The Precambrian Earth: Tempos and Events (Developments in Precambrian Geology)

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

Author: P.G. Eriksson | W. Altermann | D.R. Nelson | W.U. Mueller | O. Catuneanu

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Developments in Precambrian Geology, 12

THE PRECAMBRIAN EARTH: TEMPOS AND EVENTS


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


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)


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)


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)


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


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

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1 3 27 45 62

65 65

66

74

88 103

118 139 155

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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


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


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 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

68

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).


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

70


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


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.,

72


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.


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


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).

76


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


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


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


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.


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.


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

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

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


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-


100 (a)

==

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85

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sequence

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Inner arc sequence

10 3E

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I

I

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i

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


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).


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).


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


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

$


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

96


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).


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

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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


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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

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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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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

2.5. Granite Formation and Emplacement

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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).


105

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,


107

Temperature(C) 0 O=

200 m

I

400 m

I

600 m

I

1000

800 ,

1200

I

10,2 o

0.5

20o

30-

km

1.0 2

tO v

5%

L_

409

9 9

50-

60-

~30% ?, ,~

or}

1.5 IX.

&W"

2.0

70-

Fig. 2.5-2. Pressure and temperature diagram showing the solidus for dehydration melting of amphibolite (Wyllie et al., 1997), melt contours for percentages TTG melt from amphibolite (stippled) (Rapp, 1997), solidus for biotite dehydration melting of tonalite, and water saturated solidus for tonalite melting. Garnet is formed in the residue below the gt-in line (Wolf and Wyllie, 1993). The shaded area indicates pressure and temperature conditions at which the metamorphic facies of basalt is such (eclogite and kyanite granulite) that the density exceeds 3.3 t/m 3 (Doin and Henry, 2001), the density of the depleted harzburgite lithospheric mantle.

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).

110


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).


111

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

112


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


113

(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.

114


0

(a)

.

Temperature (C) 400 600

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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


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-


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


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


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.


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

$

I. CO

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


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


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

"~ 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

2.8. C r u s t a l G r o w t h R a t e s

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

156


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.


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.


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

161

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

162

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

164


1 i

I

:

r

'1

1

1

1

1 --r,--]-

I

~--l

i

1

1

!

0.8 ~ ..

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--

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Age (Ga) Fig. 3.2-1. Frequency distribution of juvenile continental crust based on a total volume of continental crust of 7.177 x 10 9 km 3. Juvenile crust ages are U-Pb zircon ages used in conjunction with Nd isotope data and lithologic associations. Modified after Condie (1998, 2000).

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


166

Supercontinent cycle Superplumeevent

A

Breakup

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R

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9 9

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t

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9

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. m

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Fig. 3.2-2. The supercontinent cycle. Black triangles are proposed superplume events; symbol size is proportional to event intensity. After Condie (1998). R, Rodinia; G, Gondwana; Pn, Pannotia; P, Pangea; N, new supercontinent forming today.

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


167

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

168


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).


169

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

170


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.


Fig. 3.2-3.

171

172

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.

3.3. Large Igneous Province Record

173

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

174


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",


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


176

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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.


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 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

256


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-

3.11. Rapid Continental Motion is the Late Neoproterozoic

257

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

258


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.


259

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

260


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.


261

262


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


263

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.

264

Fig. 3.11-4 (continued).



265

~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.

3.12. C o m m e n t a r y

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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.

