laser-Raman spectroscopy 524 Late Archaean cratons of North America 66 Late Neoproterozoic Marinoan (Varanger) glaciogenic succession 474 Late-orogenic basins 620 Laurentia 224 Laurentia supercontinent 680 Lava flows 275 fountaining 346 Lava-fed density currents 335 Layered convection 165 mantle circulation 162 Limpopo belt 163, 217 Limu 342 Linear granite-greenstone terranes 104 Lionel lineament 148 Lipids 541 Lithostromes 564 Loess deposits 602

932

Lomagundi event 394 Long-term 0 2 production flux 363 sedimentation rates 675 Longitudinal dune deposits 656 dunes 656 Louvicourt Group 97 Low accommodation systems tracts 732 latitude marine glaciation 602 palaeolatitudes 448 sinuosity rivers 678 Low-angle detachment surfaces 140 Low-HFSE association 84 Low-sulphate Precambrian seawater 431 Lowstand 682, 695 Lunar laser ranging 480 mare 18 nodal cycle 478 orbit 473 semimajor axis (mean Earth-Moon distance) 478 tidal friction 473 tides 600 Lyon Lake Fault sequence 354 M Macerations 519 Mackenzie dyke swarm 455 Mafic calderas 345 granulite 108 plains 174 protoliths 204 "Mafics triangle" 484 Magadiitic cherts 601 Magaliesberg Formation 658 Magma ocean 15 Magmatic diapirs 133 fountain 338 anomalies 50 Makgabeng Formation 646 Makkovik-Ketilidian and Labradorian orogens 233 Malartic segment 99 Mallina basin 679 Manchuriophycus 667

Subject Index

Mantle 256 Mantle metasomatism 200 overturn 162 plumes 45, 65, 161,168, 273,594 wedge 106 xenocrysts 186 Marble Bar greenstone belt 142 Margate terrane 244 Marginal shear zones 130 Marine swells 600 Marinoan glaciation 448,460, 474, 479 Marmora terrane 461 Mars 3,455,548, 576 Mass balance calculations 65, 108 Mass flow 661 Massive sulphide deposits 327 Master to distributary tubes 283 Matagami complex 91 Maximum clast size 628 flooding surface 682 regressive surface 682, 692 McPhee dome 129 Mean global temperature 452 recession rate 479 Mechanical erosion 483 Meimechite 304 Meiotic cell division 541 Mrlanges 320 Melilite 8 Mercury 13, 576 Meso-macrotidal conditions 659 Mesobanding 407 Metallic core 12, 1.5 Metamorphic core complexes 132 Metasedimentary granulites 237 Metasomatic fluids 197 Metasomatised lithospheric mantle 183 Metasomatism 80 Metazoan 541 Metazoan grazing 663 Meteorite impacts 152 Meteorites 3, 575 Meteoritic solar (Ne-B) component 18 Methane ices 14 Methanogens 370 Methanotrophs 370

Subject Index

Methylhopanes 541 Mg-metasomatism 493 Mg-numbers 111 Michigan 130 Micrites 554 Microbanding 407 Microbial binding 664 communities 663 filaments 522 mats 377, 593, 664, 673, 678 micro-organisms 678 sand chips 668 wrinkle marks 668 Microbialites 552 Microbially bound surface layers 666 Microfossils 73, 539 Microkrystites 37 Microlitic textures 34 Microplate tectonics 239 Microspar 554 Microtektites 33 Microtidal marine coastlines 600 Mid-ocean ridge 161,333,677 ridge basalt 17 Mid-oceanic rifts 65 Migmatitic gneiss 219 Minarets Caldera in California 153 Minas Supergroup 601,643,679 Mineral lineations 132 Minette 183 Misleading palaeotidal data 474 Mobile belt 223, 311 Moelv glaciation 460 Molasse sequences 613, 619 Molecular cloud 7 Monteville Formation 39 "Monthly inequality" of spring-tidal ranges Monzogranite 128 Moodies Group 147, 479, 631 Moon 2, 548 Moon's orbit 360 origin 457 Mount Edgar dome 130 Edgar granitoid complex 125 Narryer 21

