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

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Developments in Precambrian Geology, 13

PRECAMBRIAN OPHIOLITES AND RELATED ROCKS

DEVELOPMENTS IN PRECAMBRIAN GEOLOGY Advisory Editor Kent Condie Further titles in this series 1. B.F. WINDLEY 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. KRÖNER (Editor) Precambrian Plate Tectonics 5. Y.P. MEL’NIK Precambrian Banded Iron-formations. Physicochemical Conditions of Formation 6. A.F. TRENDALL and R.C. MORRIS (Editors) Iron-Formation: Facts and Problems 7. B. NAGY, R. WEBER, J.C. GUERRERO and M. SCHIDLOWSKI (Editors) Developments and Interactions of the Precambrian Atmosphere, Lithosphere and Biosphere 8. S.M. NAQVI (Editor) Precambrian Continental Crust and Its Economic Resources 9. D.V. RUNDQVIST and F.P. MITROFANOV (Editors) Precambrian Geology of the USSR 10. K.C. CONDIE (Editor) Proterozoic Crustal Evolution 11. K.C. CONDIE (Editor) Archean Crustal Evolution 12. P.G. ERIKSSON, W. ALTERMANN, D.R. NELSON, W.U. MUELLER and O. CATUNEANU (Editors) The Precambrian Earth: Tempos and Events

Developments in Precambrian Geology, 13

PRECAMBRIAN OPHIOLITES AND RELATED ROCKS Edited by

T.M. KUSKY Department of Earth and Atmospheric Sciences Saint Louis University St. Louis, MO 63103, USA

2004

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v

CONTRIBUTING AUTHORS

A.M. AL-SALEH Geology Department, King Saud University, PO Box 2009, Riyadh, Saudi Arabia I.I. BABARINA Geological Institute of the RAS, Pyzhevsky per. 7, Moscow 109017, Russia E.V. BIBIKOVA Vernadsky Institute of Geochemistry and Analytical Chemistry of RAS, Kosygin st. 19, Moscow 17975, Russia M.M. BOGINA Institute of Ore Deposits, Petrography, Geochemistry and Mineralogy of RAS, Staromonetny per. 35, Moscow 109017, Russia R.W. CARLSON Department of Terrestrial Magnetism, Carnegie Institute of Washington, Washington, DC, USA P.L. CORCORAN Department of Geology, University of Western Ontario, Canada J. DANN 90 Old Stow Road, Concord, MA 01742, USA ([email protected]) M. DE WIT CIGCES, Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa ([email protected], http://www.uct.ac.za/depts/cigces) J. ENCARNACIÓN Department of Earth and Atmospheric Sciences, Saint Louis University, St. Louis, MO 63103, USA (jpe@eas. slu.edu, http://www.eas.slu.edu/people/jpencarnacion/jpeslu.html) F. JUN Department of Geology, Peking University, Beijing 100871, China R. GANLEY Department of Earth and Atmospheric Sciences, Saint Louis University, 3507 Laclede Ave., St. Louis, MO 63103, USA A. GLASS Earth Sciences Department, Cardiff University, PO Box 914, Main Building, Park Place, Cardiff, South Glamorgan CF1434E, UK ([email protected]) T.L. GROVE Department of Earth, Atmospheric, and Planetary Sciences, MIT, Building 54-1224, Cambridge, MA 02139, USA ([email protected]) A. HOFMANN School of Geosciences, University of the Witwatersrand, P. Bag 3, Wits 2050, South Africa

vi

Contributing authors

X.N. HUANG Department of Geology, Peking University, Beijing 100871, China I.M. HUSSEIN Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, Germany R. HUSON Department of Earth and Atmospheric Sciences, Saint Louis University, 3507 Laclede Ave., St. Louis, MO 63103, USA ([email protected], http://www.eas.slu.edu/people/students/rlhuson/index.html) P.R. JOHNSON Saudi Geological Survey, PO Box 54141, Jiddah 21514, Saudi Arabia ([email protected]) F.H. KATTAN Saudi Geological Survey, PO Box 54141, Jiddah 21514, Saudi Arabia R. KERRICH Department of Geological Sciences, University of Saskatchewan, SK, Canada S7N 5E2 (robert.kerrich@ usask.ca) A.N. KONILOV Geological Institute of the RAS, Pyzhevsky per. 7, Moscow 109017, Russia A. KONTINEN Geological Survey of Finland, PO Box 1237, FIN-70211 Kuopio, Finland A. KRÖNER Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, Germany ([email protected], http://www.uni-mainz.de/∼kroener/) K.A. KRYLOV Department of Geological and Environmental Sciences, Stanford University, CA 94305-2115, USA (kirka@geo. tv-sign.ru) T.M. KUSKY P.C. Reinert Endowed Chair of Natural Sciences, Department of Earth and Atmospheric Sciences, Saint Louis University, St. Louis, MO 63103, USA ([email protected], http://www.eas.slu.edu/programs/geograd.html# kusky) J.H. LI Department of Geology, Peking University, Beijing 100871, China ([email protected]) J. LYTWYN Department of Geosciences, University of Houston, Houston, TX 77058, USA W.U. MUELLER Department of Geology, University of Quebec at Chicoutimi, Canada B.A. NATAL’IN ˙ Maden Fakültesi, Jeoloji Bölümü, Ayazaˇga 34469, Istanbul, ˙ ITÜ Turkey ([email protected]) X.L. NIU Department of Geology, Peking University, Beijing 100871, China S.W. PARMAN Department of Earth, Atmospheric, and Planetary Sciences, MIT, Building 54-1224, Cambridge, MA 02139, USA ([email protected])

Contributing authors

vii

P. PELTONEN Geological Survey of Finland, PO Box 96, FIN-02151 Espoo, Finland ([email protected]) J. PFÄNDER Institut für Mineralogie, Universität Münster, Corrensstr. 24, D-48149 Münster, Germany (pfaender@ mpch-mainz.mpg.de) A. POLAT Department of Earth Sciences, University of Windsor, Windsor, Ontario, Canada N9B 3P4 ([email protected]) I.S. PUCHTEL Department of the Geophysical Sciences, The University of Chicago, 5734 South Ellis Avenue, Chicago, IL 60637, USA ([email protected], http://geosci.uchicago.edu/∼ ipuchtel/) T. RAHARIMAHEFA Department of Earth and Atmospheric Sciences, Saint Louis University, 3507 Laclede Ave., St. Louis, MO 63103, USA T. REISCHMANN Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, Germany ([email protected]) A.V. SAMSONOV Institute of Ore Deposits, Petrography, Geochemistry and Mineralogy of RAS, Staromonetny per. 35, Moscow 109017, Russia A.M.C. SENGÖR ¸ ˙ Maden Fakültesi, Jeoloji Bölümü, Ayazaˇga 34469, Istanbul, ˙ ITÜ Turkey ([email protected]) A.A. SHCHIPANSKY Geological Institute of the RAS, Pyzhevsky per. 7, Moscow 109017, Russia ([email protected]) A.I. SLABUNOV Karelian Research Center of RAS, Institute of Geology, Pushkinskaya st. 11, Petrozavodsk 185610, Karelia, Russia R.J. STERN Geosciences Department, University of Texas at Dallas, Box 830688, 2601 N. Floyd Rd., Richardson, TX 75083-0688, USA (http://www.utdallas.edu/dept/geoscience/) B. YIBAS Pulles Howard and de Lange, Environmental and Water Quality Management, PO Box 861, Auckland Park 2006, South Africa ([email protected]) C. ZHENG School of Earth and Space Sciences, Peking University, Beijing 100871, China

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ix

CONTENTS

Contributing authors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

v

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

xi

Dedication . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xiii Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . T.M. Kusky

1

PROTEROZOIC OPHIOLITES AND RELATED ROCKS Chapter 1.

The Jormua Ophiolite: A Mafic-Ultramafic Complex from an Ancient Ocean-Continent Transition Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P. Peltonen and A. Kontinen

35

Chapter 2.

The 1.73 Ga Payson Ophiolite, Arizona, USA . . . . . . . . . . . . . . . . . . . . . . . . . . J.C. Dann

73

Chapter 3.

Neoproterozoic Ophiolites of the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . Robert J. Stern, Peter R. Johnson, Alfred Kröner and Bisrat Yibas

95

Chapter 4.

Neoproterozoic Ophiolites in the Arabian Shield: Field Relations and Structure . . . . . . . 129 Peter R. Johnson, Fayek H. Kattan and Ahmed M. Al-Saleh

Chapter 5.

The Wadi Onib Mafic-Ultramafic Complex: A Neoproterozoic Supra-Subduction Zone Ophiolite in the Northern Red Sea Hills of the Sudan . . . . . . . . . . . . . . . . . . . . . . 163 I.M. Hussein, A. Kröner and T. Reischmann

Chapter 6.

Tectono-Magmatic Evolution, Age and Emplacement of the Agardagh Tes-Chem Ophiolite in Tuva, Central Asia: Crustal Growth by Island Arc Accretion . . . . . . . . . . . . . . . . . 207 J.A. Pfänder and A. Kröner

ARCHEAN OPHIOLITES AND RELATED ROCKS Chapter 7.

Origin and Emplacement of Archean Ophiolites of the Central Orogenic Belt, North China Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 223 Timothy M. Kusky, Jainghai Li, Adam Glass and X.N. Huang

Chapter 8.

Re-Os Isotope Chemistry and Geochronology of Chromite from Mantle Podiform Chromites from the Zunhua Ophiolitic Mélange Belt, N. China: Correlation with the Dongwanzi Ophiolite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 275 T.M. Kusky, J.H. Li, T. Raharimahefa and R.W. Carlson

Chapter 9.

Geochemical and Petrographic Characteristics of the Central Belt of the Archean Dongwanzi Ophiolite Complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 283 R. Huson, T.M. Kusky and J.H. Li

x

Contents

Chapter 10. Microstructures of the Zunhua 2.50 Ga Podiform Chromite, North China Craton and Implications for the Deformation and Rheology of the Archean Oceanic Lithospheric Mantle 321 Xiongnan Huang, Jianghai Li, T.M. Kusky and Zheng Chen Chapter 11. Neoarchean Massive Sulfide of Wutai Mountain, North China: A Black Smoker Chimney and Mound Complex Within 2.50 Ga-Old Oceanic Crust . . . . . . . . . . . . . . . . . . . . 339 Jianghai Li, Tim Kusky, Xianglong Niu, Feng Jun and Ali Polat Chapter 12. Inferred Ophiolites in the Archean Slave Craton . . . . . . . . . . . . . . . . . . . . . . . . . 363 P.L. Corcoran, W.U. Mueller and T.M. Kusky Chapter 13. 3.0 Ga Olondo Greenstone Belt in the Aldan Shield, E. Siberia . . . . . . . . . . . . . . . . . 405 Igor S. Puchtel Chapter 14. 2.8 Ga Boninite-Hosting Partial Suprasubduction Zone Ophiolite Sequences from the North Karelian Greenstone Belt, NE Baltic Shield, Russia . . . . . . . . . . . . . . . . . . . . . . . 425 A.A. Shchipansky, A.V. Samsonov, E.V. Bibikova, I.I. Babarina, A.N. Konilov, K.A. Krylov, A.I. Slabunov and M.M. Bogina Chapter 15. The Belingwe Greenstone Belt: Ensialic or Oceanic? . . . . . . . . . . . . . . . . . . . . . . 487 Axel Hofmann and Tim Kusky MODELS FOR THE EVOLUTION OF OCEANIC CRUST WITH TIME Chapter 16. Petrology and Geochemistry of Barberton Komatiites and Basaltic Komatiites: Evidence of Archean Fore-Arc Magmatism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 539 S.W. Parman and T.L. Grove Chapter 17. Precambrian Arc Associations: Boninites, Adakites, Magnesian Andesites, and Nb-Enriched Basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 567 A. Polat and R. Kerrich Chapter 18. Archean Greenstone Belts Do Contain Fragments of Ophiolites . . . . . . . . . . . . . . . . 599 Maarten J. de Wit ANALOGS TO PRECAMBRIAN OPHIOLITES Chapter 19. Northern Philippine Ophiolites: Modern Analogues to Precambrian Ophiolites? . . . . . . . 615 John Encarnación Chapter 20. The Resurrection Peninsula Ophiolite, Mélange and Accreted Flysch Belts of Southern Alaska as an Analog for Trench-Forearc Systems in Precambrian Orogens . . . . . . . . . . 627 Timothy M. Kusky, Rose Ganley, Jennifer Lytwyn and Ali Polat Chapter 21. Phanerozoic Analogues of Archaean Oceanic Basement Fragments: Altaid Ophiolites and Ophirags . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 675 A.M.C. S¸ engör and B.A. Natal’in Chapter 22. Epilogue: What if Anything Have We Learned About Precambrian Ophiolites and Early Earth Processes? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 727 Timothy M. Kusky Subject Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 739

xi

PREFACE

Recent developments have shown that many full and partial ophiolites are preserved in Precambrian cratons. This book provides a comprehensive description and discussion of the field aspects, geochemistry, geochronology, and structure of the best of these ophiolites. The book also presents syntheses of the characteristics of ophiolites of different ages, and presents an analysis of what the characteristics of these ophiolites mean for the thermal and chemical evolution of the earth. This book emphasizes new studies of Precambrian Geology that have documented ophiolites, ophiolitic fragments, and ophiolitic melanges in many Precambrian terranes. Each chapter focuses on individual Precambrian ophiolites or regions with numerous Precambrian ophiolites, and covers field aspects, petrology, geochemistry, geochronology, and other descriptive aspects of these ophiolites, and also delves into more theoretical and speculative aspects about the interpretation of the significance of these ancient ophiolites. The first section of the book focuses on Precambrian ophiolites and associated rock suites in Proterozoic terranes. Each chapter emphasizes field relationships, structural associations, geochronology, and chemistry to present a well-rounded perspective on some of the best documented Proterozoic ophiolites in the world. These descriptions serve as a useful standard of the characteristics of the preserved Proterozoic oceanic realm that can serve also for comparison with younger and older ophiolitic sequences. The second part of the book focuses on Archean ophiolites. The first several chapters examine different aspects of the Dongwanzi ophiolite and Zunhua ophiolitic melange in the North China Craton, and correlative ophiolitic terranes in the ca. 2.5 Ga Central Orogenic Belt of the craton. The next chapter examines the controversy surrounding interpretation of several greenstone belts in northern Canada’s Slave Province as ophiolites. The following chapters document Archean ophiolites from other terranes around the world, including Middle and Early Archean belts in Asia and Africa. The final section examines several younger orogens, including the Altaids, southern Alaska convergent margin, and the Philippines, as possible modern analogs to Precambrian, and especially Archean ophiolites. A synopsis of the main points of the papers in the book is presented at the end of the Introduction, and discussed again in a concluding chapter to highlight what has been learned about Precambrian oceanic spreading systems from this compilation of descriptions of Precambrian ophiolites. A table of some of the diagnostic, characteristic, typical, and rare aspects of ophiolites of all ages is presented in order to help determine if tectonically deformed and metamorphosed sequences in ancient mountain belts may be considered ophiolites. This comparative approach is important in that it enables users to more realistically characterize an allochthonous mafic/ultramafic rock sequences as ophiolitic than some other arbitrary classification schemes that have proposed requiring three or four

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Preface

of the Penrose-style ophiolitic units to be present in Precambrian sequences for a specific rock sequence to be considered ophiolitic. Once these tectonic fragments are recognized as remnants of ancient ocean floor, great progress may be made in understanding early Earth history.

ACKNOWLEDGEMENTS First and foremost I thank my wife Carolyn, and children Shoshana and Daniel, for their patience and understanding while writing and compiling this volume. I also extend my warm thanks to the students who have helped me with editing and formatting the text, particularly Rose Ganley, Angie Bond, Rusunoko Made, and Elisabet Head. Special thanks also go to Series Editor Kent Condie for inviting me to write and compile this volume, and to the editorial staff at Elsevier including Patricia Masser and Friso Venstra. I also wish to thank the many people who have helped shape my opinions about the Precambrian and those who have reviewed chapters in this volume. Outstanding among these groups are especially Kevin Burke, Bill Kidd, John Dewey, Paul Hoffman, Celail Sengör, Win Means, Akiho Miyashiro, Declan De Paor, Bruce Marsh, Dwight Bradley, Li Jianghai, Brian Windley, Alfred Kröner, Bob Stern, Bob Tucker, Maarten de Wit, Sam Bowring, Rich Goldfarb, and Peter Hudleston. Finally, I thank the many colleagues who have worked in the field with me, both on greenstone belts of Canada, USA, Africa, Australia, China, and Arabia, and on younger ophiolites of Newfoundland, Alaska, Oman, and the Appalachians. Timothy M. Kusky

xiii

DEDICATION

This book is dedicated to Kevin Burke, whose unfaltering commitment to understanding processes of plate tectonics on the Earth through time has inspired many geologists to critically examine some of the planet’s oldest rock sequences, searching for records of ancient plate interactions. Without Kevin’s knowledge, teaching, humor, and inspiration, many careers would have taken different routes and scientific discoveries and models such as those presented in this book may never have materialized.

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1

INTRODUCTION T.M. KUSKY

Understanding the early history of the Earth and how the planet developed from a planetary nebula to its present, life-sustaining state is one of the most fundamental problems in Earth Sciences, and one that has occupied the thoughts of scholars and sages for centuries. To approach these questions it is necessary to synergize data from a variety of sources to estimate how plate tectonics, which is presently the surface expression of planetary heat loss, has evolved from a period of higher heat flow from the early Earth to its present state. This knowledge leads to a better understanding of how the surface and interior of the planet have evolved with time, and how previous interactions of the lithosphere, atmosphere, hydrosphere, and biosphere have been driven by the fundamental heat loss from the interior of the earth. These broad geodynamic goals require the integration of data from many different fields, including structural geology and tectonics, geophysics, petrology and geochemistry, sedimentology, and biology. It also involves several different scales of observation, from global geophysical data sets, to regional synthesis, through the outcrop scale, to the microscopic and molecular levels, which yield clues about the boundary conditions of formation and deformation. Understanding the early history of the Earth is thus a multi-disciplinary, multi-scale problem. To make headway in understanding how the early Earth operated, and how it may have differed or been similar to today’s Earth, it is necessary to compare studies of important Precambrian geological provinces with possible modern analogs, as well as other planets. Thus, in this book there are a number of studies that make comparisons between Precambrian and younger geological settings, and other chapters that describe active and Phanerozoic tectonic settings as potential analogs to Archean terranes, greenstone belts, and ophiolites.