4.2. Terminology o f Volcaniclastic

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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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

4.2. Terminology of Volcaniclastic

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


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


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


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


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


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


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


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

(a)

~0

i . . . . . . . . . .

o

(e)

~0~ x Bi~ '

~00 0 + x

-~

9 E-MORB N-MORB AI-depleted AI-undepleted

- ~.-~-

(.~ (D

10 o M664 9 M712

I I

o

X Z2 + 478

o

La

Ce I~r

N? Sm iu Gd l:b13y Ho iF'l"mib

1

. . . . . . . . , i l i . . . . Th La Nd Zr Eu Gd Dy Er Lu Nb Ce Sm Hf T~ Tb Ho Yb ,

(b)

,

,

,

l

,

,

,

|

9

i

,

AI20 3/TiO2

++ o

%o

o oo o(9O~

:

:

:

:

o

+

oOo o o o o~#o

0%'o 0"8 f9 o

@

oo

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I

AI203 (wt%) 0

0

;-++ Ir

Gdn 0

O+~o

8

,

,

% ~#~oo

, ** +++~.++~.+ ~++~++++

i

(Gd/Yb) n

+

+

o

* + ~o.Oo@~ +4.

0

0 0

oOoo o~ o ~ o

F

o

I

o

t

o

I

I

I

I

I

I

no2(wt%)

+{+ -it+

+

+

0.r

I

La n

+

+

+

++ +

+ + 0.5 o o ~o#oo~'# oa~,@~p " +

C~o # ~'~+~ O+ Z++o

0 0 u 0 (:~

0

o

Fig. 4.3.2-1.

"

io

' ~o ' MgO (wt %)

;o

1

~176176176~go% o o I

'

210

I

310

I

MgO (wt

%)

I

40

I

50

l~u-

294


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


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-

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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

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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.


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-

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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.

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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.

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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,

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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-

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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 ~ ,,,~

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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.

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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

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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).


310

(a) 1600

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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


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


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).


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


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

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


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


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


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


320

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Pseudomorphs after gypsum Branching and hemispherical stromatolites Xenoliths

Early mafic dykes

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-


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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


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.


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


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


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


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).

328


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-


329

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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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


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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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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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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

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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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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


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


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-


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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).

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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


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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.

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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,


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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

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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

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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


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


Characteristics and petrographic features


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)


Characteristics and petrographic fcatures


(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 $

350


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)

352


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.


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


(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


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-


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.

358


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.


359

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-

360

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

362


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


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,

364


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 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


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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).


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).


367

Fig. 5.2-2. The Dimroth-Ohmoto model for the evolution of atmospheric chemistry (a) and ocean chemistry (b).

368


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-


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

370


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)


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

372


Fig. 5.2-3. Carbon isotope records of carbonates (a) (Shield and Veizer, 2002) and organic matter (b) (Pavlov et al., 2001b; Yamaguchi, 2002).


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.

374


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


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).

376


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)


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

378


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


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

380

Chapter 5: Evolution of the Hydrosphereand Atmosphere

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


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.

382


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


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.

384


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.


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

386


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


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


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.

390


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


391

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).

392


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-


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

394


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


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


~ ,.,~


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


-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.


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.


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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

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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-

5.4. Precambrian hvn-Formation

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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

5.5. Precambrian Sulphur Isotope Record

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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


423

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-

424


S042- - .....................................

8349

i \ i BSR

~,.. | | |

SQ2S O 4 2-

| l

IA2 ......

H2S~ S ~

H2S"-I"S ~

al = BSR (2CH20 + S042- ~ " H2S + 2HC03-), 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


425

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


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


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 ,

(b)

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10 15 20 25 30 35 40 45

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0

-5

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10 15 20 25 30 35 40 45

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10 15 20 25 30 35 40 45

,

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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

900

0)

0 800 -

0

700 -

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600 -

500

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~34S 1%0) Fig. 5.5-5. 6345 variation in CAS and interbedded gypsum, Society Cliffs Formation (c. 1.2 Ga), Bylot Supergroup, northern Baffin and Bylot islands, Canada. Details are available in Kah et al. (2001, submitted).

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


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-

436


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;

5.5. PrecambrianSulphur Isotope Record

437

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

5.6. Aftermath of the "Snowball Earth" Hypothesis

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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

444


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


445

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

446


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


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).

450


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.