933

Mozambique belt 251 Muccan dome 127 granitoid complex 128 Mucilage 554 Muddy back-barrier subtidal deposits Mudstone drapes 641 Mulgandinnah 148 Munro-type 291 Muong Nong-type tektites 33 Murrurundi profile 492 Mylonites 73 Mylonitic zone 218 Myojin Knoll seafloor caldera 337 Mzumbe terrane 244

474

N Na-metasomatism 493 Na-metasomatism of feldspars 493 Nama basin 467 Namaqua belt 240 Namaqua-Natal and Maud belts 240 Namaqualand 240 Natal belt 240 Native elements 59 Nb anomalies 80 Nd isotope values 294 Neap tides 474 Neap-spring (fortnightly) cycles 474 Neap-spring-neap cycles 637 Negative 13C excursions 459 13C values for carbonates 452 Neguanee banded iron-formation (BIF) Neoproterozoic 256 Neoproterozoic Earth-Moon dynamics 598 Gariep belt 340 glaciation 448 glaciogenic successions 442, 453 ice ages 393 sedimentation rates 459 New Jersey 40 New Quebec belt 228 Noble gases 15 Noranda caldera 346 Normal grading 627 regressions 685 regressive 691

598

540

Subject Index

934

Normetal fault 95 volcanic complex 91 North China craton 405 North Pole dome 129,143 (NP-) chert 143 volcano-sedimentary complex 143 Northern Volcanic Zones (NVZ) 89 Northwest Territories, Canada 25 Nova 8 Nucleation 301 Nucleosynthesis 1 O O horizon 490 Obductive arc complexes 162 Obliquity and glacial climate 455 Obliquity of the ecliptic 455 origin 457 Ocean-atmosphere general circulation model 453 Ocean-floor volcanism 331 Oceanic crust 20 flood basalts 173 plateau 106, 162, 169, 333,677 pycnocline 360 Octupole component 455 Omphacite 61 Ontario, Canada 626 Ontong Java Plateau 112 Onverwacht Group 35, 143 Opatica belt 91 Open seas 449, 453 Open-ocean tides 657 Ophiolite complexes 163 nappes 214 Ophiolites 213, 333,678 Ophiolitic m61ange 333 Organic production 543 Ortega and Uncompahgre Groups 631 Orthogneiss 118 Ovoid 104 Oxygenic photosynthetic organisms 363 P p-process 4 Palaeoclimate

594

Palaeocurrent direction 49 Palaeoenvironments 593 Palaeohydrological parameters 600 Palaeomagnetic poles 223 Palaeomagnetism 255 Palaeontological "clocks" 473 Palaeopangea 162 Palaeoproterozoic glacial deposits 444 glaciation 452 iron-formations 478 Ketilidian mobile belt 340 Palaeorotational data 602 Palaeosalinity 600 Palaeosols 495, 545 Palaeosols, laterites and red beds 376 Palaeosols developed on igneous rocks 482 Palaeozoic granodioritic palaeosol 492 Palimpsest ripples 664 Pallasites 12 Pan-African 461 Pan-African orogen 324 tectonic cycle 244 Panorama Formation 120 Para-autochthonous 230 Parasequence 594, 682 Partial melting 206 Passive margin carbonates 678 margins 173 Pb diffusion 21 Pechenga-Varzuga 228 Pelagic sedimentation 658 Pele's hair 38 Peloids 556 Penokean orogen 130, 232 Perigean-apogean 638 Periglacial regions 450 sand wedges 452 Periodic tipping of the rotational axis 361 Periodicity in komatiite activity 45 Permian 452 Perovskite 8 Petees 666 Petrographic thin sections 519 Photosynthesis 361,547, 561,666 Phreatomagmatic eruption 328, 338