1. GENERAL CHARACTERISTICS OF OPHIOLITES Ophiolites are a distinctive association of rocks interpreted to form in a variety of plate tectonic settings, including oceanic spreading centers, back-arc basins, forearcs, arcs, and other extensional magmatic settings including those in association with plumes (Moores, 1982, 2002; Gass et al., 1984; Lippard et al., 1986; Nicolas, 1989; Parson et al., 1992; Peters et al., 1991; de Wit and Ashwal, 1997; Dilek et al., 2000). A complete ophiolite grades downward from pelagic sediments into a mafic volcanic complex comprised mostly of pillow basalts, underlain by a sheeted dike complex. These are underlain by

2

Introduction

gabbros exhibiting cumulus textures, then tectonized peridotite, resting above a thrust fault that marks the contact with underlying rock sequences (Dewey and Bird, 1971; Dewey, 1977). The term “ophiolite” refers to this distinctive rock association (e.g., Sylvester et al., 1997), although many workers interpret the term to mean structurally emplaced oceanic lithosphere rocks formed exclusively at mid-ocean ridges. Many ophiolites are altered to serpentinite, chlorite, albite, and epidote rich rocks, possibly by hydrothermal sea floor metamorphism. In the 1960s and 1970s much research was aimed at defining a type ophiolite succession, which became known as the Penrose-type of ophiolite (Anonymous, 1972). More recent research has revealed that the variations between individual ophiolites are as significant as any broad similarities between them (e.g., Casey et al., 1981; Moores, 2002; Karson, 2001; Dilek et al., 2000; Dilek and Newcomb, 2003). A classic Penrose-type ophiolite is typically 5–15 km thick, and if complete, consists of the following sequence from base to top, with a fault marking the base of the ophiolite. The base of most ophiolites consists of harzburgite, consisting of olivine + orthopyroxene (± chromite), often forming strongly deformed or transposed compositional layering, forming harzburgite tectonite. The lowest unit in some ophiolites is lherzolite, consisting of olivine + clinopyroxene + orthopyroxene, generally interpreted to be fertile, undepleted mantle. In some ophiolites, harzburgite overlies lherzolite. The harzburgite is generally interpreted to be the depleted mantle from which overlying mafic rocks were derived, and the deformation is related to the overlying lithospheric sequence flowing away from the ridge along a shear zone within the harzburgite. The harzburgite sequence may be more than 10 km thick in some ophiolites, such as the Semail ophiolite in Oman, and the Bay of Island ophiolite in Newfoundland. See Fig. 1. Resting above the harzburgite is a group of rocks that were crystallized from magma derived by partial melting of the harzburgite. The lowest unit of these crustal rocks includes crystal cumulates of pyroxene and olivine, forming distinctive layers of pyroxenite, dunite, and other olivine + clinopyroxene + orthopyroxene peridotites including wehrlite, websterite, and pods of chromite + olivine. The boundary between these rocks (derived by partial melting and crystal fractionation) and those below from which melts were extracted is one of the most fundamental boundaries in oceanic lithosphere and defines the petrologic Moho, or base of the crust. In this case, the Moho is a chemical boundary, without a sharp seismic discontinuity. A seismic discontinuity occurs about half a kilometer higher than the chemical Moho in ophiolites. The layered ultramafic cumulates grade upwards into a transition zone of interlayered pyroxenite and plagioclase-rich cumulates, then into an approximately 1 km thick unit of strongly layered gabbro. Individual layers within this thin unit may include gabbro, pyroxenite, and anorthosite. The layered gabbro is succeeded upward by several to 5 km of isotropic gabbro, which is generally structureless but may have a faint layering. The layers within the isotropic gabbro in some ophiolites define a curving trajectory, interpreted to represent crystallization along the walls of a paleomagma chamber. The upper part of the gabbro may contain many xenoliths of diabase, pods of trondhjemite (plagioclase plus quartz), and may be cut by diabase dikes.

1. General Characteristics of Ophiolites

3

Fig. 1. Schematic columnar sections of representative oceanic crust and corresponding ophiolite types (modified after Moores, 2002). (A) Complete ophiolite sequence according to the “Penrose Conference definition” (Anonymous, 1972), characteristic of a magma-rich (generally fast spreading) ridge. (B) Faulted, incomplete sequence characteristic of a magma-starved spreading where tectonic processes dominate, designated a “Hess-type”ophiolite from Hess’s (1962) characterization of oceanic crust as serpentinized peridotite. (C) Complex composite section of oceanic island-arc sequences developed on oceanic crust. Designated a “Smartville-type” for the Smartville Complex in the northwest Sierra Nevada, California (e.g., Dilek et al., 1991). (D) Possible hotspot (oceanic plateau) section of oceanic crust. (E) A new, ”transitional type” of ophiolite, recognized where oceanic crust forms over extended continental crust, with an example from the Tihama Asir area of Saudi Arabia on the Red Sea. See text for discussion.

4

Introduction

The next highest unit in a complete, Penrose-style ophiolite is typically a sheeted dike complex, consisting of a 0.5–2 km thick complex of diabasic, gabbroic, to silicic dikes that show mutually intrusive relationships with the underlying gabbro. In ideal cases, each diabase dike intrudes into the center of the previously intruded dike, forming a sequence of dikes that have chilled margins developed only on one side. These dikes are said to exhibit one-way chilling. In most ophiolites that are not severely deformed or metamorphosed, examples of one way chilling may be found, but statistically the one-way chilling may only show directional preference in 50–60% of cases. The sheeted dikes represent magma conduits that fed basaltic flows at the surface. These flows are typically pillowed, with lobes and tubes of basalt forming bulbous shapes distinctive of underwater basaltic volcanism. The pillow basalt section is typically 0.5–1 km thick. Interstices between the pillows may be filled with chert, and sulfide minerals are common. Many ophiolites are overlain by deep-sea sediments, including chert, red clay, in some cases carbonates, or sulfide layers. Many variations are possible, depending on tectonic setting (e.g., conglomerates may form in some settings) and age (e.g., siliceous biogenic oozes and limestones would not form in Archean ophiolites, before the life forms that contribute their bodies developed on the Earth).

2. VARIATIONS BETWEEN OPHIOLITES AND OCEANIC CRUST FORMED UNDER DIFFERENT CONDITIONS Research in ophiolite studies and in situ oceanic crust in the past decade has revealed a much greater diversity in the abundance and sequence of rock types present than originally known when the Penrose-definition of an ophiolite was published (e.g., Moores, 2002; Karson, 2001; Dilek et al., 2000; Dilek and Newcomb, 2003). For instance, it is now thought that Penrose-style ophiolites form at fast spreading ridges such as the East Pacific Rise, whereas slow-spreading ridges like the Atlantic may have massive serpentinite overlying harzburgite tectonite, in turn faulted against pillows, dikes and gabbros. In some of these slow-spreading systems, extension is proceeding faster than magma can upwell to replace the lost volume of oceanic crust (e.g., Dick et al., 2003). Ophiolites produced in arc environments, either the main arc, forearc, or back arc, typically have thick mafic cumulate sections, interstratified volcanoclastic, pelagic, or hemipelagic sediments, and thick lava sequences. Many arc-type ophiolites are intruded by plutons with mafic to silicic compositions, that fed upper parts of the arc sequence. Oceanic crust produced at hot spots has particular relevance to the Archean, since some models for the Archean suggest that the mantle may have been slightly hotter in earlier times, and possibly analogous to hot spots. However, hot spots are not much hotter than surrounding mantle but, but simply produce greater amounts of magma than surrounding mantle. Places like Iceland, where a hot spot is superimposed on a mid-ocean ridge may be particularly analogous to Precambrian oceanic spreading environments, where both upwelling and enhanced melting processes are occurring. Oceanic plateaus, with thick crustal sections, may have been the norm for oceanic crust in the Archean, so it is difficult, when

2. Variations between Ophiolites and Oceanic Crust Formed under Different Conditions

5

assessing oceanic crust formed in early times, to differentiate between so-called normal oceanic crust and plateau-type crust. Oceanic plateaus have crustal sections that are similar to Penrose-style ophiolites, but may reach or exceed 10–15 km in thickness. It may be best, during our present early stages of studies of ancient oceanic crust, to use the standard of the present when describing Archean sequences. In this way, any similarities or differences between the present and Precambrian can be better-assessed than approaches that use a moving target of reference, such as the “alternative Earth” model of Hamilton (2003), or the ad-hoc models of Bickle et al. (1994). Several of the ophiolites described in this volume appear to have formed within the transition from rifted continental margins to ocean spreading centers during early stages of ocean opening, then were structurally detached and/or deformed and incorporated into convergent margins during ocean closure. These ophiolites are distinctive from classical Penrose-style ophiolites and others formed in forearc and back arc environments, and we coin the term “transitional ophiolites” for this new class of ophiolite. During early stages of ocean formation, continental crust and lherzolite of the subcontinental mantle is extended forming graben on the surface, and ductile mylonites at depth. Sedimentary basins may form in the graben, and as the extension continues magmatism sometimes affects the rifted margin, either forming volcanic rifted margins, or migrating to a spreading center forming an oceanic spreading center. New asthenospheric mantle upwells along the new ridge, and may intrude beneath the extended continental crust. In some cases, wedges of extended mid-to-lower continental crust overlying mylonitic lherzolitic sub-continental mantle become intruded by numerous dikes and magmas from this new asthenospheric mantle. These dikes then feed a crustal gabbroic magma chamber closer to the surface, which in turn may feed a dike complex and basaltic pillow/massive lava section. If preserved, this unusual sequence forms a “transitional ophiolite”, grading down from subaquatic sediments, to pillow lavas, dikes, sheeted dikes, layered gabbro, dunite and pyroxenite cumulates, then remarkably into stretched, typically mylonitic granitic mylonites, underlain by lherzolite. The lherzolite tectonic may be underlain by harzburgite tectonite or harzburgite. (See Fig. 1.) Examples of this type of transitional ophiolite are found in the Proterozoic Jourma complex, and in some of the Slave Province ophiolites (see papers by Peltonen and Kontinen, 2004, and Corcoran et al., 2004). Modern analogs for such transitional ophiolites are found around the Red Sea, including at Tihama Asir, Saudi Arabia, where a 5–10 Ma old transitional ophiolite has a dike complex overlying layered gabbro, which in turn overlies continental crust. Also, on Egypt’s Zabargad Island, oceanic mantle is exposed, and it is likely that the crustal structure near this region preserves transitional ophiolites as well. The main lesson here is that ophiolites may form in many tectonic settings, from extended continental crust, to mid ocean ridges, to forearcs, arcs, back arcs, to triple junctions along convergent margins. Once we identify a sequence as ophiolitic, then we need to identify the tectonic environment in which it formed before we can make realistic comparisons to younger environments to determine how plate tectonic style has changed with time.

6

Introduction

3. PROCESSES OF OPHIOLITE AND OCEANIC CRUST FORMATION The sequence of rock types described above are a product of specific processes that occurred within the oceanic spreading centers along which the ophiolites formed (see reviews by Anonymous, 1972; Moores, 1982; Nicolas, 1989; Parson et al., 1992; Dilek et al., 2000). Variations in the rock sequence, mineralogy, chemistry, or structure of ophiolites with time may be related to variations in the processes that produced the ophiolites. With higher heat production in the early Earth, it is important to document these variations to determine how the Earth may have lost heat in early times of high heat production. It is not clear if the early mantle responded to the higher heat production by a significant increase in temperature, a change in viscosity and greater ease in convective overturn, or some other process (e.g., Abbott and Mencke, 1990; Arndt et al., 1997; Condie, 1981, 1997b; Grove et al., 1994; Helmstaedt and Shulze, 1989; Martin, 1986, 1993; McKenzie and Bickle, 1988; Nisbet et al., 1993; Parman et al., 1997; Turcotte and Schubert, 2002; Parman and Grove, 2004). Studies of Precambrian ophiolites and related rocks have great potential to unravel the secrets of early heat loss from the planet. As the mantle convects and the asthenosphere upwells beneath mid ocean ridges, mantle pyrolites, harzburgites and lherzolites undergo partial melting of 10–15% in response to the decreasing pressure. The percentage of partial melt may have been different in early times if mantle temperatures were significantly higher. So far, estimates of partial melt fractions estimated from Precambrian ophiolites have not determined whether or not the Precambrian melt fractions were on average greater than, less than, or similar to those of the younger record. The melts derived from the harzburgites rise to form a magma chamber beneath the ridge, forming the crustal section of the oceanic crust. As the magma crystallizes the densest crystals gravitationally settle to the bottom of the magma chamber, forming layers of ultramafic and higher mafic cumulate rocks. Above the cumulates, a gabbroic fossil magma chamber forms, typically with layers defined by varying amounts of pyroxene and feldspar crystals. In many examples the layering in ophiolites has been shown to be parallel to the fossil margins of the magma chamber. An interesting aspect of the magma chamber is that periodically, new magma is injected into the chamber, changing the chemical and physical dynamics. These new magmas are injected during extension of the crust so the magma chamber may effectively expand infinitely if the magma supply is continuous, as in fast spreading ridges. In slow spreading ridges the magma chamber may completely crystallize before new batches of melt are injected. Studies of ancient ophiolites therefore have the potential to estimate relative rates of extension and magma supply, a line of research that has not yet matured in Precambrian ophiolite studies. As extension occurs in the oceanic crust, dikes of magma shoot out of the gabbroic magma chamber, forming a diabasic (fine-grained rapidly cooled magma with the same composition as gabbro) to gabbroic sheeted dike complex. The dikes have a tendency to intrude along the weakest, least crystallized part of the previous dike, which is usually in the center of the last dike to intrude. In this way each dike intrudes the center of the previous dike, forming a sheeted dike complex characterized by dikes that have only one chill margin, most of which face in the same direction. In some Phanerozoic ophiolites, varia-

4. Historical Recognition of Archean Ophiolites

7

tions in the thickness and character of the dikes with depth have been related to temperature changes with depth. The dike complex thus represents a potential indicator of geothermal gradients in ancient ophiolites, but extracting such information from Precambrian ophiolites may be difficult due to the paucity of Precambrian dike complexes, plus the effects of deformation and metamorphism have obscured many original relationships. Many of the dikes reach the surface of the sea floor, where they feed basaltic lava flows. Basaltic lava flows on the sea floor are typically in the form of bulbous pillows that stretch out of magma tubes, forming the distinctive pillow lava section of ophiolites. The top of the pillow lava section is typically quite altered by sea floor metamorphism including mineralized veins that culminate in deposits of black smoker-type hydrothermal vents (e.g., Harper, 1999; Li et al., 2004). Early life forms probably flourished around these deep sea hydrothermal vents, but few studies have focused on searching for early life forms around Precambrian sea floor black smokers (but see Huang et al., 2004). Studies of lava vesicularity, and the nature of interpillow sediments, have the potential to yield clues about the water depth and hence crustal thickness of ancient ophiolites, but few studies have yet explored these potentially critical relationships. The pillow lavas are overlain by sediments deposited on the sea floor. In the Phanerozoic oceans, if the oceanic crust formed above the calcium carbonate compensation depth, the lowermost sediments may be calcareous. These would be succeeded by siliceous oozes, pelagic shales, and other deep water sediments as the sea floor cools, subsides, and moves away from the mid ocean ridge. There has been no comprehensive analysis of the types of sediments expected to be deposited on Precambrian oceanic crust as it moved away from spreading centers, nor how this sequence may have changed systematically with time. A third sequence of sediments may be found on the ophiolites. These would include sediments shed during detachment of the ophiolite from the sea floor basement, and its thrusting (obduction) onto the continental margin. The type of sediments deposited on ophiolites should have been very different in some of the oldest ophiolites that formed in the Precambrian. For instance, in the Proterozoic and especially the Archean, organisms that produce the carbonate and siliceous oozes were not present, as the organisms that produced these sediments had not yet evolved. Thus, study of the sedimentary sequences deposited on Precambrian ophiolites may yield important information about Precambrian sea water and atmospheric conditions, sedimentation processes, and about the development and evolution of life in the early oceans.

4. HISTORICAL RECOGNITION OF ARCHEAN OPHIOLITES Very few complete Phanerozoic-like ophiolite sequences have been recognized in Archean greenstone belts, leading some workers to the conclusion that no Archean ophiolites or oceanic crustal fragments are preserved (e.g., Bickle et al., 1994). However, as emphasized by Sylvester et al. (1997), the original definition of ophiolites (Anonymous, 1972) includes “dismembered”, “partial”, and “metamorphosed” varieties, and there is no justification for new arbitrary definitions that attempt to exclude portions of Archean

8

Introduction

greenstone belts that contain two or more parts of the full ophiolite sequence, especially in structurally complex settings such as found in greenstone belts (e.g., Harper, 1985; de Wit et al., 1987; Kusky 1990, 1991). Similarly, many Proterozoic ophiolites are dismembered, or partial sequences (Kröner, 1985; Berhe, 1990; Dann, 1991). Archean oceanic crust was possibly thicker than Proterozoic and Phanerozoic counterparts, resulting in accretion predominantly of the upper section (basaltic) of oceanic crust (Burke et al., 1976; Hoffman and Ranalli, 1988; Moores, 1986; Burke, 1995; Kusky and Vearncombe, 1997). The crustal thickness of Archean oceanic crust may in fact have resembled modern oceanic plateaux (e.g., Sleep and Windley, 1982; Kusky and Kidd, 1992). If this were the case, complete Phanerozoic-like MORB-type ophiolite sequences would have been very unlikely to be accreted or obducted during Archean orogenies. In contrast, only the upper, pillow lava-dominated sections would likely be accreted (Kusky and Polat, 1999). Portions of several Archean greenstone belts have been interpreted to contain dismembered or partial ophiolites. Accretion of MORB-type ophiolites has been proposed as a mechanism of continental growth in a number of Archean, Proterozoic, and Phanerozoic orogens (Kusky and Polat, 1999). It is worthwhile to investigate these claims to better understand the crustal structure and tectonic setting in which these Archean ophiolites formed. Several suspected Archean ophiolites have been particularly well-documented. One of the most disputed is the circa 3.5 Ga Jamestown ophiolite in the Barberton greenstone belt of the Kaapvaal craton (Brandl and de Wit, 1997). de Wit et al. (1987) describe a 3 km thick tectonomagmatic sequence including a basal peridotite tectonite unit with chemical and textural affinities to Alpine-type peridotites, overlain by an intrusiveextrusive igneous sequence, and capped by a chert-shale sequence. This partial ophiolite is pervasively hydrothermally altered and shows chemical evidence for interaction with sea water with high heat and fluid fluxes (de Wit et al., 1990). SiO2 and MgO metasomatism and black-smoker like mineralization is common, with some possible hydrothermal vents traceable into banded iron formations, and subaerial mudpool structures. These features led de Wit et al. (1982, 1992) to suggest that this ophiolite formed in a shallow sea, and was locally subaerial, analogous to the Reykjanges ridge of Iceland. In this sense, Archean oceanic lithosphere may have looked very much like younger oceanic plateaux lithosphere. Maarten de Wit (2004) presents an updated description and interpretation of the Barberton belt and Jamestown ophiolite in this volume, incorporating years of additional mapping. Several partial or dismembered ophiolites have been described from the Slave Province, although these too have been disputed (e.g., King and Helmstaedt, 1997). From the Point Lake greenstone belt in the central Slave Province, Kusky (1991) described a fault-bounded sequence grading downwards from shales and chemical sediments (umbers) into several kilometers of pillow lavas intruded by dikes and sills, locally into multiple dike/sill complexes, then into isotropic and cumulate-textured layered gabbro. The base of this partial Archean ophiolite is marked by a 1 km thick shear zone composed predominantly of mafic and ultramafic mylonites, with less-deformed domains including dunite, websterite, wehrlite, serpentinite, and anorthosite. By using down-plunge projections and sectionbalancing techniques, Kusky (1991) estimated that the shear zone at the base of this ophi-

4. Historical Recognition of Archean Ophiolites

9

olite accommodated a minimum of 69 km of slip. Although this still allows the ophiolite to have formed at or near extended older continental crust that forms the Anton terrane to the west (Kusky, 1989), the actual amount of transport was probably much greater. Synorogenic conglomerates and sandstones were deposited in several small foredeep basins, and are interbedded with mugearitic lavas (and associated dikes), all deposited/intruded in a foreland basin setting. Kusky (1990) suggested that portions of the Cameron and Beaulieu River greenstone belts of the southern Slave Province contain ophiolitic components. The belts are cut by numerous layer-parallel shear zones, but some sections are composed mostly of tholeiitic pillow basalts, others contain approximately equal quantities of pillows and dikes, and a few sections consist of nearly 100% mafic dikes. The bases of these greenstone belts are marked by up to 500 m thick shear zones (locally containing mélanges), with tectonic blocks of gabbro, mafic volcanics, peridotite, and slivers of the underlying quartzofeldspathic gneiss (with extensive mafic dike complexes) and its autochthonous cover. Original relationships between dikes in the basement complex and dikes in the basal parts of the greenstone belts have not been established, but older-generation mafic dikes do not cut intervening sedimentary sequences nor the shear zone that separates the greenstone belt from the basement. Helmstaedt et al. (1986) describe a pillow lava sequence that grades down into sheeted dikes and gabbro from the Yellowknife greenstone belt, but interpreted the basal contact of the belt as an unconformity on a banded iron formation, an interpretation questioned by Kusky (1987). The dikes and pillow lavas are geochemically similar to MORB (MacLaughlin and Helmstaedt, 1995), although Isachsen et al. (1991), and Isachsen and Bowring (1997) have shown that the Yellowknife greenstone belt contains several different, and probably unrelated volcanic and sedimentary sequences, separated by as much as 50 Ma and spanning an age interval of 200 Ma. In this volume, Patricia Corcoran, Wulff Mueller, and Tim Kusky present a review of the current status of the ophiolitic interpretation of some of the greenstone belts in the Slave Province, synthesizing fifteen years of debate on the interpretation. Harper (1985) and Wilks and Harper (1997) describe rocks of the South Pass area in the Wind River Range, Wyoming, as containing a dismembered metamorphosed Archean ophiolite. This ophiolite contains all of the units of a complete ophiolite except the basal peridotite tectonite, and contacts between all units are shear zones. Cumulate textures in ultramafic rocks and gabbros are present, as are small exposures of a sheeted dike complex. Pillow lavas are associated with metapelites and banded iron formation. It has been argued (Bickle et al., 1994) that the paucity of well-developed sheeted dike complexes known from Archean greenstone belts indicates that they are not ophiolites. But sheeted dikes are not well-preserved in many Phanerozoic ophiolites, especially when they are metamorphosed and deformed to the extent that most Archean greenstone belts are. Abbott (1996) argues that sheeted dikes are not necessarily formed in every ocean floor sequence. Despite this, sheeted dike complexes have been discovered in several of the ophiolitic greenstone belts described above. Well-developed sheeted dike complexes have also been mapped in several locations in the Kalgoorlie terrane of the Yilgarn craton (Fripp and Jones, 1997). Multiple cooling units of dolerite, high-Mg mafic rocks, and ser-