452


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 40~ were confined to middle and high southern palaeolatitudes (Crowley et al., 1989; Gibbs et al., 2002). In present low latitudes ( 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


456

High-latitude glaciation

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0 10 20 30 40 50 60 70 80 90 Obliquity ~ (o)

3

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Relative mean annual insolation

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 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).

594

Chapter 7: Sedimentation ThroughTime

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).

596

Chapter 7: Sedimentation Through Time

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

597

7.1. Introduction

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

598


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

7. I. Introduction

599

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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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

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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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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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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.

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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-

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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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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

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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.

634


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).


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).

636


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).


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

638


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


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


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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9

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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


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

642


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

7. 6. Sedimentary Dynamics of Precambrian Aeolianites

643

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

644


7.6. Sedimentary Dynamics t~'Precambrian Aeolianites

645

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).

7.6. Sedimentary Dynamics of Precambrian Aeolianites

647

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

648


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.


649

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.

650


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


651

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-

652


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).


653

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.

654


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),


655

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

656


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


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

I~ I I I

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


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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

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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.

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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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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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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

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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

672


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

7.10. Microbial M a t Features in Sandstones

673

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

674


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).

7.11. S e d i m e n t a t i o n R a t e s

675

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 ( , ,-, 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


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


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


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


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


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,


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


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


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.


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


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).


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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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


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


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-


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,

734


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

736


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

738


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.


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 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-

744


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,

746


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,

748


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

750


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

752


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-

754


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

9.6. Evolution of Precambrian Life and Bio-Geology

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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).

9. 7. Sedimentation Regimes Through Time

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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

9.8. Sequence Stratigraphy through Time

761

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).


762

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-

9. 9. Tempos and Events in Precambrian Time

763

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.

764

Fig. 9.9-2.



765

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.



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.

768


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.


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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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


The Precambrian Earth: Tempos and Events (Developments in Precambrian Geology)

The Precambrian Earth: tempos and events

Archean greenstone belts (Developments in Precambrian geology)

Principles of Precambrian Geology

Precambrian Geology of the USSR

Iron Formation: Facts and Problems (Developments in Precambrian geology)

Earth's Oldest Rocks, Volume 15 (Developments in Precambrian Geology)

Archaean Geochemistry: Symposium Proceedings, 1977 (Developments in Precambrian geology)

Precambrian Ophiolites and Related Rocks, Volume 13 (Developments in Precambrian Geology)

Precambrian Ore Deposits of the East European and Siberian Cratons (Developments in Economic Geology)

Developments and interactions of the Precambrian atmosphere, lithosphere, and biosphere

Precambrian of the Southern Hemisphere

Precambrian plate tectonics

The Geology of South Australia (Bulletin) Volume 1 The Precambrian

Precambrian Continental Crust and its Economic Resources

Precambrian continental crust and its economic resources

Precambrian Continental Crust and its Economic Resources

Carbonate sedimentation and diagenesis in the evolving Precambrian world

Developments in Precambrian Geology, Volume 16: Neoproterozoic-Cambrian Tectonics, Global Change and Evolution (A Focus on South Western Gondwana)

Early Precambrian Processes (Geological Society Special Publication)

Precambrian geology of Finland: key to the evolution of the Fennoscandian shield

Precambrian crustal evolution in the North Atlantic Region

Physical Geology: Exploring the Earth

Petroleum Geology (Developments in Petroleum Science)

Extinction events in earth history: proceedings of the Project 216, Global Biological Events in Earth History

Precambrian sedimentary environments: a modern approach to ancient depositional systems

Geochronology, Dating, and Precambrian Time: The Beginning of the World As We Know It (The Geologic History of Earth)

Bentonites: Geology, Mineralogy, Properties and Uses (Developments in sedimentology)

Tempos Líquidos

TEMPOS LIQUIDOS

Terminal Eocene Events (Developments in Palaeontology and Stratigraphy)

The Precambrian Earth: Tempos and Events (Developments in Precambrian Geology)