Subject Index

PIA (plagioclase index of alteration) 497 Picrites 87 Piecemeal configuration 354 Pilbara 66, 582 Pilbara craton 20, 89, 111, 174, 319, 321,390, 541 Supergroup 120 Pilgangoora syncline 132 Pillow breccia 69 Pillowed 327 Plagioclase spherulites 303 Planar deformation features (PDFs) 47 Planetary accretion 14 embryos 1 obliquity 455 scale circulation cells 601 Planetesimals 3 Platform carbonates 390 Platinum group elements 38 Platte-type river deposits 662 Playa 657 "Plughole" model 161,206 Plume heads 162 "Plume tectonics" 161 Plumes 255 Point-bar deposits 599 Polycyclic aromatic hydrocarbons (PAHs) 575 Polygonal convection systems 207 systems 206 Polymictic conglomerate 70 Polyphase deformation 80 Pongola Supergroup 542 Pontiac sedimentary rocks 89 Popigai impact event 46 Porcupine sedimentary rocks 97 Porphyry stocks 620 Port Askaig tillite 460 Port Nolloth zone 461 Positive 13Ccarb 454 13Ccarb excursions 466 Powell Tuff 351 Pre-caldera basalt volcanic base 354 subaqueous basalt plain 353 Pre-Devonian fluvial systems 660 Pre-Fountain palaeosol 492 Pre-vegetation sandy systems 662

935

Pre-vegetational fluvial systems 600 Precambrian carbonate platforms 542 depositional systems 594 glaciations 602 length of day (LOD) 473 Paleobiology Research Group 527 sulphur isotope record 421 tidal, wave and storm shelf dynamics 598 tidal periods and Palaeorotation 474 Precipitation of calcite 543 Preissac-Lacorne batholith 102 Present low latitudes 452 Pretoria Group 658, 724 Primitive mantle 190 Prodelta to open shelf sedimentation 599 Prokaryotes 540 Proterozoic geomagnetic field 455 glaciations 360, 448 glaciomarine deposition 448 global environment 452 greenstone belts 312 large obliquity 455 Proto-continents 677 Proto-Earth 14 Proto-stars 4 Proto-Sun 8 Protocontinents 207 Protoplanetary disk 1 Pseudotachylites 52 Pualco Tillite 450, 476 Pull-apart or strike-slip 619 Ridge Basin, California 677 Pycnocline 419, 678 Pyke Hill 307 Pyrite conglomerates 544 pebbles 374 Pyroclast suspensions 340 Pyroclastic 274 Pyroclastic debris 355 tuff 330 Pyroxene 309

Q Quadrilatero Ferrifero 118 Quadrupole component 455

936

Quartz arenite 497, 613 enrichment 497 Quartzite 72 Quench textures 53 Quenching 34 R

r-process 3 Radial 34 Radiogenic heat 201 Radiometric dating 677 Raman spectroscopy 524 Ramsay Lake Formation 452 Rapitan glaciation 448 Group 460 iron-formation (IF) 420, 454 Raquette Formation 618 Rate of chemical weathering 483 lunar recession 481 mechanical erosion 483 subsidence 675 the Earth's rotation 660 Ravinement surface 693 Reactivation surfaces 637 Recurrence interval (RI) 41 Recycling 65 Red beds 496, 544 giants and supergiants 8 Red Sea-type spreading 238 Reducing atmosphere 364 Regional-scale fold interference 119 Regression 682 Regressive 686 Regressive ravinement surface 692 surface of marine erosion 682, 692 Regressive-transgressive sequence 682 Relative rates of chemical weathering and mechanical erosion 490 Resonances of the fluid core 480 Retroarc foreland basin 679 Reworked deposits 334 Reworking and recycling 675 Reynella Siltstone 474 Rhythmic bedding 633 Rhythmites 600