10

Introduction

pentinite are truncated at an angle between 35◦ –80◦ by an unconformably overlying mafic volcanic breccia, pillow breccia, and lenses of pillow lava that strike parallel to bedding in overlying sedimentary rocks in the Kanowna Lake area. Fripp and Jones (1997) interpret this unit as a sheeted dike complex overlain by a volcanic carapace. At the Cowan Lake Six Islands locality, Fripp and Jones (1997) describe lherzolite and dunite that grade up into websterite and gabbro with pyroxenite layers. These rocks are overlain by high-Mg mafic and picritic basalts that occur in multiple tabular cooling units, interpreted as sheeted dikes that exhibit both one-way and two-way chill margins. These are overlain by chert, silicified mudstone, shale and graywacke turbidites, which locally occur as partially assimilated xenoliths (containing zircons) within the intrusive rocks. Fripp and Jones (1997) interpret the Lake Cowan greenstone locality to include the peridotitic lower plutonic sequence that marks the transition zone between mantle and crust in ophiolite suites. This transition zone sequence is overlain by a sheeted dike complex, but the extrusive magmatic carapace is omitted by faulting at this locality. Fripp and Jones (1997) note the many similarities between the Kalgoorlie ophiolites and Phanerozoic ophiolites such as the Samail, Troodos, and Bay of Islands massifs. Kimura et al. (1993) interpret parts of the Larder Lake and Beardmore-Geraldton greenstone belts in the Abitibi and Wabigoon subprovinces of the Superior Province to include ophiolitic fragments accreted in arc environments, in a manner analogous to the setting of the basalt/chert slivers of the Schreiber-Hemlo belt (Kusky and Polat, 1999). The Larder Lake belt occurs in the southern part of the Abitibi greenstone belt, and consists of pillow basalts and banded iron formation (BIF) tectonically stacked with terrigeneous turbidites. The pillow basalts and BIF are interpreted to be the upper part of an oceanic plate stratigraphy, offscraped and interdigitated with trench turbidites in an accretionary wedge setting similar to Alaska (Kusky et al., 1997a, 1997b, 2004a; Kusky, 2004) or Japan (Isozaki et al., 1990). Kimura et al. (1993) and Williams et al. (1991) also suggest a similar exotic origin for basalts and iron formation tectonically interleaved with terrigeneous turbidites in the Beardmore-Geraldton area in the southern part of the Wabigoon subprovince. In both of these examples, the accreted trench turbidites and ophiolitic slivers are intruded and overprinted by arc-related plutons and lavas, formed when the trench stepped back and intruded its own accretionary wedge. A similar accretionary wedge setting and oceanic crustal origin for slivers of basalt in greenstone belts of the Pilbara craton has been proposed by Isozaki et al. (1992). Evidence for the creation and obduction of oceanic crust in the Archean is not limited to field relationships as described above. Jacob et al. (1994) report that the geochemistry of diamondiferous eclogites from the Udachnaya Mine, Siberia (Puchtel et al., 1997), are most consistent with derivation from subducted slabs of Archean oceanic crust that were extensively hydrothermally altered prior to subduction. Similarly, many eclogite samples from South African kimberlites are also interpreted as remnants of subducted Archean oceanic crust (e.g., Jagoutz et al., 1984; MacGregor and Manton, 1986; Carlson et al., 2000). In summary, dismembered ophiolites are a widespread component of Archean greenstone belts, and many of these apparently formed as the upper parts of Archean oceanic crust. Most of these appear to have been accreted within forearc and intra-arc tectonic set-

5. Historical Recognition of Proterozoic Ophiolites

11

tings. The observation that Archean greenstone belts have such an abundance of accreted ophiolitic fragments compared to Phanerozoic orogens suggests that thick, relatively buoyant, young Archean oceanic lithosphere may have had a rheological structure favoring delamination of the uppermost parts during subduction and collisional events (Hoffman and Ranalli, 1988; Kusky and Polat, 1999). 5. HISTORICAL RECOGNITION OF PROTEROZOIC OPHIOLITES In contrast to the Archean, a number of ophiolites have been recognized and generally accepted in Proterozoic terranes for a number of years. The Late Proterozoic ArabianNubian Shield hosts a number of ophiolite-decorated sutures, and boasts one of the highest ophiolite densities known for a Proterozoic terrane on the planet. The Arabian-Nubian Shield is part of the East African Orogen that stretches from the Arabian-Nubian Shield in the north, through parts of India, east Africa and Madagascar, and has uncertain links with Neoproterozoic orogens in Antarctica and elsewhere around the globe. The East African Orogen has a complex history including a record of the break-up of Rodinia at circa 900– 800 Ma, and the evolution of numerous arc systems, oceanic plateaux, oceanic crust, and sedimentary basins. Neoproterozoic closure of the Mozambique Ocean sutured east and west Gondwana along the length of the East African Orogen (Stern, 1994; Kusky et al., 2003a, 2003b). An accretionary collage of arc, ophiolitic, and microcontinental terranes formed during closure of the Mozambique ocean is now preserved in the Arabian-Nubian shield. Some of the arc terranes appear to represent juvenile additions to the continental crust during this time period, whereas others may have been built on older continental basement on the margins of the Mozambique Ocean. Many of the ophiolites in the East African Orogen and Arabian-Nubian shield record different aspects of this complex history. Careful field mapping, geochronology, geochemistry, and structural analyses of some of these ophiolites has led to significant improvements in understanding of heat loss from the Neoproterozoic Earth, continental growth, and preservation of juvenile crustal terranes. Paleo and Meso-Proterozoic ophiolites are less abundant than the Neoproterozoic ophiolites so well preserved in the Arabian-Nubian Shield. A few Mesoproterozoic ophiolitic terranes have been described from the Karelian Shield, Cape Smith Belt, West Africa, and the SW USA (e.g., Abouchami et al., 1990; Boher et al., 1992; Dann, 2004; Scott et al., 1991, 1992; St-Onge et al., 1989). However, until recently, no Mesoproterozoic ophiolites were known. The first section of this book presents descriptions of several recently recognized Proterozoic ophiolite sequences, as well as reviews of some of the more classic Proterozoic ophiolitic terrains. 6. THE ROLE OF OCEANIC LITHOSPHERE IN CONTINENTAL GROWTH Although the rate of continental growth is a matter of geological debate (Fyfe, 1978; Dewey and Windley, 1981; Armstrong, 1991; McCulloch and Bennett, 1994; Taylor and McLennan, 1995; Rudnick, 1995; Artemieva and Mooney, 2001), most geological data

12

Introduction

indicates that the continental crust has grown by accretionary and magmatic processes taking place at convergent plate boundaries since the early Archean (Burke et al., 1976; Sleep and Windley, 1982; Taylor and McLennan, 1995; Friend et al., 1988; Card, 1990; Condie, 1994, 1997a; Sengör and Natal’in, 1996; Windley et al., 1996). Arc-like trace element characteristics of continental crust suggest that subduction zone magmatism has played an important role in the generation of the continental crust (Taylor and McLennan, 1995; Hofmann, 1988; Tarney and Jones, 1994; Rudnick, 1995); a corollary of this observation is that oceanic crust and lithosphere must have been created in order to be subducted. Convergent margin accretionary processes that contribute to the growth of the continental crust include oceanic plateau accretion, oceanic island arc accretion, normal ocean crust (mid-ocean ridge) accretion/ophiolite obduction, back-arc basin accretion, and arc-trench migration/Turkic-type orogeny accretion (Ben-Avraham et al., 1981; Kusky and Polat, 1999). These early accretionary processes are typically followed by intrusion of late stage anatectic granites, late gravitational collapse, and late strike slip faulting (Kusky, 1993; 1998). Together, these processes release volatiles from the lower crust and mantle and help to stabilize young accreted crust and form stable continents (see also Abbott, 1991; Abbott and Mooney, 1995; Abbott et al., 1997), and also alter in various ways the original chemical and physical relationships that make the primary tectonic setting of the rock suites difficult to determine. Thus, in order to understand the earliest physical and chemical conditions in potential Precambrian ophiolite suites it is typically necessary to first unravel a complex suite of later superimposed structures and mineral assemblages. Sengör et al. (1993) and Sengör and Natal’in (1996, and see also this volume) proposed a new type of orogeny, so-called “Turkic or accretionary-type orogeny”, for continental growth. These orogenic belts possess very large sutures (up to several hundred km wide) characterized by subduction-accretion complexes and arc-derived granitoid intrusions, similar to the Circum-Pacific accreted terranes (e.g., Alaska, Japan). These subductionaccretion complexes are composed of tectonically juxtaposed fragments of island arcs, back-arc basins, ocean islands/plateaux, trench turbidites, and micro-continents (Sengör, 1993; Sengör and Natal’in, 1996). Another important feature of these orogens is the common occurrence of orogen parallel strike-slip fault systems, resulting in lateral stacking and bifurcating lithological domains (Sengör and Natal’in, 1996). In these respects, the accretionary-type orogeny may be considered as unified accretionary model for the growth of the continental crust. In the concluding section of this book, Sengör and Natal’in (2004) present an analysis of the Altaids orogen and compare ophiolitic fragments therein with similar ophiolitic fragments in Precambrian terranes, especially Archean granitegreenstone terranes. Their similarity is remarkably, suggesting a common origin.

7. CONVERGENT MARGIN ANALOGS TO ARCHEAN GREENSTONE BELTS The southern Alaska convergent margin consists of belts of accreted PaleozoicCenozoic flysch, mélange, and ophiolites. Paleogene near-trench plutons intruded the margin diachronously from 61 Ma in the west to 50 Ma in the east during migration of a

7. Convergent Margin Analogs to Archean Greenstone Belts

13

trench-ridge-trench triple junction. Recognizing these plutons as a product of ridge subduction has several implications for forearc evolution and interpretation of linear belts of plutons, flysch, ophiolites, and mélange in Precambrian orogens. Forearcs are not necessarily places characterized exclusively by high-P, low-T, metamorphic series and a lack of plutonism, but may contain high-T, low-P metamorphism in association with belts of plutonic rocks if the forearc was affected by ridge subduction. Similarly, belts of magmatic rocks in Precambrian orogens (e.g., Wylie et al., 1997; Rapp, 1997) may not necessarily represent individual arc terranes, but could be a paired arc/forearc system that experienced ridge subduction. The record of ridge subduction events varies considerably depending on plate geometry and rates of triple junction and slab window migration. However, some of the hallmark signatures of ridge subduction in forearcs include along strike diachronous intrusion of TTG (tonalitic, trondhjemitic, granodiorite), to granitic plutons, high-T metamorphism, emplacement of ophiolites in the forearc, diachronous gold mineralization, and belts of anomalous complex faulting. Structural, thermal and magmatic aspects of the Chugach terrane are similar to the geology of Archean granite-greenstone terranes. In both, deformation is locally melt-dominated, and plutons follow a low-K series from diorite to trondhjemite. Metamorphism is of a high-T, low-P series. Most Archean granite-greenstone terranes acquired their first-order structural and metamorphic characteristics at convergent plate margins, where large accretionary wedges similar in aspect to the Chugach, Makran, and Altaids grew through offscraping and accretion of oceanic plateaux, oceanic crustal fragments, juvenile island arcs, rifted continental margins, and pelagic and terrigeneous sediments. Some suites of TTG in these terrains appear to have been generated during ridge subduction events, suggesting that ridge subduction is an important process in continental growth. Ridge subduction was likely an important process in the Archean, when the total number of plates was higher, and the number of ridge-trench encounters was greater. The southern Alaska margin serves as a relatively modern example of processes important in Archean forearc evolution and continental growth. 7.1. Near Trench Magmatism Associated with Ridge Subduction Near-trench magmatic rocks in southern Alaska related to subduction of the Kula-Farallon ridge include muscovite-biotite granodiorites, leucotonalites, trondhjemites, and dikes with basaltic to rhyolitic compositions. The ages of these plutons are diachronous along strike (Bradley et al., 1994, 2003; Kusky et al. 1997a, 1997b, 2003a, 2003b), and track the migration of the Kula-Farallon-North America triple junction as it migrated along the Cordilleran margin. The plutons can be used as time-markers to help construct a model for the structural evolution of the wedge during ridge subduction (Kusky et al. 1997a, 1997b, 2003a, 2003b). 7.2. Generation and Emplacement of Ophiolites at the Triple Junction A geochemical study and petrologic model for the generation of the Resurrection and Knight Island ophiolites (Lytwyn et al., 1997; Kusky et al., 2004b, 2004c, 2004d), formed

14

Introduction

at the triple junction and soon after incorporated into the accretionary wedge concluded that the T-MORB’s show geochemical variations related to derivation from and mixing of multiple parental magmas, derived through near-fractional melting of variably depleted mantle sources in a subaxial melting column. Geochemical evidence for contamination of the magmas by sedimentary material similar in composition to nearby flysch, supports the interpretation that these ophiolites formed in a near-trench triple junction environment. Kusky and Young (1999) relate emplacement of the ophiolite to ridge subduction, and show that the ophiolite was formed at a triple junction, incorporated into the accretionary prism, and intruded by near-trench magmas, all within 3.6 Ma. Magmas of the Aialik pluton intruded along a shear zone formed during incorporation of the ophiolite into the accretionary prism. The sedimentary geochemistry of the shale/turbidite sequence overlying the ophiolite shows a progressively increasing terrigenous source up-section, as the future Resurrection ophiolite approached the Chugach accretionary prism. 7.3. Comparison to Crustal Growth Mechanisms in the Precambrian Comparison of crustal growth processes in accretionary prisms affected by ridge subduction with crustal growth as seen in Archean shield areas offers insight about processes in Archean subduction zones. From this, it should be possible to eventually estimate if ridge subduction was more common in Precambrian terranes, when plate boundary lengths were greater (and total number of plates greater), and subducting lithosphere was on average younger. Ridge subduction related processes should have been more important in the older geologic record, as the number of triple junctions increases nearly twice as fast as the number of plates. With the probable larger number of plates accommodating greater heat production in the Archean, the number of triple junctions would increase, and the role of ridge subduction is thus expected to have been more important in the growth of the continental crust in Precambrian times.

8. SYNOPSIS OF THIS VOLUME The first section of this book contains several chapters that present descriptions and interpretations of Proterozoic ophiolites, including reviews of previously described ophiolites, and some descriptions of sequences interpreted as ophiolites for the first time. 8.1. Proterozoic Ophiolites and Related Rocks Until the recent discovery of the 2.5 Ga Dongwanzi ophiolite, many students of Precambrian geology regarded the circa 1.96–2.0 Ga Jourma ophiolite of Finland and the 1.998 Ga Purtuniq ophiolite from the Cape Smith Belt as the world’s oldest, nearly complete well-preserved ophiolites. The Jourma ophiolite is a remarkable polyphase ophiolite that preserves evidence for formation at an ocean-continent transition. In this volume, Petri Peltonen and A. Kontinen present a detailed description and review of recent work on the

8. Synopsis of This Volume

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Jourma ophiolite by the Finnish Geological Survey and other teams. They conclude that the Jormua Ophiolite is an allochtonous mafic-ultramafic rock complex, thrusted onto the Karelian Craton margin, that formed within a passive margin environment ∼ 100 km southwest from its present position. This complex consists of two distinct units: (1) fragments of ancient subcontinental lithospheric mantle that became exposed beneath the Archean craton by detachment faulting following the final break-up of the craton, and (2) alkaline and tholeiitic igneous suites that were emplaced within and through the lithospheric mantle at ∼ 2080 Ma and 1950 Ma, respectively. The mantle peridotites had yielded melt already before they were intruded by the oldest suite of dikes at > 2800 Ma. These old dikes are “dry” clinopyroxene cumulates being products of an Archean magmatic episode. Later, during the initial stages of continental break-up at ∼ 2080 Ma, this same piece of mantle became extensively intruded by hydrous alkaline magmas that resulted in formation of high-pressure hornblendite-garnetite cumulates deep in the ophiolite stratigraphy and fine grained OIB-type dikes at more shallow levels. Simultaneously, the residual peridotites became metasomatized due to porous flow of the melt in the peridotite matrix. Alkaline magmatism was soon followed by lithospheric detachment faulting that exposed the subcrustal peridotites at the seafloor, where they became covered by tholeiitic (EMORB) pillow and massive lavas and intruded by coeval dikes and gabbros. Since transitional contacts between all main ophiolite units can be demonstrated, the Jormua Ophiolite Complex is interpreted to represent a practically unbroken sample of seafloor from an ancient oceancontinent transition (OCT) zone, strikingly similar to that reported from the Cretaceous West Iberia non-volcanic continental margin. From Payson Arizona, Jesse Dann (2004) presents maps, photographs, and data on the 1.73 Ga Payson Ophiolite. This well-exposed but unusual ophiolite is disposed as a shallow-dipping, layered sequence of coeval gabbro, sheeted dikes, and submarine volcanic rocks, partly disjointed by later intrusion and deformation. A sheeted dike complex is spectacularly exposed on cliffs and in stream sections in shallow canyons. The dike complex is rooted in underlying gabbro, as shown by gabbro-dike mingling and mutual intrusion. An unusual continuous zone of intense alteration marks the transition from sheeted dikes to submarine volcanics. The Payson ophiolite has many arc-like characteristics. A tonalite/dacite magmatic suite occurs as rare lavas and as dikes and hypabyssal plutons mutually intrusive with the basaltic sheeted dikes and gabbro. An older basement complex occurs as roof pendants in gabbro and screens in the sheeted dike complex. Dann suggests a model for the Payson ophiolite in which an intra-arc basin formed by seafloor spreading along an arc-parallel strike-slip fault system, and relates its emplacement within the arc and accretion to North America to events associated with the ca. 1.70 Ga Yavapai Orogeny. Proterozoic ophiolites were first widely recognized from the Arabian-Nubian Shield. In this volume, Robert Stern, Peter Johnson, Alfred Kröner, and Bisrat Yibas present a review of the Neoproterozoic ophiolites of the Arabian-Nubian Shield, whereas Peter Johnson, Fayak Kattan, and Ahmed Al-Saleh describe the field characteristics of some of the better-exposed ophiolites in more detail. Ophiolites of the Arabian-Nubian Shield range in age from 690 to 890 Ma and in the northern part of the shield, occur as nappe complexes marking suture zones between terranes. Although dismembered and altered, all of the diag-

16

Introduction

nostic components of ophiolites can be found, including harzburgite, cumulate ultramafics, layered and higher level gabbro and plagiogranite, sheeted dikes, and pillow basalt lavas. Allochthonous mafic-ultramafic complexes in the southern part of the shield in Ethiopia and Eritrea are interpreted as ophiolites, but are more deformed and metamorphosed than those in the north. Reconstructed ophiolitic successions have crustal thicknesses of 2.5 to 5 km. The Arabian-Nubian shield ophiolitic mantle was mostly harzburgitic, and exhibits chemical compositions comparable to modern forearcs and distinctly different from midocean ridges and backarc basin peridotites. Arabian-Nubian shield ophiolites are often associated with a thick (1–3 km) sequence of cumulate ultramafic rocks, which define a transition zone between the seismic and petrologic Mohos. These cumulates are dominated by dunite, with subordinate pyroxene-rich lithologies. Cumulate ultramafics grade upwards into layered gabbro. Both tholeiitic and calc-alkaline affinities are present, and a significant, although subordinate, amount of boninites have been identified. The ArabianNubian shield ophiolitic lavas include both LREE-depleted and LREE-enriched varieties, but as a group are slightly LREE-enriched. On a variety of discrimination diagrams, the lavas plot in fields for MORB, BABB, arc tholeiite, and boninite. Nd-isotopic compositions indicate derivation from a long-depleted mantle source. Mineral and lava compositions are consistent with the hypothesis that most Arabian-Nubian shield ophiolites formed in ‘suprasubduction zone’ (SSZ) settings, and the high Cr# of Arabian-Nubian shield ophiolitic harzburgites suggests a forearc environment. Stern et al. (2004) conclude that studies of deep water sediments deposited on Arabian-Nubian shield ophiolites are needed to better characterize and understand the Neoproterozoic ocean where the ophiolites formed. In a chapter that is complimentary to the overview by Stern and others, Peter Johnson, Fayek Kattan, and Ahmed Al-Saleh describe field relationships from a number of the Arabian Shield ophiolites. Where most complete, they consist of serpentinized peridotite, gabbro, dike complex, basalt, and pelagic rocks. However, because of folding and shearing, the majority of the ophiolites lack one or more of these diagnostic lithologies. Nonetheless, the incomplete assemblages are identified as ophiolites because they minimally include peridotite and gabbro, in many cases are associated with basalt, and in all cases show evidence of emplacement by thrusting and shearing rather than intrusion. The ophiolites range in age from ∼ 870 Ma to ∼ 695 Ma, documenting a 200-million year period of oceanic magmatism in the Arabian shield, and are caught up in ∼ 780 Ma to ∼ 680 Ma suture zones that reflect a 100-million year period of terrane convergence. All the ophiolites are strongly deformed, metamorphosed, and altered by silicification and carbonatization. Low-grade greenschist facies metamorphism predominates, but in places the rocks reach amphibolite grade. Alteration resulted in the development of listwaenite, particularly in shear zones, and locally the only evidence that mafic-ultramafic rocks underlie a given area is the presence of upstanding ridges of listwaenite that are resistant to erosion. Kinematic indicators in shear zones indicate that the ophiolites were affected by both strike-slip and vertical displacements. Variations in senses of shear observed along and across strike demonstrate considerable strain partitioning during deformation. However, prevailing senses of shear can be discerned for several of the ophiolites that, in conjunction with other structural

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observations, indicate the main shear trajectories of the shear zones containing the ophiolites. Jabal Ess, Jabal Tharwah, and Bi’r Umq ophiolites were emplaced during periods of dextral transpression on the Yanbu and Bi’r Umq sutures, respectively. The Bi’r Tuluhah ophiolite was emplaced during sinistral transpression of the Hulayfah-Ad Dafinah-Ruwah suture, and the Halaban ophiolite was emplaced during west-directed convergence on the Halaban suture. In a final paper on the Arabian-Nubian Shield, Ibrahim Hussein, Alfred Kröner, and Thomas Reischmann describe aspects of the 808 ± 14 Ma Wadi Onib mafic-ultramafic complex, located within the Onib-Sol Hamed suture in the northern Red Sea Hills of the Sudan, and relate these features to suprasubduction zone processes. The Wadi Onib ophiolite consists, from bottom to top, of a basal peridotite unit, an exceptionally thick (2–3 km) transitional zone of interlayered cumulates, isotropic gabbroic with plagiogranite bodies, a sheeted basic dike complex, and pillowed basaltic lavas containing fragmentary lenses of ribbon chert and/or graphitic to shaly carbonates. Whereas the basal unit is strongly serpentinized and/or carbonatized, the transitional zone comprises abundant and well preserved pyroxenites. The transition zone also shows a polycyclic cumulate arrangement that possibly originated from multiple magma pulses rather than from tectonic interslicing. Moreover, mineral grading, gravity stratification and a spectrum of folds with varying geometrical dispositions and amplitudes within discrete layers as well as a vertical metamorphic zonation (suggesting seafloor hydrothermal processes) are evident within the Onib ophiolitic sequence. In particular, the volcanic component is Ti-rich, has a transitional IAT/MORB character and is indistinguishable from anomalous MORB and/or marginal basin basalts. Thus, the Onib is envisaged to be of arc/back-arc (marginal) basin affiliation, and it classifies as a supra-subduction zone (SSZ) rather than normal MORB-type ophiolite. The ophiolitic sequence probably resulted from parental magma(s) generated through multi-stage partial fusion of mantle peridotite. In a paper summarizing the youngest ophiolite described in this volume, Joerg Pfänder and Alfred Kröner (2004) present field, geochronologic, and geochemical data on the tectono-magmatic evolution, age and emplacement of the 570 million year old Agardagh Tes-Chem ophiolite in Tuva, Central Asia. This ophiolite is located in the Palaeozoic Central Asian Mobile Belt which formed during subduction-accretion processes lasting from the early Neoproterozoic to the late Palaeozoic. The ophiolite was obducted onto the TuvaMongolian Massif (microcontinent?) in the early Palaeozoic towards the SE along N- and NW-dipping faults and is embedded within a tectonic mélange and thus is part of an accretionary wedge. Dating of three small zircon fractions from a plagiogranite by the evaporation technique yielded a 207 Pb/206Pb age of 569.6 ± 1.7 Ma, which reflects the crystallization age of the plutonic section of the ophiolite. Geochemical data reveal an island arcrelated origin for the ophiolite, where typical island arc volcanic rocks predominate over MORB-like pillow lavas. In contrast to the highly incompatible element-enriched volcanic rocks, all plutonic rocks of the ophiolite are depleted, and mineral compositions of ultramafic cumulates indicate the presence of boninitic parental melts. The ophiolite therefore consists of an association of island arc and back-arc related sequences that have been amalgamated during subduction-accretion and collisional obduction. Isotopic and trace element

18

Introduction

data reveal the existence of a depleted and refractory mantle source beneath Central Asia, from which the volcanic and plutonic rocks of the ophiolite were formed. However, source contamination took place by sediment subduction, before the parental melts of the island arc volcanic rocks were formed. 8.2. Archean Ophiolites and Related Rocks The second main section of the book presents descriptions of a number of well-exposed Archean mafic-ultramafic sequences that have been suggested to be possible ophiolites. Recognizing ophiolites in the Archean record has been more controversial than calling similar sequences in the Proterozoic record ophiolites, and many of the papers in this section discuss this bias of some workers against calling Archean ophiolite-like sequences ophiolites. In a series of four chapters on the Dongwanzi ophiolite, Timothy Kusky, Jianghai Li, and their co-workers describe various aspects of a remarkable complete but dismembered ophiolite sequence discovered in the North China craton in 2001. In the first overview paper, Timothy Kusky, Jianghai Li, Adam Glass, and Xiongnan Huang describe the general field characteristics of the 2.5 Ga Dongwanzi ophiolite, and discuss its regional tectonic setting. Banded iron formation structurally overlies several tens of meters of variably deformed pillow lavas and mafic flows. These are in structural contact with a 2 km thick mixed gabbro and sheeted dike complex with gabbro screens, exposed discontinuously along strike for more than 20 km. The dikes consist of metamorphosed diabase, basalt, hb-cpx-gabbro, and pyroxenite. Many have chilled margins developed on their NE sides, indicating one-way chilling. The dike/gabbro complex is underlain by several kilometers of mixed isotropic and foliated gabbro, which develop compositional layering approximately two kilometers below the sheeted dikes, and then over several hundred meters merge into strongly compositionally layered gabbro and olivine-gabbro. The layered gabbro becomes mixed with layered pyroxenite/gabbro marking a transition zone into cumulate ultramafic rocks including serpentinized dunite, pyroxenite and wehrlite, and finally into strongly deformed and serpentinized olivine and orthopyroxene-bearing ultramafic rocks interpreted as depleted mantle harzburgite tectonites. A U/Pb zircon age of 2.505 Ga from gabbro of the Dongwanzi ophiolite makes it the world’s oldest recognized, laterally-extensive complete ophiolite sequence. Characteristics of this remarkable ophiolite may provide the best constraints yet on the nature of the Archean oceanic crust and mantle, and offer insights to the style of Archean plate tectonics and global heat loss mechanisms. In a companion paper, Rachael Huson, Timothy Kusky, and Jianghai Li (2004) compare major and trace element concentrations of rocks from the Dongwanzi ophiolite with known concentrations from well-studied ophiolites and rocks from tectonic settings to determine tectonic environment of formation. Major element analysis shows samples are subalkalic (in particular, calc-alkaline) to alkalic. Trace element analysis shows enrichment of large ion lithophile elements as well as depletion of high field strength elements relative to mid-ocean ridge basalts. Calc-alkaline geochemical characteristics of oceanic rocks have predominantly been identified in suprasubduction zone settings and their occurrence in the

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Donwangzi ophiolite suggests a similar tectonic setting. Trace element signatures are also similar to suprasubduction zone ophiolites indicating formation above a subduction zone. The Dongwanzi ophiolite is but one of the largest well-preserved greenstone belts in the Central Orogenic belt that divides the North China craton into eastern and western blocks. More than 1,000 other fragments of gabbro, pillow lava, sheeted dikes, harzburgite, and podiform-chromite bearing dunite occur as tectonic blocks (tens to hundreds of meters long) in a biotite-gneiss and BIF matrix, intruded by tonalite and granodiorite, in the Zunhua structural belt. Blocks in this metamorphosed Archean ophiolitic mélange preserve deeper levels of oceanic mantle than the Dongwanzi ophiolite. The ophioliterelated mélange marks a suture zone across the North China Craton, traced for more than 1,600 km along the Central orogenic belt. Many of the chromitite bodies are localized in dunite envelopes within harzburgite tectonite, and have characteristic nodular and orbicular chromite textures, known elsewhere only from ophiolites. The chromites have variable but high chrome numbers (Cr/Cr + Al = 0.74–0.93) and elevated P, also characteristic of suprasubduction zone ophiolites. The high chrome numbers, coupled with TiO2 wt% < 0.2 and V2 O5 wt% < 0.1 indicate high degrees of partial melting from a very depleted mantle source and primitive melt for the chromite. As reported in the chapter by Kusky, Li, Raharimahefa, and Carlson (2004d), a Re-Os model age from the chromites indicates an age of 2547 ± 10 Ma, showing that they are the same age as the Dongwanzi ophiolite. The range in initial Os isotopic compositions in the chromites in these ophiolitic blocks is small and well within the range seen in modern ophiolites. The chondritic to sub-chondritic initial ratio also is interesting in that it shows more similarity to the values found for abyssal peridotites than OIB’s, pointing to an ocean-ridge rather than plume setting for the initial formation of these peridotites. The ultramafic and ophiolitic blocks in the Zunhua mélange are therefore interpreted as dismembered and strongly deformed parts of the Dongwanzi ophiolite. Xiongnan Huang, Jianghai Li, Timothy Kusky, and Zheng Chen describe a remarkably well-preserved suite of microstructures from the Zunhua podiform chromite, and discuss implications for the deformation and rheology of the Archean oceanic lithospheric mantle. The Zunhua podiform chromite preserves typical magmatic fabrics including nodular and orbicular textures, and magmatic flow structures. The magmatic textures indicate that the Zunhua podiform chromite was formed through five-stages of evolution, with the following time sequence: disseminated chromite, net-like veins, antinodular, orbicular and nodular textures. The evolution of the texture series can be interpreted to result from fast flowing magmatic flowing systems. They result from the vertical accretion of the oceanic mantle. The podiform chromite ores show strong deformation with development of pull-apart structures, banding, folds, and mylonitic foliation. These structures were formed at high temperature in the oceanic mantle during the oceanic ridge spreading as the ores were caught up by plastic flow and sheared transversely. The Zunhua podiform chromite bodies result from active magmatic accretion and strong high-temperature plastic flow, therefore a fast spreading oceanic ridge is suggested for its formation. Silicate mineral inclusions within the chromium spinel and geochemical characteristics of the Zunhua ophiolite support a geological setting in a suprasubduction belt.

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Introduction

Several hundred kilometers to the southwest of the Dongwanzi ophiolite, Jianghai Li, Timothy Kusky, Niu Xianglong, and Feng Jun describe the textures and mineralogy of a Neoarchean massive sulfide deposit in the Wutai Mountains, recognizing a black smoker chimney and mound complex within 2.50 Ga old oceanic crust. The Wutai VMS is one of largest sediment-hosted sulfide deposits in China. It forms small lenses, thin sheets, and tabular bodies of massive to layered sulfide, disseminated through a forearc mélange belt. Although they are reworked by late deformation, sulfide deposits formed at different crustal levels still can be identified, including relicts of chimneys, pyritic siliceous exhalite, massive crystallized sulfides, talus of massive sulfides and stockwork zones. The country rock of the Wutai VMS ores show intense silicification and chloritization. Epidosites have been identified within mafic rocks. Under microscopic observations, porous sulfides show a mineralogical zonation around micro-conduits. The colloform textures developed delicate banding and concentric textures. The vuggy cavities are commonly lined by concentric layers consisting of idiomorphic pyrite and silica. The Wutai VMS are spatially associated with convergent plate boundaries, formed in the upper sequence of a former Neoarchean oceanic basin. They have been overthrust by foreland-thrust belts following closure of an oceanic basin. The preliminary studies reveal the presence of black smoker chimneys preserved in the Wutai Mountains, which suggest that seafloor black smoker activity at about 2.50 Ga plays an important role for generation and accumulation of Wutai VMS. In addition, the Wutai VMS is quite similar to Besshi-type deposits, it is inferred to be generated in the setting of forearc, later tectonically transported in mélange belts during continental collision. The possibility that some mafic greenstone belts in the Slave Craton of northern Canada and Nunavut may be ophiolites has been a contentious issue with debate spanning much of the late 1980’s, 1990’s, and early 2000’s. In the sixth chapter in this section, three authors (Patricia Corcoran, Wulff Mueller, and Timothy Kusky) with different views on this subject have co-written a paper that attempts to lay-out the most important observations, and reconcile possible interpretations with these observations. This chapter reviews three distinct areas in the Slave craton and assesses their potential of containing ophiolite sequences. These include the (1) Yellowknife, (2) Point Lake, and (3) Beaulieu and Cameron River volcanic belts. Since in all these case, the mafic-ultramafic sequences that have some ophiolitic characteristics rest structurally over continental crust, the best modern analogy for Slave Province ophiolites may be Tethyan-type ophiolites. The 6 km thick Chan Formation of the Yellowknife volcanic belt resembles modern ophiolites with tholeiitic massive to pillowed flows, abundant gabbro dikes and sills, interflow sedimentary rocks, and a mafic sheeted dike swarm. The base of this crustal-floored sequence is sheared and locally stitched by late-tectonic plutons and the dunite-peridotite-gabbro segment is lacking, so if it is an ophiolite, it only contains the upper parts of the sequence. The inferred base of the Point Lake volcanic belt is composed of mafic mylonite with low-strain domains of gabbro, pyroxenite, dunite, and peridotite. The mafic mylonite is overlain by gabbro, layered gabbro, minor mafic dikes, pillowed flows, massive flows, hyaloclastite, and local chert. A well-defined sheeted dike swarm is absent although mafic dikes are locally preserved in the crustal sequence, and other dikes cut underlying granite. The Beaulieu and

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Cameron River volcanic belts are spatially associated with mafic dike swarms that intrude the Sleepy Dragon basement complex. However, the currently juxtaposed dike swarm and mafic volcanic belt are not necessarily directly related, since they are everywhere separated by a major shear zone with significant displacements. Mafic massive and pillowed flows and sub-volcanic sills are predominant above the sheared basement contact. In the strict sense, these belts or belt segments do not fit the definition of a complete ophiolite, but do meet the general requirements in that they are allochthonous mafic sequences consisting of submarine volcanics and intrusives. Ophiolites form in numerous tectonic settings and complete preservation from tectonized mantle to surficial ocean floor products is highly unlikely, especially for Archean rocks. Therefore, the nature of the basement contacts is particularly significant. If the contacts are tectonic, then parts of the ophiolitic sequences may have been sheared off, which is commonly the case for the mafic-ultramafic intrusive component. Recent models have compared parts of certain Slave Province greenstone belts with supra-subduction zone settings including arcs and back-arcs, extensional settings such as mid-ocean ridges. Some of the ophiolitic Slave Province greenstone belts have characteristics that suggest they formed along ocean-continent transition zones, similar to the Jourma ophiolite of Finland and the Cretaceous west Iberian continental margin. One of the important points that Corcoran, Mueller, and Kusky develop is that identifying ophiolite sequences based solely on geochemistry is overly simplistic, and regional geological context, structure, and stratigraphy is required (e.g., Pearce, 1987; Wood et al., 1979). Ophiolites are generally considered a distinct suite of obducted ocean floor rocks with a highly varied geochemical affinity depending on tectonic setting (Sylvester et al., 1997; Dilek et al., 2000). As pointed out by Eldridge Moores (2002) the “ophiolite conundrum” marks a discrepancy between structural and stratigraphic setting, and geochemical characteristics. Moores argues that the mantle is heterogeneous at all scales and geodynamic settings, and that a distinct geochemical signature for ophiolites is lacking. This has ramifications especially for the Archean, in which volcano-sedimentary sequences are generally incomplete, structures are complex, Fe-tholeiites and komatiites are abundant (rare to non-existent in modern ophiolites), and mantle compositions and temperatures were possibly different. Identifying the tectonic setting of a dismembered mafic (-ultramafic) volcanic sequence thus becomes enigmatic. Must there be a complete stratigraphic sequence (i.e., dunite-peridotite, tectonite, gabbro, sheeted dykes, pillows, pelagic sedimentary rocks) to qualify a specific sequence as an ophiolite? What portion of the succession is necessary in order to be called an ophiolite? Is there a distinction between Phanerozoic and Archean ophiolites? Further north in the Aldan Shield in eastern Siberia, Igor Puchtel (2004) describes a 3.0 billion year old partial ophiolitic sequence from the Olondo greenstone belt. The Olondo greenstone belt is distinguished from the other greenstone belts in the Aldan Shield by an abundance and a great facies diversity of mafic-ultramafic rocks. The rocks are relatively well preserved both geologically and geochemically compared to other Archean ophiolitelike sequences worldwide, and thus can be regarded as valuable witnesses of the early history of the Earth. The Olondo greenstone belt contains one of the oldest ophiolite-like sequences on the planet. The age of the Olondo greenstone belt at 3.0 Ga is intermediate

22

Introduction

between the two most commonly cited periods of global crust-forming activity, namely, 2.7 and 3.4 Ga (Condie 1995, 1998). Thus, the study of this belt can help fill the gap in our understanding the significance of the tectonothermal and chemical evolution of the Earth during the time period between the early and late Archean. Andrey Shchipansky and others (2004) describe Neoarchean subduction-related assemblages of the North Karelian greenstone belt, in the northeast part of the Baltic Shield, Russia. This belt contains some of the world’s oldest known boninite series rocks, occurring in at least in two areas of the belt. The first area, referred to here as the Khizovaara structure, shows evidence of a late Archean ocean-island volcanic arc collage formed during two tectonic episodes nearly 2.8 Ga ago. The second area, named the Iringora structure, preserves distinctive features of an ophiolite pseudostratigraphy, including not only gabbro and lava units, but also remnants of a sheeted dike complex. The major and trace element chemistry of the Iringora ophiolitic gabbro, dike and lava units suggests a comagmatic series with a continuous compositional variation from more primitive mafic to strictly boninitic melts. In terms of major and trace element abundance, the boninite series of the North Karelian greenstone belt is practically indistinguishable from the Group I and II of the Troodos upper pillow lavas. These occurrences strongly suggest that Neoarchean subduction-related processes including boninite-hosting supra-subduction zone ophiolites have not changed substantially over the past 2.8 Ga. The Belingwe belt in Zimbabwe is probably the best-known well-preserved late Archaean greenstone belt in the world. Despite the presence of well exposed rocks of very low metamorphic grade and low strain, the tectonic evolution of the Belingwe belt has been a matter of much controversy, with debates focusing on whether the parts of the greenstone succession are oceanic in nature, or whether they were erupted through underlying continental crust. Axel Hoffman and Tim Kusky (2004) present a synthesis of the geology of the Belingwe belt, and assess various tectonic models for the belt’s origin. The Belingwe greenstone belt comprises two distinct greenstone successions. The lower, 2.9–2.8 Ga old Mtshingwe Group consists of four stratigraphic units, an intermediate to felsic volcanic and volcanoclastic unit, an ultramafic to mafic lava plain sequence, a conglomerate-shale sedimentary sequence, and a unit of tectonically imbricated sedimentary and volcanic rocks. Although geochronological, geochemical and lithological characteristics are broadly known, the tectonic evolution of the Mtshingwe Group remains a matter of speculation. Controversy surrounds the intensely studied, 2.7 Ga old Ngezi Group, which consists of a thin basal sedimentary sequence, a thick ultramafic to mafic volcanic sequence, and an upper sedimentary succession. The basal unit rests unconformably on up to 3.5 Ga old granitoid gneisses and Mtshingwe Group rocks, consists of fluvial to shallow-marine sedimentary rocks, and is similar to cratonic cover successions. The structurally overlying volcanic unit is a submarine lava plain sequence of massive and pillow basalts with komatiites near the base and andesites near the top. The upper sedimentary unit represents a foreland basin sequence and consists of karstified carbonate ramp limestones overlain by deeper-water turbidite deposits. Autochthonous versus allochthonous models have been proposed for the tectonic evolution of the Ngezi Group. Proponents of the autochthonous model regard the Ngezi Group as a conformable sequence that formed in an ensialic rift