Subject Index

Richtersveld terrane 241 Rift basins 678 Rift-related, hydrothermally influenced basins 454 Ring faults 130 Rio de la Plata craton 471 River regime, wave energy and tidal range 599 Rivers and fans 661 Rodinia 162, 257 Rodinia supercontinent 393 breakup 454 Rolled up mat fragments 668 Roof pendant 132 Rotation of the early Earth 601 rRNA 516 Runoff rates 599 S s-process 3 Sagduction 66 Saldania belt 470 Salgash Subgroup 143 Saline lake deposits 601 Salinity stratification 660 Sand cohesive and even thixotropic 673 sheet deposits 601 wedges 450 Sand-wedge polygons 454 Sandsheet 650 Sandsheet deposits 643, 645, 650, 654, 656 Sandwave deposits 598 Sandy braided river system 663 Sanukitoid 106, 109 S~o Francisco craton 199, 679 Satonda Island 552 Saturation index 546 Schist 78 Sea level 165, 677 Sea-floor hydrothermal alteration 80 Seamount 325 Seamount Six 342 Seasonal changes 450 cycle 456 temperature ranges 452 Seasonality paradox 450 Seawater 546 Secular carbon isotope curve 399

Subject Index

change 457 change of obliquity 457 Sediment accumulation rates 543 binding and trapping 666 bypassing 675 recycling 677 reworking 673 supply 675 Sediment-gravity flows 340 Sedimentary basin geometry 148 structures 602, 626 Sedimentation rate 675 Seif dune 645 SEM 520 Semi-perennial fluvial systems 600 Semidiurnal bundles 638 laminae 474 tides 474 Sensitive High-Resolution Ion MicroProbe (SHRIMP) 20 Sequence 681 Sequence hierarchy 699 stratigraphy 681 Shallow subduction 108 water 327 Shaw 130 Shaw granitoid complex 128 Sheaf-like structures 305 Sheaths of filamentous cyanobacteria 667 Sheet flow 279, 327 hyaloclastites 342 Sheet-floods 662 Shelf and epeiric seas 657 environments 139 Shelf-breaks 657 Shelf-like palaeoenvironments 657 Shield morphology 329 volcano 316 Shocked quartz 47 zircon 58

937

Shoreface deposits 597 dynamics 597 Shoreline 685 Shoreline forced regression 692 regression 685 transgression 685 Short-lived nuclides 1 Shoshonites 318 Siberian cratons 224 Siderophile elements 38 Sigmoidal foreset bundles 637 Silicate weathering 452, 547, 550 Silicon carbide 11 nitride 11 Siltstone-shale couplets 633 Silverton Formation 658 "Single giant impact" hypothesis 457 Sioux Lookout belt 626 Skeletal 301 Skeletal spinel 34 Slave Province 174, 498, 679 Slope 150 "Slushball" Earth 453 Sm-Nd isotopes 677 Small scale structures 625 Snowball Earth 584 Snowball Earth hypothesis 360, 440, 452, 602 Snowbird tectonic zone 185 Soda ocean 543,547 Soil profiles 599 Solar luminosity 602 Ne 17 radiation 455 system 1 Solar-like Ne 17 South Australia 38 South Kittys Gap volcano-sedimentary complex 143 South Saskatchewan or Brahmaputra-type rivers 662 Southern Cross granite-greenstone terrane 21 Southern Volcanic Zone (SVZ) 89, 350 Southwestern USA 680 Soutpansberg volcanism 221 Spheroids 521 Spherules 2