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setting above a mantle plume. Other workers regard the volcanic sequence as an allochthonous unit of oceanic crust that was obducted onto continental basement. A great number of arguments have been proposed in recent years from structural, sedimentological, and geochemical studies for and against the different models. A critical reappraisal of the various arguments indicate the lack of convincing evidence for an ensialic and autochthonous origin. Arguments for an allochthonous origin are strong, whereas an oceanic origin can only be inferred by assuming that modernistic plate tectonic processes were operating in the late Archean. 8.3. Models for the Evolution of Oceanic Crust with Time Chapters in the third section of the book focus on melting and petrological processes in the Archean mantle, sub-ridge, and sub-arc environments, and on models for the evolution of oceanic crust with time. Authors of papers in this section have synthesized data from several different belts, and place important constraints on the nature of the Archean mantle. Archean greenstone belts are known for their hallmark deposits of komatiites, magnesium rich lavas that many petrologists have suggested indicate significantly higher temperatures for the Archean mantle. These estimated temperatures were in turn used by many workers to derive unusual non-uniformitarian and non-actualistic models for tectonics on the early Earth. Steve Parman and Tim Grove show field and petrologic data from the Barberton Greenstone belt that suggests an alternative interpretation, that the Archean mantle may not have been so different from that of today. The paper by Parman and Grove is therefore very significant in that it removes any reason for assuming that plate tectonics should have been drastically different from today. The Barberton Greenstone Belt is one of several mid- to late-Archean greenstone belts that lie along the eastern margin of the Kaapvaal craton (Brandl and de Wit, 1997). With an age of 3.49–3.46 Ga (Lopezmartinez et al., 1992), the BGB is among the oldest of the Kaapvaal Craton’s greenstone belts and is part of the nucleus around which the Late Archean greenstone belts to the north (e.g., Murchison and Giyani) and to the south (e.g., Nondweni and Commondale) were attached. Parman and Groves focus their discussion on the komatiites and related basaltic komatiites from the Komati and Hoogenoeg formations. These two formations form a continuous stratigraphic section and have been the main focus of their research, though reference is also made to komatiites in the Barberton Greenstone Belt’s smaller and less well preserved komatiite-bearing Sandspruit, Theespruit, Mendon and Weltevreden sequences. In the end Parman and Groves put the Barberton data in the context of the global komatiite data set, showing that komatiites do not require exceptionally high mantle temperatures to form. Ali Polat and Robert Kerrich (2004) synthesize data on known occurrences of boninites, adakites, magnesian andesites, and Nb-enriched basalts, and related these to Precambrian arc associations. Boninitic lavas have recently been reported from several Precambrian terranes, including the ∼ 3.8 Ga Isua terrane of West Greenland; 2.8 Ga Opatica and the 2.7 Ga Abitibi terranes of the Superior Province; the 2.8 Ga North Karelian terrane of the Baltic Shield; and the 1.9 Ga Flin Flon terrane in the Trans-Hudson orogen. In the Isua belt, boninitic flows coexist with pillow basalts and picrites. Boninitic lavas, and low-Ti

24

Introduction

tholeiitic basalts, outcrop over a 300 km corridor in the Abitibi volcanic-plutonic subprovince. They are intercalated with a stratigraphically lower ocean plateau association of komatiites and basalts, and an upper volcanic arc association of tholeiitic to calc-alkaline arc basalts; accordingly there was contemporaneous eruption of neighboring plume and arc magmas. The 2.8 Ga Opatica boninitic lavas are spatially and temporally associated with arc-type volcanic rocks. The 2.8 Ga Baltic Shield boninitic rocks are related to a supra-subduction ophiolite complex. All of these Precambrian boninitic lavas share the low-TiO2, high Al2 O3 /TiO2 ratios, U-shaped REE patterns, and negative Nb but positive Zr anomalies of Phanerozoic counterparts; however, SiO2 contents are variable. Boninites of Phanerozoic age occur in ophiolites or intra-oceanic island arcs, such as the Izu-Bonin-Mariana arc system. These primary liquids are interpreted as second-stage high-temperature, low-pressure melting of a depleted refractory mantle wedge fertilized by fluids and/or melts, above a subduction zone. Precambrian boninitic lavas are likely products of the same conjunction of processes. Low-Ti tholeiites lack the LREE enrichment coupled with negative Nb anomalies of the boninites. They had a similar depleted wedge source, but without a subduction zone component. An association of adakites, magnesian andesites (MA), and Nb-enriched basalts (NEB) with “normal” tholeiitic to calc-alkaline basalts and andesites has recently been described from the 2.7 Ga Wawa and Confederation volcanic-plutonic terranes of the Superior Province. Cenozoic adakites are considered to form by slab melting; MA the product of hybridization of adakite liquids with the peridotitic mantle wedge; and NEB melting of the residue of the MA wedge source. This volcanic association is found in Cenozoic arcs characterized by shallow subduction of young, hot oceanic lithosphere. Archean equivalents likely formed under comparable tectonic settings. U-shaped REE patterns in conjunction with positive Zr anomalies of Archean and Phanerozoic boninites can be modeled by a depleted peridotitic wedge fertilized by adakite liquids and/or hydrous fluids in a convergent margin. Consequently, Phanerozoic type arcs were operating in Archean convergent margins. Imbrication of komatiite-basalt ocean plateau volcanic sequences with arcs solves the apparent Mg#, Ni deficit of some models for Archean upper continental crust. Higher geothermal gradients in Archean subduction zones may have played an important role for the growth of continental crust. In a final, concluding chapter for the this section, Maarten de Wit synthesizes data from Archean greenstone belts around the world and concludes that “Archean greenstone belts do contain fragments of ophiolites” (a direct pun on an earlier, and flawed paper by Bickle et al., 1994). Maarten de Wit notes that most Archean greenstone belts are so severely tectonized so that reconstruction of their rock assemblages revealing original autochthonous relationships is a daunting task (de Wit and Ashwal, 1997; Kusky and Vearncombe, 1997). There are about 260 individual Archean greenstone belts worldwide. Few of these have been studied in sufficient detail to provide relatively reliable information about pre-2.5 Ga geological processes (de Wit and Ashwal, 1997). Greenstone belts represent some of the earliest records of Earth history, but they are not restricted to the Archean. For example, the large Neoproterozoic Arabian-Nubian shield has an Archean-like cratonic crust with at least 7 major greenstone belts, most of which

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comprise island arc-like successions and associated (but often dismembered) ophioliteassemblages (Berhe, 1997; see Stern et al., 2004; Johnson et al., 2004; and Hussein et al., this volume). Similarly, the Baltic shield contains greenstone belt sequences ranging in age from > 3.1 Ga (Mesoarchean) to 1.9 Ga (Mesoproterozoic). Some of the Mesoproterozoic greenstone belts share characteristics of many Archean greenstone belts (e.g., abundant komatiites), whilst others share characteristics of Phanerozoic ophiolites (Sorjonen-Ward et al., 1997, and this volume). A wide spectrum of tectonic environments is preserved within Archean greenstone belts, and many individual belts are mixtures of components from different tectonic environments and in particular from island arc terrains (de Wit and Ashwal, 1997; Kusky and Vearncombe, 1997). It is claimed nevertheless by some that oceanic crustforming environments are not preserved amongst this mixture of tectonic regimes because in their views no rocks assemblages in Archean greenstone belt sequences exhibit sufficient features to warrant definitive classification as an ophiolite (Bickle et al., 1994; Hamilton, 1998, 2003). The difficulty in recognizing and even defining ophiolites has been acknowledged widely and is not addressed here (Anonymous, 1972; and this volume). In his short contribution de Wit outlines some probable and some possible ophiolite sequences that have been reported from a number of Archean greenstone belts around the world. He also comments on the likely tectonic implications of these examples to better resolve Archean processes. 8.4. Analogs to Precambrian Ophiolites The final section of the book encompasses three important Phanerozoic analogs to processes that are currently producing rocks assemblages and structures that resemble Archean ophiolitic greenstone belts. John Encarnación (2004) describes the northern Philippines as a possible modern analogue for some Precambrian greenstone belts. It has a ∼ 150 Myr history of multiple and overlapping periods of oceanic crust generation, arc volcanism, sedimentation, and deformation dominated by wrench tectonics. At least five ophiolite complexes of distinct age make up most of the basement—all having a distinct suprasubduction zone signature, relationships reminiscent of the Yellowknife Belt in the Slave Province. Arc plutons are predominantly of the diorite-tonalite series with minor alkali-feldspar bearing rocks. Sedimentary basins probably floored by oceanic crust are dominated by immature sediments and volcaniclastics and are locally up to ∼ 10 km thick. The whole arc and ophiolitic complex is in the process of being accreted to Eurasia, where it may be preserved in a broad “suture zone” between Eurasia and Australia and/or the Americas. Southern Alaska’s Mesozoic-Cenozoic Chugach-Prince William terrane is an unusual forearc in that it contains belts of graywacke-dominated flysch, mélange, and ophiolitic fragments all intruded by a suite of tonalite-trondhjemite-granodiorite plutons, and large parts of the accretionary prism are metamorphosed to the greenschist, amphibolite, or granulite facies. In the second chapter of this section, Tim Kusky, Rose Ganley, Jennifer Lytwyn, and Ali Polat (2004c) describe the overall structural geometry, abundance and

26

Introduction

types of rocks and rock suites present, the petrogenetic relationships between rock suites, and the metamorphic style are all strongly reminiscent of Archean granite-greenstone terranes. As such, the southern Alaska forearc represents one of the world’s best modern analogs to early stages in the evolution of Archean granite-greenstone terranes. Belts of flysch, mélange and accreted ophiolites are described, and particular attention is paid to details of the geology of the 57 ± 1 Ma Resurrection Peninsula ophiolite as a remarkable analog to some Archean greenstone belts. The Resurrection ophiolite formed in a neartrench environment as the Kula-Farallon ridge was being subducted beneath North America. The magmatic sequence includes pillow lavas, sheeted dikes, gabbros, trondhjemites, and a poorly-exposed ultramafic section. The lavas show mid-ocean ridge basalt and arclike geochemical signatures, interpreted to reflect compositionally diverse melts derived from near-fractional melting of a variably depleted mantle source, mixed with variable amounts of assimilated continentally-derived flysch. A sedimentary sequence overlying the ophiolite preserves a continuous record of turbidite sedimentation deposited on the ophiolite as it was transported to North America and emplaced in the Chugach accretionary prism. The top of the sedimentary section is truncated by the Fox Island shear zone, a 1 km thick, greenschist-facies, west-over-east thrust related to the emplacement of the ophiolite into the accretionary wedge. The Fox Island shear zone is intruded by a 53.4 ± 0.9 Ma granite, showing that the ophiolite formed, was transported to the North American continent, overthrust by a major accretionary prism-related thrust, and intruded by granite all within 3.6 ± 1.4 Ma. Geological relationships in the southern Alaska forearc are instructive, in that if similar relationships were found in an Archean granite-greenstone terrane, they would probably currently be interpreted to reflect calc-alkaline mafic-felsic volcanic-plutonic complexes intruded and erupted through a complex metasedimentary sequence. Many Precambrian forearc ophiolites and accretionary prisms may have gone unrecognized because the processes of forearc ophiolite emplacement and intrusion by near-trench magmas at triple junctions has been poorly documented. In the final chapter of the book, A.M. Celail Sengör and Boris Natal’in synthesize Phanerozoic analogues of Archean oceanic basement fragments in the Altaids, distinguishing between nearly complete ophiolites, and severely dismembered ophiolitic bodies they term ophiorags. Sengör and Natal’in note that in the minds of most geologists, orogenic belts are linear/arcuate, long and narrow zones of intense deformation. That is why, irregularly shaped areas of widespread “orogenic deformation” interspersed with abundant fragments of the members of the ophiolite association in the Precambrian, but especially in the Archaean, have been thought as products of processes no longer operative. However, the geology of the Altaid orogenic system in Asia greatly resembles in its overall map aspects, lithological content, structural characteristics, and in the distribution and types of fragments of floors of former oceans to the Archaean granite-greenstone terrains. In the Altaids, ophiolites are now encountered in three main settings: (1) Ophiolites that occur as basement of ensimatic arcs, (2) ophiolites and ophirags that occur in former forearcs now entrapped within transform sutures: (a) Ophiolites as backstop to accretionary wedges, (b) Ophirags within accretionary wedges, and (3) Ophiolites and ophirags in collisional

References

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suture zones that have usually evolved from members of the second category. Ophiolites and ophirags have a widespread distribution within the orogenic edifice. This distribution was brought about by processes that shaped the Altaid edifice, namely, generation of suprasubduction zone forearc basements created by pre-arc spreading, back-arc basin opening, subduction-accretion, trench-linked strike-slip faulting including arc slicing and arc shaving faults and associated ocean floor spreading processes, and collision of buoyant pieces along suture zones. Without appreciating the nature and sequence of these processes and their superimposition, it is impossible to understand the rules that govern the distribution of oceanic basement fragments in the Altaids. These processes have led to a tremendous degree of structural shuffling of previously distant environments and a large degree of dismembering of formerly more complete ocean floor fragments. The preservation is highly selective and favors upstanding and buoyant segments of ocean floors. Such pieces are embedded most commonly in metapelitic/metapsammitic or, more rarely, in serpentinitic matrices in mélange/wildflysch complexes. Sengör and Natal’in contend that the same rules apply to the greenstone belts of the Precambrian and greatly hinder their deciphering in the absence of biostratigraphic control.

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Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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THE JORMUA OPHIOLITE: A MAFIC-ULTRAMAFIC COMPLEX FROM AN ANCIENT OCEAN-CONTINENT TRANSITION ZONE P. PELTONENa AND A. KONTINENb a Geological Survey b Geological Survey

of Finland, P.O. Box 96, FIN-02151, Espoo, Finland of Finland, P.O. Box 1237, FIN-70211, Kuopio, Finland

The Jormua Ophiolite is an allochtonous mafic-ultramafic rock complex, thrusted onto the Karelian Craton margin, that formed within a passive margin environment ∼ 100 km southwest from its present position. This complex consists of two distinct units: (1) fragments of Archean subcontinental lithospheric mantle that became exposed from beneath the Karelian craton by detachment faulting following the final break-up of the craton, and (2) alkaline and tholeiitic igneous suites that were emplaced within and through the lithospheric mantle at ∼ 2.1 Ga and 1.95 Ga, respectively. At the prerift stage of continental breakup (c. 2.1 Ga), residual lithospheric peridotites became intruded by alkaline melts that formed “dry” clinopyroxene cumulate dikes. Slightly later, this same piece of mantle became extensively intruded by hydrous alkaline magmas that resulted in formation of high-pressure hornblendite-garnetite cumulates deep in the ophiolite stratigraphy and fine grained OIB-type dikes at more shallow levels. Simultaneously, the residual peridotites became metasomatized due to porous flow of the melt in the peridotite matrix. Alkaline magmatism was soon followed by lithospheric detachment faulting that exposed the subcrustal peridotites at the seafloor, where they at 1.95 Ga became covered by tholeiitic (EMORB) pillow and massive lavas and intruded by coeval dikes and gabbros. Since transitional contacts between all main ophiolite units can be demonstrated, the Jormua Ophiolite Complex is interpreted to represent a practically unbroken sample of seafloor from an ancient ocean-continent transition (OCT) zone, strikingly similar to that reported from younger similar tectonic settings, such as the Cretaceous West Iberia nonvolcanic continental margin.

1. INTRODUCTION Though included only as a connotation in the current de facto ophiolite definition (Anonymous, 1972), it is generally accepted that ophiolite complexes represent fragments of oceanic lithosphere that formed in a number plate tectonic settings (oceanic spreading ridges, island arcs, back arc basins, leaky transforms, nascent ocean basins, etc.). The DOI: 10.1016/S0166-2635(04)13001-6

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special importance of Precambrian ophiolites is that their presence or absence in the rock record is usually considered one of the critical evidences for or against operation of modern type plate tectonism in Precambrian. In addition, the chemical composition of associated basalts provide an uncontaminated window into the Precambrian convective mantle, and in rare cases—such as Jormua—the petrology of the mantle rocks can be studied in situ from excellent outcrops. The 1.95 Ga Jormua complex was the first early Proterozoic ophiolite ever reported (Kontinen, 1987). Since then a few early Proterozoic and even Archean ophiolites have been discovered and described but still their number is notably low (Scott et al., 1992; Dann, 1991; Kusky et al., 2001). Being in many respects similar to the Northern Apennines ophiolites (e.g., Rampone and Piccardo, 2001) and sharing several of the salient features of the modern oceancontinent transition zones (Louden and Lau, 2001) the Jormua Ophiolite Complex has been interpreted as a break-up-related “passive margin ophiolite” (Kontinen, 1987; Peltonen et al., 1996, 1998). It is made up of two components of distinct origin: (a) subcontinental (> 2.8 Ga) lithospheric mantle (SCLM) component, and (b) asthenospheric component that consists of remnants of the 1.95 Ga mantle diapir and various types of igneous rocks emplaced within and through the mantle tectonites ∼ 1.95–2.1 Ga. Thus, the main importance of Jormua for understanding the evolution of Precambrian plate tectonic processes is the evidence it provides of continental rifting and break-up related tectonic and magmatic processes. The other ophiolitic rocks in Finland—the Outokumpu and Nuttio complexes (Fig. 1)—which appear to share the same large-scale tectonic setting along the Karelian Craton margin, provide us samples from more mature oceanic basin and island-arc setting, respectively (Koistinen, 1981; Vuollo and Piirainen, 1989; Hanski, 1997). Together, in a synthesis that remains to be compiled, these occurrences provide a tantalizing opportunity to interpret the break-up of the Karelian Continent, and the igneous processes that took place in the Svecofennian ocean from its birth until its closure. This contribution is intended to be a compact overview of the geology of the Jormua Ophiolite Complex and a state-of-art summary of what is known and inferred about its genesis. Emphasis is placed on the field description of the main ophiolite units. Chemical and isotopic compositions of both the crustal and mantle rocks have extensively been recently described elsewhere (Peltonen et al., 1996, 1998) and only some salient points shall be touched in this context. For the same reason no analytical data is included. There are still many lines of research that have not yet been applied to this rare piece of ancient oceanic and continental lithosphere. We encourage more research on Jormua so that more of its messages from the geological past will be revealed. In the meantime, we believe that the tectonic framework upon which the future work will bounce is already well established.

2. THE REGIONAL SETTING OF THE JORMUA OPHIOLITE COMPLEX The Jormua Ophiolite Complex is the northernmost and the most completely preserved example of the ophiolite fragments within the Paleoproterozoic North Karelia Schist Belt

2. The Regional Setting of the Jormua Ophiolite Complex

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Fig. 1. Geological map of the eastern part of the Fennoscandian Shield emphasizing the location of the Jormua Ophiolite Complex and other early Proterozoic ophiolite fragments. Archaean blocks: PC = Pudasjärvi Complex; IC = Iisalmi Complex; EFC = Eastern Finland Complex. Modified from Koistinen et al. (2001) and Hanski (1997).

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Chapter 1: The Jormua Ophiolite

and the Kainuu Schist Belt in the central part of the Fennoscandian Shield (Fig. 1). These two Paleoproterozoic schist belts, which consist mainly of 2.3–1.90 Ga metasediments (grouped under the term Karelian in the Finnish literature), are located 0–100 km to the east of the suture between their Archaean basement structure (Karelian Craton) and the 1.93–1.80 Ga Svecofennian island arc domain (Fig. 1). Approximately 1.9 Ga ago the southwest margin of the Karelian Craton was covered by thrusted allochthonous complexes and deformed by the related compressional and subsequent wrench/thrust tectonics. The thrusting took place before the D1 deformation as a response to the collision between the Karelian Craton and the cratonized Svecofennian arc collage (Gaál and Gorbatschev, 1987; Kärki and Laajoki, 1995; Kohonen, 1995; Korsman et al., 1999). The thrust belt contains a 200 km long chain of ophiolite fragments whose distribution is related to the early thrusting, further modified by the later multistage regional deformation (Koistinen, 1981; Kontinen, 1987; Kärki and Laajoki, 1995). The Kainuu Schist Belt that encloses the Jormua Ophiolite occupies a structural depression between two rigid blocks of the basement structure: the Eastern Finland Complex and Pudasjärvi-Iisalmi Complexes (Fig. 1). The belt is up to 30 km wide but nowhere more than 2–3 km thick on the basis of structural and gravimetric data. It thus represents a relatively thin veneer of faulted and folded autochthonous and allochthonous supracrustal rocks and ophiolite fragments on the thick Archaean basement structure. The belt is dissected by a folded N-S trending strike-slip fault zone along which the Archean basement blocks have moved both vertically and horizontally (Kärki and Laajoki, 1995). The central part of the Kainuu Schist Belt around the Jormua Ophiolite comprises three major unconformity or thrust-bound lithofacies: (1) the autochthonous, cratonic to epicratonic Jatuli (2.3–2.1 Ga) sequence, consisting predominantly of quartzites derived from fluvial to shallow marine, mature quartz-rich sands; (2) the 2.1–1.95 Ga rift-phase related “lower Kaleva” assemblage, characterized by metaturbiditic conglomerates, quartz wackes, graywackes and shales, as well as turbidite-hosted P-Mn-C-rich silicate-carbonate iron formations and abundant, metal-rich (Cu, Ni, Zn) graphitic black schists; and (3) the allochthonous “upper Kaleva” sequences, dominated by deep marine metaturbiditic graywacke-shale deposits whose depositional age is bracketed by the age of the Jormua (∼ 1.95 Ga; Peltonen et al., 1998) and the youngest detrital zircon SHRIMP ages of 1.92 ± 0.12 Ga (Claesson et al., 1993). Many of the detrital zircons in these “upper Kalevian” metagraywackes are thus younger than the mafic sequence of the Jormua Ophiolite. The Jatulian mature arenites, and the “lower Kalevian” sequence, which are presumably rift-related deposits contain mainly recycled Archean detritus, whereas the younger, flyschlike “upper Kalevian” metagraywackes and shales contain also abundant Proterozoic material. An additional important geological unit, probably closely related to the origin of the Jormua Ophiolite Complex, is the ∼ 1.96 Ga Otanmäki gneissic, peralkaline-alkaline Atype granite to the SSW of Jormua (Fig. 1). These granites have intruded Archaean gneisses and Jatuli-type cover sediments and are present as a narrow strongly foliated tectonic slice. At their southern margin the Otanmäki gneissic granites are in a faulted contact with the rocks of the Kainuu Schist Belt and the Archaean rocks of the Iisalmi Complex. We are

3. Principal Features of the Main Blocks

39

tempted to believe that the Otanmäki gneissic alkaline granite is an allochthonous unit not belonging to the basement structure of the Kainuu Schist Belt.

3. PRINCIPAL FEATURES OF THE MAIN BLOCKS The earliest stages of the Svecofennian deformation, related to the early tectonic processes that contributed to the detachment of the Jormua Ophiolite Complex from oceanic environment and its subsequent thrusting across the foreland, involved tectonic disruption of the original ophiolite assemblage. Consequently, the Jormua Ophiolite now consists of four major fault-bounded blocks (Fig. 2), which represent diverse parts through the ancient ocean-continent transition zone (Peltonen et al., 1996, 1998). Extensive shearing along the Kainuu Schist Belt in the latest stages of the Svecofennian tectonism (Kärki and Laajoki, 1995) and associated parasitic faulting and folding further disrupted and deformed the ophiolitic blocks, some of which now have forms of shear-controlled “megaaugens”. The salient characteristics of these blocks are summarized in Table 1. The eastern block consists of several fault-bounded slices of serpentinized mantle tectonites (mainly harzburgites) and minor dunitic pods, some of which enclose small nodularand orbicular-textured podiform chromitite lenses. Mantle peridotites have been intruded by gabbro stocks and basaltic sheeted dike complexes. The eastern block is practically the only block that is associated with the extrusive seafloor sequence. Reconstructed stratigraphy of the ophiolite suggests that the dike complex and locally also mantle peridotites are directly overlain by metabasaltic pillow lavas intercalated with some massive lavas and pillow breccias. According to diamond drilling data the pillow lavas at the western margin of the eastern block are overlain by basic tuffs intercalated with sedimentary carbonates. These metatuffs are—at least in a tectonostratigraphic sense—overlain by the typical upper Kaleva metaturbiditic black shales and graywackes. Thus, in terms of igneous stratigraphy, the eastern blocks expose the most complete ophiolite within the Jormua igneous complex. The northern block is scantily exposed and hence the least studied of the Jormua blocks (Fig. 2). Nevertheless, the presently available data suggests that this block is fairly similar with the eastern block. The main exposed components are mantle tectonites, gabbroic feeder dikes, and more extensive gabbro-diorite intrusions and some massive metabasaltic amphibolites, probably of sheeted dike origin. Highly altered, epidote and iron sulfide-rich gabbros are known from one locality. The central block resembles the eastern block in some respects but several fundamental differences are evident (Table 1). Both extrusive rocks and gabbroic pods and dikes are uncommon within the central block. Instead, the mantle tectonites have been intruded by abundant fine grained EMORB dikes. The abundance of these dikes progressively increases from NW to SE (Fig. 2). In the northwest single anastomosing dikes intrude peridotites. In the following, these dikes are referred to as “deep dikes” because of their apparent location deep in the ophiolite stratigraphy. Towards southeast “deep dikes” gradually coalescence into thicker units finally forming a spectacular sheeted dike complex with only rare interdike screens of mantle peridotites or gabbro. The mantle peridotites (serpentinites) of the

40

Chapter 1: The Jormua Ophiolite

Fig. 2. Geological map of the Jormua Ophiolite Complex. Modified after Kontinen (1998).

3. Principal Features of the Main Blocks

41

Table 1. Characteristics of the Jormua Ophiolite complex Rock type

Western block

Central block

Eastern block

Northern block

Crustal units Extrusive unit Sheeted dike complex Gabbros/plagiogranites Ultramafic cumulates

− − − −

− ++ − −

++ ++ ++ ?

+ + + −

Intrusive to mantle tectonites EMORB-dikes Gabbroic feeder dikes Chromitite pods OIB-type dikes Clinopyroxenite mantle dikes Hornblendite mantle dikes Garnetite veins Carbonatitic veins

? − − ? ++ ++ + +

++ + − + + − − −

++ ++ + − − − − −

+ + − − − − − −

Mantle tectonites Lherzolites (> 3 wt% Al2 O3 ) Depleted lherzolites (1
Al2 O3
3 wt%) Harzburgites and dunites (< 1 wt% Al2 O3 )

+ ++ −

− ++ +

+ ++ +

? + +

+, present; ++, abundant; −, absent.

central block are intruded by an additional generation of ultramafic-mafic dikes which are absent in the eastern block. These dikes, which have OIB-type geochemical characteristics and resemble ultramafic lamprophyres, are intruded by the more voluminous EMORB dikes and are therefore labelled as “early dikes”. While the EMORB dikes intersect the mantle tectonite foliation at high angles the general trends of the OIB-type dikes are subparallel. This clearly implies that EMORB and OIB magmas were emplaced during distinct episodes, and that between the emplacement of OIB and EMORB dikes, the tectonite foliation was rotated from subvertical to almost horizontal. The contact between the western block and central blocks is unexposed. Thus, the possibility remains that they form a single long continuous sheet, but their significant internal differences suggest that this is probably not the case (Fig. 2). First, the western block is not associated with “upper Kaleva” graywackes, but instead is bounded by slices of the Archaean basement and rift-related “lower Kaleva” sediments. Second, the mantle peridotites are less depleted in their basaltic constituents and harzburgitic and dunitic residues are uncommon. Furthermore, gabbros, basaltic dikes and pillow lavas are absent from the western block. Instead, mantle peridotites are extensively intruded by clinopyroxenite and hornblendite dikes and pods representing igneous cumulates that crystallized in melt channels and pathways within the upper mantle. Isotopic data suggest that these cumulates are broadly coeval and closely related to the OIB-type dikes found in the central block

42

Chapter 1: The Jormua Ophiolite

(Peltonen et al., 1998). Overall, the western block shares more common features with orogenic lherzolite massifs (i.e., subcontinental lithospheric mantle) than with true ophiolites.

4. RECONSTRUCTED STRUCTURE OF THE COMPLEX Although the Jormua Ophiolite Complex has been tectonically disrupted into several distinct blocks, the reconstruction of the original igneous stratigraphy can be made with reasonable confidence. This is because transitional contacts between the main ophiolitic units, i.e., extrusive rocks, sheeted dikes complex, gabbros and mantle peridotites can be demonstrated in the field. A revised stratigraphic reconstruction for the Jormua Ophiolite Complex is presented in Fig. 3. The upper part of the stratigraphic column is based on the geology of the eastern and central blocks of the Jormua Ophiolite Complex. The extrusive unit consists of pillow and massive lavas capped by some basic tuffite and sedimentary carbonates. The extrusive unit is relatively thin (0–400 m) and one may argue that locally lavas were deposited directly onto mantle tectonites exposed at the seafloor. In modern seafloor, mid-ocean ridges with mantle peridotite outcrops are characterized by slow spreading rates and deep axial valleys that are expected to form at magma-poor ridge regions where a substantial fraction of the oceanic lithosphere is made of tectonically uplifted mantle material (Cannat, 1993). Downwards in the stratigraphic section, pillow lavas occur as interdike screens within the uppermost sheeted dikes complex, and some dikes are present within the lava unit implying coeval formation of lavas and dikes. Thick mafic-ultramafic cumulate layers are absent from the Jormua Ophiolite Complex. This is, however, not because of incomplete preservation of the original sequence, but is a characteristic feature of slow spreading-ridge ophiolites where thick axial lithosphere prevents magma pooling at crustal depths. In Jormua, the sheeted dike complex is not rooted in the cumulates but instead is rooted in the mantle tectonites (Figs. 2 and 3). Going downwards the coherent sheeted dike complex gradually gives way to dike swarms with abundant mantle tectonite interdike screens and finally into sparse anastomosing basaltic dikes (“deep dikes”) intruding mantle tectonites. Furthermore, indirect evidence for the absence of large magma chambers in Jormua is provided by the geochemistry of the basalts; the analysed lava and sheeted dike samples cannot be related to each other by fractional crystallization processes but represent rather unmodified melt fractions which did not undergo fractionation at an intermediate magma storage (Peltonen et al., 1996). Isotropic gabbro (+ plagiogranite) pods characterize the middle section of the stratigraphic column (Fig. 3). Importantly, these gabbros are frequently cross-cut by sheeted dikes, but are never observed to intrude the dike complex. This suggests that gabbros and seafloor magmatism represent distinct magmatic episodes and that the U-Pb zircon ages provided by the gabbros give a maximum age for the seafloor volcanism. However, the absolute time difference is believed to be small, with the gabbro stocks being related to the initial stages of the continental rifting, while lavas and dikes record a slightly more advanced stage of the oceanization. Distinct types of gabbros are found beneath the large high-level gabbro-plagiogranite pods. These occur in the form subvertical dikes that

4. Reconstructed Structure of the Complex

43

Fig. 3. Igneous stratigraphy of the Jormua Ophiolite Complex. Age data from Huhma (1986), Kontinen (1987), Peltonen et al. (1998), and Peltonen et al. (2003, unpublished).

intrude the mantle tectonites (hereafter: gabbroic feeder dikes). Their internal fractionation is indicative of them being feeder dikes for upper-level magmas. They have yielded equal crystallization age (1953 ± 2 Ma; Peltonen et al., 1998) with the gabbro stocks (1960 ± 12 Ma) and plagiogranites (1954 ± 11 Ma; Kontinen, 1987) and it is not evident whether they are comagmatic with high-level gabbro stocks or with sheeted dikes or lavas. The lower part of the stratigraphic column refers to the western block of the complex. Since the western block is lithologically distinct from other blocks and the contact between the western and central block is unexposed a continuous stratigraphic column was not drawn. This part of the Jormua Ophiolite Complex does not resemble a true ophiolite but bears striking similarities with fragments of subcontinental lithospheric mantle (SCLM), such as orogenic lherzolite bodies of the French Pyrenees. Thus, the Jormua Ophiolite consists of two distinct parts; one similar to slow-spreading type ophiolites (oceanic lithosphere) and another that is similar to orogenic lherzolites (subcontinental lithospheric mantle). Importantly, juxtaposition of these parts is not coincidental since sim-

44

Chapter 1: The Jormua Ophiolite

ilar c. 1.95 Ga old dikes facies are present in both fragments. This suggests that these two distinct parts share a common history and led Peltonen et al. (1998) to suggest that as a whole the Jormua Ophiolite Complex records an almost continuous sequence across an ocean-continent transition (OCT) zone.

5. ALTERATION AND METAMORPHISM Seafloor metamorphism, alteration during obduction and tectonic transport, and finally the Svecofennian regional metamorphism, have together resulted in extensive destruction of the primary mineralogy of the mafic and ultramafic lithologies of the Jormua Ophiolite. Fortunately, however, chemical changes with respect to most elements are far less pronounced. The metamorphic history of the ophiolite commenced already at the seafloor stage—evidence of which has mostly been lost by later imprints. Hydrothermal circulation on the seafloor resulted in alteration of basalts and veining of the basaltic dikes by albite-filled fractures (Fig. 4f). Lavas, in turn, are anomalously depleted in Fe, a feature that has been related to the seafloor weathering (Peltonen et al., 1996). It is not clear anymore to what extent the mantle peridotites were serpentinized at this stage. However, since the basaltic lid is thin in Jormua and field evidence suggests that some mantle tectonites were exposed at the seafloor, it is likely that mantle peridotites became at least partially serpentinized before obduction. The alteration that took place during the obduction and tectonic transport is only poorly characterized. The serpentinization of the mantle tectonites continued due to their interaction with meteoric waters. This alteration proceeded at relatively low temperature conditions resulting in extensive replacement of olivine and pyroxenes by pseudomorphic lizardite and local carbonatization and silicification of the margins of the tectonite massifs. Basaltic and gabbroic dikes in contact with peridotites became rodingitized due to their interaction with serpentinizing fluids. This resulted in considerable loss of silica and alkalies from the dikes (Fig. 8) and replacement of the primary igneous mineral parageneses of the gabbro dikes by hydro-grossular garnet, diopside, epidote, and chlorite. Much of the present mineral parageneses of the Jormua mafic and ultramafic rocks, however, represent the metamorphic equilibria attained during the Svecofennian regional metamorphism. Within the Kainuu Schist Belt metamorphism culminated under low-P high-Ttype conditions in the amphibolite facies between 1.87 and 1.85 Ga. The cooling was a prolonged event and it was not until ∼ 1.80 Ga that temperatures fell below 500 ◦ C. Thus, the metamorphism significantly outlasted the deformation that was essentially over by 1.86 Ga (Tuisku, 1997). The metamorphic mineral paragenesis in the metabasalts typically is sodic plagioclase + actinolitic hornblende ± epidote ± chlorite. Due to the activity of CO2 -rich metamorphic fluids some lava samples became slightly depleted in LREE. This is evident in the Sm-Nd isochron diagram where basalt samples form an isochron with a slope corresponding to an age of 1.72 ± 0.12 Ga (Peltonen et al., 1996). The slope of this isochron is strongly controlled by the three pillow-lava samples having the most LREE-depleted patterns. Since this “errorchron” age is significantly less than the U-Pb zircon age (∼ 1.95 Ga)

5. Alteration and Metamorphism

45

(a)

(b)

(c)

(d)

(e)

(f)

Fig. 4. Outcrops of (a) pillow lava, (b) pillow breccia, (c) “deep dikes”, i.e., EMORB dikes intruding mantle tectonites at the root zone of the sheeted dikes complex. Note the anastomosing shape of the dikes and the dark alteration selvages at their margins due to interaction with adjacent serpentinites (former residual mantle peridotites), width of the photo ∼ 2 m, (d) sheeted dikes complex with 100% dike-in-dike sets, width of the photo ∼ 3 m, (e) sheeted dikes with gabbro interdike screens, (f) “deep dike” with seafloor alteration-related albite veins overprinted by late alteration selvage between dike and enclosing mantle tectonite at left, width of the photo ∼ 30 cm. (f) Reprinted with the permission from Journal of Petrology, vol. 37, Oxford Univ. Press.

46

Chapter 1: The Jormua Ophiolite

of the Jormua Ophiolite, or the age of the obduction (∼ 1.90–1.87 Ga), it is clear that the LREE depletion of these three samples cannot be of igneous origin. In mantle tectonites, the regional metamorphism resulted in dehydration and replacement of the earlier low-T pseudomorphic lizardite serpentine by non-pseudomorphic antigorite serpentine. Present stable metamorphic parageneses of serpentinites are dependent on the bulk rock composition and include antigorite+magnetite and antigorite+tremolite+magnetite in the eastern and central block serpentinites. Olivine is added to the parageneses in the western block. On the basis of published mineral stability curves (e.g., Will et al., 1990) we estimate that these parageneses imply peak-metamorphic temperatures of approximately 480 and 530 ◦ C for pressures of 2 and 5 kb, respectively, in the west and slightly lower temperatures in the east. As a result of nearly complete serpentinization, the analysed H2 O(tot)-contents of peridotite samples now range between 9.4 and 12.0 wt%. However, since SiO2 /MgO ratios of Jormua serpentinites are still very similar to fresh or only slightly serpentinized mantle peridotites elsewhere, it is highly probable that serpentinization largely conserved both SiO2 and MgO, in which case the volume of the peridotites must have increased (O’Hanley, 1996). Within the Jormua Ophiolite Complex talc-carbonate rocks occur as altered marginal variants of serpentinite massifs. Their mineralogy is dominated by carbonate and talc in approximately equal proportions, together with some magnetite and sulfides. Extensive talc-carbonate alteration is restricted to narrow marginal zones of serpentinites. Talccarbonate alteration may be a post-metamorphic process (Eckstrand, 1975) or, alternatively, antigorite-carbonate-talc and carbonate-talc assemblages could have been stabilized under prograde conditions but at significantly higher XCO2 than the carbonate-free mineral assemblages. Similar metamorphic zonation in serpentinite bodies has been described from metakomatiites (Gole et al., 1987). By analogy, the concentric metamorphic mineralogy of serpentinite massifs could have been created when the breakdown of premetamorphic serpentines (XH2 O = 1) and the infiltration of serpentinite margins by CO2 rich fluid (originating from adjacent calcareous metasediments and pre-metamorphic alteration zones undergoing decarbonatization) generated gradients in the composition of metamorphic fluid. Carbonate-free assemblages represent equilibration beyond the limit of CO2 -infiltration. Peltonen et al. (1998) discussed in length the mobility of REE during serpentinization or deserpentinization reactions. They came to the conclusion that in the absence of suitable ligands (such as sulfate or carbonate) in the fluids, the alteration— although severe—did not seriously affect the REE characteristics of the peridotites.

6. THE CRUSTAL UNIT 6.1. Lavas and Hyaloclastites The extrusive crustal part of the Jormua Ophiolite is rather thin, approximately averaging only 100–400 m (Fig. 3). The thickness of the whole basaltic lid of the ophiolite that includes lavas, sheeted dikes and high-level gabbro stocks was variable (< 500 to > 1.5 km).