938

Spherulitic aggregates 34 morphology 301 Spinel 8 Spinifex 298 Spinifex paradox 307 Ridge 279 textures 277 Stable platform 613 Stanovoy Province 236 Stars 3 Steep Rock sequence 319 "Stepwise oxidation" of the Proterozoic atmosphere 428 Stillwater Complex 602 Storm waves, wind-driven surface currents 657 Stratified ocean 419 Stratovolcanoes 325 Strelley Pool chert 120, 532 Strike-slip basins 678 Stromatolite classification 565,567 decline 543,569 definition 564 stratigraphy 569 taxonomy 565 Stromatolite-bearing carbonate 614 Stromatolites 172, 518, 542, 545,564, 570, 600 Stromatolitic carbonates 465 evaporites 678 Strongly directional ocean-type palaeocurrents 657 seasonal climate 450 Strontium isotopic composition 467 Structural amplification 133 Sturgeon Lake caldera 346 Sturtian and Marinoan glaciations 395 glaciation 448, 460, 479 glaciogenic succession 476 iron-formation 454 Subaerial unconformity 681 Subaqueous density currents 334 eruptions 334, 336 plains 315 pyroclastic flows 334 rhyolite domes 330

Subject Index

Subducting slabs 199 Subduction 65, 161 Subduction-related arc volcanism 66 Submarine fan 618 hydrothermal fluids 378 ridges 173 Sulphate reduction 549 Sulphate-reducing bacteria 368, 382 microbes 371 Sulphide-oxidising bacteria 671 Sulphidic Proterozoic ocean model 438 Sulphur isotope record 382 Springs Group 125 Supercontinent 602, 679 Supercontinent breakup 175 cycle 162, 167, 215,402, 602 tectonics 239 Supercontinental cycle 680 fragmentation 679 Supercooled komatiite 308 Superior craton 679 Province 66, 174, 405, 626 Supernova 1, 8 Superplume 161,163 Superplume event 163, 679 Supersaturated 307 Supra-subduction settings 237 zone 215 Surtseyan-style eruptions 340 Sutures (collisional orogens) 223 Svecofennian 324 Svecofennian accretionary orogen 232 Syenogranite 128 Syndepositional tectonics 660 Synformal and antifonnal structures 243 Synorogenic basins 317 sequences 615 Synrift and craton cover sequences 613 Synsedimentary 147 Synvolcanic faults 351 Systems tracts 682 Systems tracts, transgressive and regressive

683

Subject Index

T T-Tauri 8 Talga-Talga Subgroup 143 Taltson magmatic zone 197 Taltson-Thelon orogenic belt 230 Talus scree 617 Tasiusarsuaq terrane 77 Taupo backarc zone of New Zealand 153 Tectonic decoupling 230 Tectonised sedimentary rocks 629 Tectonothermal reworking events 480 Tectosphere 260 Tektites 33 TEM 520 Tempestites 660 Terrane juxtaposition 205 Terrestrial biomats 377 impact structures 2 O isotope fractionation line 14 Tethyan ophiolites 319 Theia 14 Thermal conductivity 237 Thermo-mechanical erosion 280 Thermocline 660 Thickness of a palaeosol 490 Thrombolites 554 Thrust-nappe 73 Ti enrichment 502 Tidal action 597 bedding (flaser, wavy and lenticular) 633 channel deposits 598 deposits 598 flat sediments 598 inlets and tidal deltas 598 range 657, 659 resonance 657 rhythmite data 473 rhythmite records 473 rhythmites 450, 473, 598, 633 sand ridges 598 Tidalites 448, 614 Tidally influenced Precambrian shelves 598 Tide-dominated shallow marine systems 598 Tide-influenced cherty 150 Tides 621 Tilt of the rotational axis 660 Timeball Hill Formation 658 Times of core resonance 480