6. The Crustal Unit

47

Locally, it is apparent that the basaltic flows were deposited directly onto mantle peridotites. Most of the extrusive metavolcanic rocks of the Jormua Ophiolite Complex occur within the eastern block (Fig. 2; Table 1) as tectonic slices up to 400 m thick and kilometers in length, and typically in faulted contact with adjacent lithologies. Typically, lavas are bordered by serpentinites and talc-carbonate rocks or gabbros-plagiogranites in the structural footwall (to the west), and by upper Kalevian black schists and metagraywackes in the structural hanging wall (to the east). Most of the extrusive part of the Jormua consists of pillow lavas (∼ 50%; Fig. 4a) together with substantial amounts of pillow breccias (Fig. 4b) and hyaloclastites that make up to 25% of the extrusive unit. The remaining 25% consists of massive lava flows or flow parts. Some dikes and gabbro intrusions, one > 15 m thick differentiated gabbro-pyroxenite sill, for example, are present within the extrusive sequence. Importantly, interstitial or intercalated terrigenous sedimentary material is completely absent. Based on the presence of hyaloclastite interpillow matrix, minor pillow breccias, and the vesicle-rich nature of some of the pillowed flows Kontinen (1987) argued that the lavas erupted in a relatively shallow-water environment. Relict porphyritic and glomeroporphyritic textures are commonly recognizable in lavas which at present consist essentially of recrystallized plagioclase and nematoblastic calcic amphibole. The chemical composition of the lavas, such as their high Mg# [Mg/(Mg+Fe2+ tot ) = 0.59 to 0.73] and high Cr and Ni abundances suggest that lavas are not strongly fractionated. On the Cr vs. Y diagram (Fig. 5), for example, lava samples plot along or only slightly below the partial melting trend suggesting that the primary magma has undergone only minor fractionation (Fig. 5). Since chromite saturates early in MORB (Fisk and Bence, 1980) fractionation should rapidly deplete the residual melt in chromium. Most lava samples cannot be related to each other (or to the sheeted dikes) by fractional crystallization. Instead, the chemical composition of the lavas is consistent with them representing individual melt fractions produced through variable degrees of partial melting from which only minor amounts of chromite, olivine, and plagioclase have been segregated. Such chemical characteristics are in perfect agreement with the igneous stratigraphy inferred from field observations that demonstrated the absence of large magma chambers where pre-eruptive fractionation could have taken place. The trace element composition of Jormua basalts has been extensively discussed by Kontinen (1987) and Peltonen et al. (1996). Lavas (and most of the dikes) form a rather homogeneous group that has a composition closely similar to that of enriched mid-ocean ridge basalts (EMORB) of the modern seafloor. They are characterized by flat REE (Fig. 6) and, e.g., Nb, Ta, and Th abundances in excess of that typical for NMORB. On the basis of trace element and Nd isotope data Peltonen et al. (1996) concluded that the chemical composition of Jormua lavas and sheeted dikes result from mixing of NMORB and OIB mantle sources. Furthermore, on the Th/Yb vs. Ta/Yb diagram (Pearce, 1983) lava (and dike) samples form a coherent group of analyses that plot along the diagonal mixing array between depleted and enriched mantle sources. This implies that they all contain broadly uniform amounts of enriched endmember, and are free of any geochemical crustal (arctype) signature.

48

Chapter 1: The Jormua Ophiolite

Fig. 5. Cr vs.Y plot for lavas and basaltic dikes of the Jormua Ophiolite. Mineral vector calculations after Peltonen et al. (1996). The subhorizontal line is the partial melting trend of Pearce (1982).

6.2. Sheeted Dike Complex The sheeted dike complex, with its > 1 km thickness and ∼ 10 km2 areal coverage, is the major crustal unit of the Jormua Ophiolite. The central block (Fig. 2) emphasizes one of the most spectacular phenomena of Jormua: the “rooting” of the sheeted dike complex in the uppermost mantle. Stratigraphically the lowest part of the complex consists of individual anastomosing basaltic dikes (“deep dikes”) that intrude the mantle tectonite (Fig. 4c). Margins of thick “deep dikes” are characterized by 10–30 cm wide, dark green, chlorite rich alteration selvages while thin “deep dikes” may be thoroughly altered. Compared to dike interiors the alteration margins are strongly depleted in, e.g., Si, Ca, and alkalies, and enriched in Mg, Cr, Ni, and loss of ignition reflecting their pervasive hydration and double diffusive interaction with mantle peridotite (Peltonen et al., 1996). This interaction took place either during serpentinization of the peridotite host (rodingitization), or later during regional metamorphism, or both. Moving upwards in the stratigraphy “deep dikes” coalescence into groups of multiple dikes with progressively smaller mantle tectonite and gabbro interdike screens demonstrating the consanguinity of “deep dikes” to the coherent sheeted dikes complex. Presence of lava, gabbro and mantle tectonite as interdike screens

6. The Crustal Unit

49

Fig. 6. Chondrite-normalized (Boynton, 1984) rare earth element plots. (a) Lavas and basaltic dikes. Two distinct basalt suites are present: EMORB-type basalts with flat REE patterns and “early suite” OIB-type lavas with fractionated HREE and high LaN/YbN. (b) Gabbros and plagiogranites of the Jormua Ophiolite. The gabbros include samples from both from high-level gabbro stocks and gabbroic feeder dikes which have similar patterns and equal range of absolute concentrations.

(septa) in the Jormua sheeted dikes complex demonstrate that the contacts between the main ophiolite units are transitional, as required for ophiolite recognition by, e.g., Bickle et al. (1994).

50

Chapter 1: The Jormua Ophiolite

At higher crustal levels the sheeted dike complex consist of 100% of subparallel metadolerite and metabasalt dikes (Fig. 4d). Dikes in the sheeted complex are generally 20– 120 cm thick, aphyric or plagioclase-phyric with sharp chilled mutual contacts. The older dikes, which tend to be thick and coarse grained, were intruded by thinner and finer grained younger dikes, and became transformed to interdike screens of “half”, “one-way chilled”, and “marginless” dikes. Branching and apophyses along dike margins are common. Some of the dikes contain abundant plagioclase concentrated by magmatic flow in the dike interiors. No other microtextures have survived the recrystallization during regional metamorphism. The occasional presence of widely separated serpentinite and gabbro septa attest in a striking way the magnitude of extension during the formation of the mafic lid of the Jormua Ophiolite Complex (Fig. 4e). Although the “deep dikes”, sheeted dikes and overlying basalts are clearly coeval magmatic rocks with broadly similar chemical and isotopic compositions (Kontinen, 1987; Peltonen et al., 1996), minor differences emphasize the intricacy of this magmatism. The Cr (ppm) vs. Y (ppm) diagram (Fig. 5) emphasizes that while the lavas represent almost unmodified melt fractions, most of the “deep dikes” and sheeted dikes are significantly more fractionated and cannot represent feeders for the lavas. Probably this is related to the rate of the magma flow in the dike. Those dikes that acted as feeders for the lavas were characterized by high flow rates which prevented extensive fractional crystallization of the melt and resulted in eruption of lavas with primitive composition. Most of the dikes are, however, “blind” and never fed any basalts but instead underwent fractionation during ascent. 6.3. Gabbros A thick layered gabbro unit, characteristic of many classic young ophiolites is distinctly absent in the Jormua Ophiolite Complex. Instead, all gabbro occurrences in Jormua, even the largest ones (0.5 × 1.5 km in areal extent), represent stocks and dikes intrusive into the uppermost mantle. Most gabbro outcrops are made of isotropic or “varied-textured”, coarse to pegmatoid gabbro (Fig. 7a) and any clear modal layering or banding is absent even in the largest occurrences. Gabbro bodies in the central part of the eastern block are frequently intruded by fine grained to aphanitic, usually aphyric, typically 0.2 to 1.2 m wide basaltic dikes. These are separated by cm to meter wide gabbro screens (septa) that comprise 30–60% of many of the gabbro outcrops (Fig. 4e). Locally subparallel dikes form several meter wide zones of dike-in-dike sets. Many of the dikes show clear chilled margins against the gabbro screens and older dikes. Two types of gabbros have been distinguished in the field: grayish green Mg-gabbros, and dark green Fe-gabbros characterized by relatively low MgO, low SiO2 and distinctly high Fe and Ti reflected in abundance of ilmenite. Mineral assemblages in gabbros are usually thoroughly metamorphic: of the primary phases only An-rich plagioclase is sporadically preserved in the Mg-gabbros and coarse ilmenite in the Fe-gabbros. The most voluminous Fe-gabbros are present in the central part of the eastern block. Transitions from adjacent Mg-gabbros take place within a few meters by abrupt increase in ilmenite

6. The Crustal Unit

51

(a)

(b) Fig. 7. Outcrops of (a) varied-textured ilmenite-bearing high-level gabbro stock, and (b) leucotonalite (plagiogranite) dikes and veins in fine grained high-level gabbro.

content. Grain-size is characteristically variable ranging from fine to coarse and even pegmatoid over short distances. Irregular fine grained dike-like parts in some ferrogabbro outcrops suggest that coeval Fe-basalt dikes are rooting in the Fe-gabbro stocks. However, no Fe-rich lavas or dikes in the sheeted dike complex have been recognized so far. On the AFM ternary diagram gabbro samples follow a tholeiitic fractionation trend showing pronounced Fe (+ Ti) enrichment in the most evolved samples of the igneous suite (Fig. 8). Incompatible element abundances (such as Ti and Zr) are significantly lower in all gabbros compared to lavas or sheeted dikes of the complex (Kontinen, 1987). This implies that the gabbros are cumulates with only small amounts of intercumulus liquid remaining. Either substantial post-cumulus growth has taken place or, maybe more likely, the gabbroic cumulates have been depleted in residual liquid by filter pressing due to their crystallization in a dynamic environment. Furthermore, the compatible elements Cr and

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Chapter 1: The Jormua Ophiolite

Fig. 8. AFM triangular plot for gabbros and plagiogranites of the Jormua Ophiolite Complex emphasizing the alkali depletion of gabbroic feeder dikes due to their rodingitization and, on the other hand, the extreme alkali (sodium) enrichment of the plagiogranites. Discrimination line after Irvine and Baragar (1971).

Ni are also present in significantly lower concentrations than is common for the lavas and dikes with the same Mg level. This suggests that the parental magma for the gabbros was already relatively evolved and implies that they probably are not coeval with lavas or dikes. 6.4. Plagiogranites Plagiogranites (i.e., leucotonalites) are well exposed only at one locality in the eastern block where they are closely associated with gabbros, microgabbros and diorites. Two main types of occurrences are present. First, some plagiogranite is present as a few meters wide and at least several tens of meters long zones of multiple, successive, and apparently syn-magmatically deformed dike injections within high-temperature ductile shear zones in Mg-gabbros. Plagiogranite in these zones is present as networks of irregular, cm to dm

7. The Mantle Section

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thick, branching and cross-cutting dikes. Outcrop features suggest coeval emplacement of microgabbroic-basaltic dikes with the plagiogranite dikes and mingling of these magmas. Second, plagiogranite is also found as dike networks within microgabbro-diorite that occurs as marginal facies of large ferrogabbro bodies (Fig. 7b). Outcrop features suggest emplacement of the microgabbros-diorites and plagiogranites as a multistage progressive process involving mingling of the various pulses of magma. The leucotonalitic to trondhjemitic segregations (plagiogranites) of Jormua Ophiolite Complex have the chemical characters of ocean ridge granites being, e.g., very high in Na2 O and low in K2 O (Fig. 8). Chondrite normalized REE patterns of the leucotonalite segregations are somewhat fractionated at a relatively high level of REE abundances and have negative Eu minimas (Fig. 6b). The high Y and Nb concentrations of the dioritesleucotonalites places them in the within-plate granite field in the Nb vs. Y discrimination diagram of Pearce et al. (1984), which is in line with the overall EMORB character of the Jormua mafic rocks. Zircons from one leucotonalite segregation yielded a somewhat unprecise U-Pb age of 1954 ± 12 Ma. In addition, εNd (1.95 Ga) for this sample is +1.9, which is close to the +2 average for the lavas and dikes implying that plagiogranite origin is intimately related to the oceanic crust-forming magmatism.

7. THE MANTLE SECTION 7.1. Mantle Peridotites (Metaserpentinites) Mantle peridotites comprise approximately 55% of the area of the Jormua Ophiolite Complex. At present, the peridotites are mainly antigorite metaserpentinites whose mineralogical composition is controlled by bulk rock compositions and metamorphic grade. The eastern and central block peridotites are still characterized by outcrop textures typical for residual mantle tectonites, and by the absence of any well-defined textural or compositional layering. Serpentine pseudomorphs (bastite recrystallized to antigorite) after residual orthopyroxene occur as high relief lumps in the weathered outcrops (Fig. 9a). They are embedded within a “groundmass” consisting of somewhat darker serpentine + magnetite dust that is replacing mantle olivine. Locally, pyroxene pseudomorphs form mm to cm thick bands and strained chromite grains form schlierens thus defining the tectonite foliation (Fig. 9b). This foliation is cross-cut by 1950 Ma old gabbroic feeder dikes implying that the foliation is of mantle origin and not due to ∼ 1880 Ma regional deformation. Chromite is the only primary mineral that is preserved in the mantle peridotites (Liipo et al., 1995). Most of the disseminated chromite grains in the serpentinites have thoroughly been altered to ferri-chromite and chromian magnetite. Only occasionally do the serpentinites contain disseminated grains with unaltered chromite cores. In these grains the mutual boundary between the chromite core and ferri-chromite outer shell is sharp both optically and chemically. Ferri-chromite in turn gradually grades to chromian magnetite towards the grain margin. Electron microprobe analyses imply that at the scale of hand samples, there is only minor intra and inter grain variation in Cr# [Cr/(Cr + Al)] of

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(a)

(b)

(c)

(d)

Fig. 9. (a) “Knobby”-textured serpentinite (former residual harzburgite) where bastite pseudomorphs after orthopyroxene remain as high-relief knots in the antigorite serpentine matrix after mantle olivine, (b) chromite (now largely chromian magnetite) striations define mantle foliation, (c) residual mantle tectonite infiltrated by basaltic melt. Such samples are responsible for the relatively high REE abundances in some peridotite samples and their flat chondrite-normalized patterns (Fig. 11a), (d) small podiform chromitite body enclosed by talc-carbonate altered eastern block mantle tectonite. (a) Reprinted with the permission from Journal of Petrology, vol. 39, Oxford Univ. Press.

the chromite cores. Mg# [Mg/(Mg + Fe2+ )], in turn, varies greatly being higher in the interiors of the large chromite cores and lower in interiors of the small ones and gradually decreases from the chromite cores towards their margins (Kontinen and Peltonen, 1998). This implies that Mg# have been strongly modified from the original mantle values by alteration and metamorphism. However, the apparent immobility of Cr and Al in the chromite cores suggests that the variation in the Cr# (45–70) may still closely reflect the primary mantle melting-controlled values. Relatively high and variable Cr# suggests that the eastern and central block peridotites represent residues after variable but generally high degrees of partial melting. Chemical composition of the metaserpentinites imply similar origin as was deduced from the compositions of chromite. On the Pd/Ir vs. Ni/Cu metal ratio diagram of Barnes et al. (1988) talc- and carbonate-free serpentinite samples form a tight cluster close to the

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mantle field (Fig. 10a). Two talc-carbonate altered samples are displaced from the mantle field: a serpentinite-talc schist contact zone sample (0.73 wt% CO2 ) contains 170 ppm Cu (> 5 times that of primitive mantle) and one of the massive serpentinites (1.72 wt% CO2 ) is strongly depleted in Pd. Uniform and low Pd/Ir is inconsistent with a cumulate origin for any of these samples. Due to similar partition coefficients between melt and residual sulfides during partial melting the Pd/Ir is not sensitive to variations in the degree of mantle melting. This is, instead, illustrated by Cr/Al vs. Ni/Al plot (Fig. 10b). In this diagram— where both ratios increase as a function of mantle melting—Jormua serpentinites record a wide range of ratios indicating variable degrees of melt extraction from peridotites. Incompatible element abundances bear evidence for a complex post-melting history of the Jormua mantle peridotites. Chondrite-normalized REE patterns bring out major differences between peridotites from different blocks of the Jormua Ophiolite Complex (Fig. 11). Serpentinites from the eastern block yielded two distinct types of patterns. First, several samples are characterized by U-shaped chondrite-normalized patterns—a form that is typical for dunites and harzburgites from the basal sections of ophiolites elsewhere (e.g., McDonough and Frey, 1989). Such patterns indicate that peridotites have first been depleted in LREE and MREE during mantle melting and later enriched in LREE. Second, some eastern block samples contain much more REE and yield flat patterns, similar to those of Jormua EMORB, but at a lower level. Some of these samples may represent small dunitic cumulate pods but some are characterized by typical residual mantle outcrop textures and more likely represent residual peridotites with substantial amounts of infiltrated and trapped basaltic melt (Fig. 9c). Western block peridotites yield truly distinct forms of chondrite-normalized patterns. They are characterized by relatively steep patterns between HREE and MREE but somewhat depleted LREE resulting in upward-concave patterns. These enriched peridotites yield similar initial 143 Nd/144Nd as the OIB-type “early dikes” and hornblenditic mantle dikes that intrude the western block peridotites. This was explained by Peltonen et al. (1998) by coeval flow of alkaline melt in dikes (conduit flow) and in the residual peridotite matrix (porous flow) that was at least partly driven by filter pressing of the melt from the dikes. 7.2. Chromitites Massive chromitite bodies and peridotites rich in disseminated chromite are known only from the eastern block where they occur within a 700 m long talc-carbonate altered peridotite slice between two pillow lava slices (Fig. 2). Most of the chromitite-bearing lithology in this relatively poorly exposed zone comprises brecciated serpentinized dunites with scattered small (cm to dm size) broken fragments of massive, nodular, or orbicular chromitite. Only two larger than one meter-size chromitite bodies are currently known (Fig. 9d). The largest of the presently known pods, which is about 0.8 m wide and at least 5 m long is located in heavily carbonated serpentinite fringed by talc-carbonate rocks. Margins of this folded and boudinaged chromitite lens comprise strongly fractured and altered chromite, the fractures being filled with carbonate and chlorite, whereas the core part of the body contains surprisingly fresh coarse grained chromite (Fig. 12). This chromite has a moder-

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Fig. 10. (a) Pd/Ir vs. Ni/Cu plot of Barnes et al. (1988) for serpentinite samples from the Jormua Ophiolite. With the exception of some talc-carbonate altered samples they all plot within the mantle field consistent with their residual origin. Low Pd/Ir is inconsistent with cumulate origin for any of the samples. (b) Cr/Al vs. Ni/Al diagram illustrating the compositional variability of the Jormua mantle tectonites due to extraction of variable melt fractions. The trend from the Lizard (least depleted) to Vourinous peridotites (most depleted) illustrates that produced by increasing degree of partial melting of mantle peridotite. Reference fields after Roberts and Neary (1993). Symbols in (a) and (b): open circle (eastern block peridotites), filled circle (central block peridotites), and black triangle (western block peridotites).

7. The Mantle Section

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(a)

(b) Fig. 11. Primitive mantle (McDonough and Sun, 1995) normalized REE patterns for (a) eastern block serpentinites and (b) western block serpentinites. Note the large range in REE abundances for the eastern block samples indicating that some peridotites contain substantial amounts of infiltrated basaltic liquid (see also Fig. 9c) while some have U-shaped patterns more typical for oceanic peridotites. The pattern shapes for the western block peridotites are distinct and mimic those of alkaline (hornblenditic) dikes of the western block (Fig. 15b) suggestive of them being metasomatized by corresponding melt or fluid.