939

Timing of glaciations 459 Tonalite 20 Tonalite-trondhjemite-granodiorite (TTG) 315 Tonalite-trondhjemite-granodiorite (TTG) suites 104 Tonalitic 180 Top indications 627 Trace fossils 541 Trans-Hudson 706 Trans-Hudson orogen 185, 333 terranes 324 Transformist 201 Transgression 682 Transgressive 682, 686 Transgressive ravinement surface 694 surface 692 Transgressive-regressive sequence 682 shorelines 328 Transitional and offshore mud belts 658 Transpressive orogeny 220 Transvaal basin 162 Supergroup 658, 659, 724 Transverse aeolian dunes 649 dunes 645,654 Tree of Life 516 Trollheim-type fan deposits 661 Trondhjemite 20 Trondhjemitic 180 Troy Quartzite 653 True polar wander 256 Truncation of the secondary mineral assemblages of palaeosol 489 Tsunami 48, 660 Tube-shaped komatiites 279 "Tuff-wacke" problem 626 Tugela terrane 244 Turbidite 71,678 Turbidite deposits 617 Turbidity current 658 U U-Pb dating 20 Ultrapotassic magmatic events provinces 199 rocks 183

183

Subject Index

940

Umkondo magmatism 254 Unconfined braid-plain systems 661 Unconformities 619 Undercooling 301 Underplating 115 Uniformitarian 201 Uniformitarianism 593 Unimodal sequences 324 Uplift rate 675 Upper Bhander Sandstone 654 Upper mantle 17 Upper Mount Guide Quartzite 631 Upwelling deep ocean water 378 mantle 181 of deep waters 454 Uraniferous quartz-pebble conglomerates 380 Uraninite 371,544 Uranium deposits in quartz-pebble conglomerates 376 Urey-Reaction 548 V "Vaalbara" supercontinent 679 Validity of tidal rhythmite data 478 Varanger tillites 460 Variable salinities 657 Variations of the C isotopic composition of seawater proxies 459 Varioles 298, 303 Varvites 452 Vegetation-free landscapes 597 Vendian acritarch populations 454 Vendozoa 541 Ventersdorp 174 Ventersdorp Supergroup 601 Ventifacts 601 Venus 18, 140, 576 Vertical tectonics 66 Vesicular 288 Vesicularity 337 4 Vesta asteroid 12 Vetreny belt 228 Ville Marie palaeosol 490 Viscosity 277 Volcanic cyclicity 312 plains 18 rises 18 terminology 273 Volcaniclastic 273

Volcaniclastic debris 678 Volcano-sedimentary belts 229 greenstone sequences 104 sequences 613 Vredefort structure 27 W Warrawagine 128 Warrawoona Group 20, 89, 542 Warrawoona syncline 132 Water 14 Water depth 139 Waterberg Group 631 Wathaman batholith 197 Wave and tidal action and turbidity currents 678 height and water depths from wave ripple forms 658 Wave- and storm-dominated shallow marine systems 594 Waves 621 Wavy-crinkly laminae 669 Way-up indications 627 Weathering of igneous rocks 483 profile 483,496 profiles (palaeosols) 360 rates 599 Weeli Wolli Formation 478 West Greenland 25, 66 Westem Australia 20 Western Greenland 680 Wet interdune deposits 657 Wharton Group 650 White smokers 153 Whiterock Member 601 Whole-mantle convection 162, 165 Wide, low-angle shelves 597 Wilgerivier Formation 600 Wilson cycles 679 Wind stress 600 systems 601 Wind-ripple laminae 646, 654 migration 653 stratification 642

Subject Index

Wit Mfolozi Formation 542 Wittenoom layer 35 Witwatersrand basin 162, 602, 657 Supergroup 601,614, 631 Wolf-Rayet stars 8 Wopmay accretionary orogen 233 Wyman Formation 120 Wyoming craton 199 X Xenoliths

108, 219

941

Y Yavapai-Mazatzal-Midcontinent orogen 233 Yilgalong granitoid complexes 128 Yilgam craton 20, 66, 180, 319, 390, 593 Yule granitoid complex 128 Z Zimbabwe 104 Zimbabwe craton 116, 205,679 Zircon 4 Zonal surface winds 456 Zonations in both bulk chemical and mineralogical compositions 487

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