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Fig. 12. Photomicrograph of podiform chromitite from the eastern block of the Jormua Ophiolite Complex. Transparent light, width of the field ∼ 3 cm, photo by J. Väätäinen.

ately aluminous composition with average Mg# and Cr# of 76 and 55, respectively, and low TiO2 content of ∼ 0.24 wt%. Small silicate inclusions were present in many of the chromite grains but are now commonly replaced by secondary minerals. The preserved ones comprise relatively sodic tsermakitic hornblende having compositions strikingly similar with the amphibole inclusions in chromitites of some oceanic and ophiolitic peridotites (e.g., McElduff and Stumpfl, 1991). Interestingly, the chromitites yielded initial γOs values of +0.8 ± 0.5 and 3.0 ± 0.1 consistent with their derivation from a convective MORB-like oceanic mantle at the time of the ophiolite formation ∼ 1.95 Ga (Tsuru et al., 2000). 7.3. Gabbroic Feeder Dikes The eastern block mantle peridotites are intruded by a suite of gabbroic dikes. Typically these dikes are couple of meters wide and show symmetric texture indicative of them representing subvertical channels. They range from subparallel to discordant relative to the mantle foliation and branches from the main dikes cross-cut the foliation at high angle (Fig. 13). Locally, the veining of the host peridotite has resulted in detachment of peridotite xenoliths from the dike wall, leaving them ”floating” in the gabbro. Typically, the dike margins are composed of coarse grained (up to 5 cm long) often plastically deformed clinopyroxene crystals enclosed by fine grained dark green intercumulus material, whose primary composition remains obscure because of its pervasive alteration to chlorite. Locally, the clinopyroxene crystals are aligned parallel to the dike margins indicative of the

7. The Mantle Section

Fig. 13. A sketch of an outcrop of gabbroic feeder dike. Note that the dike is broadly parallel to the mantle foliation, but in the meantime brecciates the peridotite—features that are indicative of semibrittle environment for gabbro emplacement. 59

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magma flow in the dike. Towards their core parts the dikes comprise increasingly coarse grained or pegmatoidal clinopyroxene + plagioclase + ilmenite cumulates. In many places the plagioclase has been thoroughly replaced by grossular garnet and epidote, and ilmenite by sphene and rutile as a result of rodingitization and subsequent regional metamorphism. The irregular contacts of the gabbros suggest that the mantle tectonite was undergoing deformation at the time of the emplacement, and the coarse grain size of these rather thin dikes imply their slow cooling at the high ambient temperatures of the host peridotite. Importantly, too, the alteration history of the gabbroic feeder dikes and basaltic “deep dikes” are distinct. The former are often pervasively rodingitized with abundant grossular garnet while the latter are altered (epidotized) but not rodingitized sensu stricto (never garnet). This is believed to be indicative that gabbros intruded fresh (i.e., nonserpentinized) peridotites that later became serpentinized resulting in coeval rodingitization of the enclosed gabbro dikes. In contrast, the absence of rodingitization reactions in the “deep dikes” suggests that they intruded already extensively serpentinized peridotites. Rare earth element patterns of the gabbroic feeders are identical to those of the highlevel gabbros (Fig. 6b). Large ranges in the elemental abundances testify to the strong across-dike fractionation that took place in the dikes and that is broadly similar with the fractionation of the high level gabbros. Importantly, samples from both the gabbro dikes and high level stocks yield similar initial 143 Nd/144Nd values and fit along a well-defined Sm-Nd isochron. The slope of the isochron corresponds to an age (1936 ± 43 Ma) that is equal to the U-Pb zircon ages of both the gabbro dikes and high-level stocks (Peltonen et al., 1998). As a summary, we believe that the internal structure, alteration, and chemical and isotopic composition of these gabbroic feeder dikes imply that they represent feeders for the high-level gabbro bodies. However, because gabbros are frequently cross-cut by EMORB dikes, but are never observed to intrude the basaltic dikes, the gabbros are not directly related to the extrusive rocks in Jormua. More likely, their emplacement preceded (probably by a few million years) that of the sea floor volcanism. This is consistent with the evidence of the rheological properties of mantle at the time of the emplacement of the gabbro dikes (above) which suggests that they intruded the peridotites at relatively high ambient temperatures, perhaps before the mantle was exposed beneath the continental crust by lithospheric detachment faulting. 7.4. “Early” OIB-Type Dikes A distinct suite of fine grained dikes with chemical characteristics similar to oceanic island basalts (OIB) have been observed in the central block only (Figs. 3 and 6a). They occur as subvertical, NNE-trending, < 10–200 cm wide dikes that run subparallel with chromite foliations of the enclosing mantle tectonite. Their mineralogy and LILE element abundances have been strongly modified due to rodingitization. At present they mainly consist of chlorite, ilmenite, magnetite, sphene, apatite, carbonate and trace sulfides. Immobile elements, such as REE, Zr, Nb, Y, and Al can still be applied to determine their origin and led Peltonen et al. (1996) to argue that they represent primitive alkaline melts

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(a)

(b)

(c)

(d)

Fig. 14. (a) An outcrop illustrating the cross-cutting relationship of “early” OIB-type dikes and deep dikes (EMORB) within the central block mantle tectonites, (b) clinopyroxenitic cumulate dike intruding mantle tectonite, (c) thick hornblenditic mantle dike (gray) with garnetite and garnet rich veins (mottled), (d) a close-up of garnetite vein with > 50 vol% garnet crystals (white; now pseudomorphosed) growing inside from the conduit wall defining comb-layering. (c) Reprinted with the permission from Journal of Petrology, vol. 39, Oxford Univ. Press.

(ultramafic lamprophyres) from a ∼ 2.3 Ga [Nd(TDM)] plume source. Also “early dikes” are completely devoid of any subduction-related geochemical component. Since abundant EMORB dikes cross-cut the OIB-dikes at high angles, it is evident that they represent distinct episodes of magmatism (Fig. 14a). Preliminary U-Pb zircon ages determined by SIMS from one OIB-dike are consistent with field evidence that emplacement of OIB-dikes indeed preceded that of tholeiitic gabbro magmatism and ocean floor volcanism by some tens of millions of years (Peltonen et al., in preparation). Most likely their emplacement was related to the initial stages of continental rifting. These dikes are among the oldest Precambrian alkaline rocks described in the literature (Blichert-Toft et al., 1996).

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7.5. Clinopyroxenite Feeder Dikes Cumulus-textured clinopyroxenite dikes are common within the western block of the Jormua Ophiolite (Figs. 3 and 14b). Such dikes are present but rare in the central block and are completely absent from the eastern block. These dikes are typically from a few dm up to one meter wide having relatively straight and linear sharp contacts against the enclosing mantle tectonites. Locally they occur in thick dike-in-dike swarms or may form small cumulate pods within the mantle tectonite. They are subparallel with the mantle foliation but in detail are discordant. These dikes consist of clinopyroxene ortho- or mesocumulates with only minor intercumulus material and are collectively called “dry” clinopyroxenites to distinguish them from hydrous hornblendite-garnetite veins (next section). Most of the primary clinopyroxene crystals, however, have been retrogressively replaced by fibrous actinolite. The high abundances of, e.g., Mg, Cr and Ni, convex-upward primitive mantle normalized REE patterns (Fig. 15a) and low abundances of incompatible elements all reflect the accumulation of calcic pyroxene and low modal amount of intercumulus melt in these dikes. Particularly close analogies for such mantle dikes can be found from the orogenic lherzolite massifs of the French Pyrenees (e.g., Conquéré, 1971; Bodinier et al., 1987a). Two clinopyroxene cumulate dikes from central and western blocks yielded magmatic, growth-zoned zircons with relatively large spread of 207 Pb/208Pb ages between 3106 and 2718 Ma. Sm-Nd isotope data, however, clearly suggests that the crystallization age of the dikes is Proterozoic. This implies that Archean zircon grains in these dikes are xenocrysts inherited from deeper sources of the continental mantle (Peltonen et al., 2003). Therefore, central and western block mantle tectonites (which are intruded by such dikes) must represent ancient subcontinental lithospheric mantle (SCLM) that became exposed from underneath the Karelian craton during the 1.95 Ga rifting event. Since the clinopyroxenitic mantle dikes bear no evidence of having gone through melting after their formation, the mantle peridotites exposed in the central and western blocks cannot be the source for the Jormua gabbros, dikes or lavas. 7.6. Hornblendite Mantle Dikes and Garnetite Veins In addition to “dry” clinopyroxenites the western block tectonites are characterized by abundant hydrous intrusives. Frequently, the hydrous veins are subparallel with the “dry” dikes, but in detail cross-cut them. They form a somewhat heterogeneous suite of dikes and veins with significant modal and grain size variations within individual dikes. Hydrous mantle dikes include at least the following rock types (in approximate order of abundance): pure medium-grained hornblendites (Fig. 14c), garnetite veins (Fig. 14d), pegmatitic hornblendites, magnetite-ilmenite-zirconolite-rich cumulates and carbonatitic segregations. By their chemical composition the hydrous dikes are more evolved than the clinopyroxenites. Igneous mineralogy is not well-preserved but primitive mantle normalized REE patterns are particularly informative for the origin of hydrous dikes (Figs. 15b, c). These patterns

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Fig. 15. Primitive mantle (McDonough and Sun, 1995) normalized REE patterns for cumulustextured alkaline mantle dikes from the Jormua Ophiolite: (a) clinopyroxene cumulate dikes, (b) hornblendites, and (c) garnet-rich veins.

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largely reflect the mineralogical composition of the dikes. Hornblendites are characterized by steep fractionated patterns with convex-upward LREE part of the spectrum, while garnetite vein patterns clearly reflect the accumulation of garnet that strongly partitions HREE. These types of dikes, especially the garnet-bearing ones that imply high crystallization pressures, do not occur in the mantle sections of true ophiolites. They are, however, common in orogenic lherzolite massifs (SCLM), such as Lherz and Freychinéde, French Pyrenees (Conquéré, 1971; Bodinier et al., 1987b; Fabriés et al., 1991; Woodland et al., 1996) where they represent high-pressure cumulates of the continental magmatism that passed through the uppermost subcontinental lithospheric mantle. In Jormua, the hydrous dikes are intimately associated also with sporadic occurrences of carbonatite-like vein material. In fact, petrographic observations suggest that there exists a complete sequence from hornblendite veins with intercumulus carbonatite to veins consisting of more than 50% carbonates. This may be indicative that the carbonatitic material represents an extreme differentiation product of the hornblenditeproducing alkaline magma. Carbon isotope values for these carbonatites are typical for upper mantle (unpublished) and clearly distinct from secondary (metamorphic) carbonates of the talc-carbonate rocks. At present, carbonatitic segregations consist of carbonates (> 50 vol%), amphiboles, phlogopite, chlorite, magnetite, apatite, ilmenite, zircon, zirconolite and baddeleyite. Microscopic vugs consisting of carbonatitic material have been described from mantle xenoliths and mantle xenocrysts (e.g., Ionov et al., 1996; Zhang and Liou, 1994) but, to our knowledge, the carbonatitic veins within the Jormua mantle tectonites are the first mesoscopic carbonatite occurrence described from mantle samples. Zircons from one carbonatite vein have been dated by SIMS (Peltonen et al., in preparation). This sample yielded a bimodal distribution of ages. Most of grains record ages close to 1950 Ma that equals the age of the EMORB magmatism in Jormua. Some grains, however, are significantly older being ∼ 2.1 Ga. The results indicate that the igneous age of the carbonatites (and by corollary hydrous dikes in general) is close to 2.1 Ga but that most of the grains were recrystallized, as indicated by their morphology and lack of color, in the 1950 Ma thermal event. Such a time sequence is supported by field observations that imply that the only EMORB dike which is known from the western block cross-cuts not only the “dry” clinopyroxenites but also hydrous dikes. Importantly, the age of this carbonatite vein equals that of the central block OIB-dikes upper in the ophiolite stratigraphy, which indicates that hydrous cumulate dikes of the western block represent deep-seated equivalents of the OIB-type alkaline dikes (Fig. 3).

8. GEODYNAMIC SETTING OF THE JORMUA OPHIOLITE COMPLEX AND EVOLUTION OF THE KARELIAN CONTINENTAL MARGIN The internal stratigraphy of the Jormua Ophiolite Complex (e.g., absence of thick cumulate layers; thin basaltic lid), together with the presence of subcontinental lithospheric mantle imply that Jormua formed at the final stages of continental rifting representing

8. Geodynamic Setting of the Jormua Ophiolite Complex and Evolution of the Karelian Continental Margin 65

the first sea floor to be generated. On the basis of the structure of the Jormua alone, it cannot be judged whether this break-up led to development of a major ocean. Therefore, two alternative paleogeographic settings of origin are possible: (a) either the Jormua Ophiolite Complex formed between the Eastern Finland and Pudasjärvi-Iisalmi Archean complexes (Fig. 1) within a continental rift zone that never developed into major ocean, or (b) Jormua formed within the westernmost passive margin of the craton and has been tectonically transported over the Pudasjärvi-Iisalmi complexes (Fig. 1). The rift zone model (a) seriously contradicts the lithofacies of the associated (“upper Kaleva”) metasediments. Instead, we prefer the passive margin model (b) because there is no evidence of rift sedimentation around 1.95 Ga within the Kainuu Schist Belt, and because the Jormua Ophiolite (as Karelidic ophiolite fragments in general) has been obducted together with deep water slope-rise graywackes (Koistinen, 1981; Kontinen and Sorjonen-Ward, 1991) that were deposited less than 1920 ± 10 Ma ago (Huhma, 1986; Claesson et al., 1993). The monotonous deep water turbiditic lithofacies, complete absence of any type of volcanic interbeds or synsedimentary intrusions, and the presence of 1.92–1.97 old sedimentary source component from a remote unknown terrain (Claesson et al., 1993) clearly exclude the origin of “upper Kaleva” as a continental rift fill. In the passive margin model the Jormua Ophiolite can be envisaged as screens of oceanic lithosphere within a recently developed (∼ 1950 Ma) continental margin which, following the thermal subsidence, became covered by the “upper Kaleva” slope-rise turbidites (Fig. 16). This model implies that both the Jormua Ophiolite and the somewhat older (early rifting) Otanmäki alkaline gneisses are allochthonous and that their present appearance as mega-boudinaged slices within basement and cover sediment slices is due to imbrication and shearing. The deposition of the upper Kaleva metasediments as well as ophiolite obduction took place somewhere between 1920 ± 20 Ma and 1871 ± 5 Ma. The lower limit is provided by the youngest detrital zircons in these metasediments (Claesson et al., 1993), and the upper limit by the age of the oldest granite intruding the “upper Kaleva” schists (Huhma, 1986). Evidence of the flat-lying nappes that were responsible for the thrusting of the Jormua klippe onto the craton some 1.9 Ga ago has mostly been destroyed by the subsequent regional deformation. Though there is evidence of a 2 Ga mantle plume in the south-east are Fennoscandian Shield (e.g., Puchtel et al., 1998), the limited amount of 1950 Ma volcanism at the westernmost margin of the Karelian Craton suggests that it developed as a non-volcanic continental margin. The presence of both exposed continental lithospheric mantle and asthenosphere-derived igneous rocks in the Jormua Ophiolite Complex suggest a heterogeneous and asymmetric stretching and rifting process, which resulted in delamination and exposure of the subcontinental lithospheric mantle in an early stage of the ocean opening (Fig. 16; Whitmarsh et al., 2001). Phanerozoic ophiolites similar to the Jormua include the Ligurian/western Alps ophiolites that also have been related to continental break-up by asymmetric passive rifting (e.g., Lemoine et al., 1987; Rampone and Piccardo, 2001, and references therein). Modern environments that may be analogous to the provenance of the Jormua include the West Iberian non-volcanic margin where the continent-ocean transition zone consists of partially serpentinized continental mantle tectonites veined by

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Fig. 16. Proposed tectonic setting for the protolith of the Jormua Ophiolite Complex within an ocean-continent transition zone. Modified after Whitmarsh et al. (2001) emphasizing the situation at the present West Iberia passive margin.

EMORB dikes, gabbros and pyroxenites akin to those in Jormua (Chian et al., 1999; Cornen et al., 1999). Sedimentary deposits along this ∼ 1.95 Ga non-volcanic continental margin that could be related to the break-up/ocean opening are rare. One of the reasons for the obvious sediment starved nature is the peneplain nature of the Karelian continent at that time and thus lack of higher mountainous terrains along the break-up zone to supply large quantities of detritus. Furthermore, there is evidence that suggests that the whole Karelian Craton may have been submerged 1980 Ma ago shortly before the opening (Puchtel et al., 1998). The absence or very thin Jatuli-type sedimentary cover along the continental margin indicates that the break up zone probably rose above sea level and was subjected to some erosion before its rifting. Despite the obvious sediment starved nature of the opening, there still is puzzlingly little evidence of the break-up along the craton margin, and most of this evidence comes from thrusted units. One explanation could be that the break-up zone and inferred thinned margin is buried below the Svecofennian terrane in the west of the suture (Fig. 1). This scenario is however not supported by isotope data from syn- and postcollisional granites to the west of the exposed suture (Huhma, 1986; Lahtinen and Huhma, 1997; Rämö et al., 2001). Therefore, perhaps the Svecofennia-Archaean suture represents a major strike-slip fault

9. Epilogue

67

along which much of the thinned continental margin and underlying mantle was removed already before the amalgamation of the Svecofennian terrane. In this case the thrusting of the eastern Finland ophiolites and related units may have been a tectonic episode that significantly preceded and was unrelated to the actual Svecofennian collision. The present tectonic setting of the Jormua Ophiolite Complex (within klippes transported far from their inferred tectonic root west of the suture) provides few clues for unraveling the possible causes of detachment of the Jormua from the seafloor and its subsequent obduction on to the continent. One and purely speculative suggestion would be that maybe there was a shortlived event of attempted subduction beneath the continental margin preceding the obduction. This may have resulted in uplifting of the thin transitional crust that in turn facilitated the bulldozing of ultramafic seafloor and “upper Kaleva” turbidites from somewhat more distant oceanic realm onto the continent. Continental mantle in rift zones lies open to serpentinization (e.g., Perez-Gussinye and Reston, 2001) that could have made the provenance of Jormua boyonant and facilitated the obduction.

9. EPILOGUE The Jormua Ophiolite is a truly unique mafic-ultramafic rock complex that consists of two distinct components: (1) Archean subcontinental lithospheric mantle that became exposed on the seafloor due to detachment faulting, and (2) a suite of convective mantlederived alkaline and tholeiitic igneous rocks that intruded the SCLM before and during the continental rifting between 2.1 and 1.95 Ga. The mantle exposed in Jormua thus mainly represents the uppermost sub-crustal lithosphere of the Archean Karelian Craton. Intriguingly, mantle xenoliths representing this same continental mantle have been recently recovered from ∼ 500 Ma kimberlite pipes that intrude the craton margin 100 km SSE of Jormua. The mantle xenolith suite consists garnet peridotites and mantle eclogites derived from a depth range of ∼ 110–240 km (Peltonen et al., 1999; Kukkonen and Peltonen, 1999). Combined, an astonishing window to understand the evolution of the Karelian mantle for over ∼ 3 Ga period and over its entire vertical depth seems to be at hand. Unfortunately, the high degree of alteration and metamorphism of the Jormua mantle samples decreases their value; primary minerals are generally no longer available and the multistage alteration hampers study of the most subtle geochemical aspects. Nevertheless, relatively good exposure of glacially polished outcrops over an area of 50 km2 will provide many new insights to the structure and composition of Precambrian mantle.

REFERENCES Anonymous, 1972. Penrose field conference on ophiolites. Geotimes 17, 24–25. Barnes, S.-J., Boyd, R., Korneliussen, A., Nillsson, L.-P., Often, M., Pedersen, R.B., Robins, B., 1988. The use of mantle normalization and metal ratios in discriminating between the effects of partial melting, crystal fractionation and sulphide segregation on platinum-group elements, gold,

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nickel and copper: examples from Norway. In: Prichard, H.M. (Ed.), Geo-Platinum. Elsevier, pp. 113–143. Bickle, M.J., Nisbet, E.G., Martin, A., 1994. Archean greenstone belts are not oceanic crust. Journal of Geology 102, 121–138. Blichert-Toft, J., Arndt, N.T., Ludden, J.N., 1996. Precambrian alkaline magmatism. Lithos 37, 97– 111. Bodinier, J.-L., Guiraud, M., Fabriés, J., Dostal, J., Dupuy, C., 1987a. Petrogenesis of layered pyroxenites from Lherz, Freychinéde and Prades ultramafic bodies (Ariége, French Pyrenees). Geochimica et Cosmochimica Acta 51, 279–290. Bodinier, J.-L., Fabriès, J., Lorand, J.-P., Dostal, J., Dupuy, C., 1987b. Geochemistry of amphibole pyroxenite veins from the Lherz and Freychinède ultramafic bodies (Ariège, French Pyrenees). Bulletin de Mineralogie 110, 345–358. Boynton, W.V., 1984. Geochemistry of rare earth elements: meteorite studies. In: Henderson, P. (Ed.), Rare Earth Element Geochemistry. Elsevier, pp. 63–114. Cannat, M., 1993. Emplacement of mantle rocks in the seafloor at mid-ocean ridges. Journal of Geophysical Research 98, 4163–4172. Chian, D., Louden, K.E., Minshull, T.A., Whitmarsh, R.B., 1999. Deep structure of the oceancontinent transition in the southern Iberia Abyssal Plain from seismic refraction profiles: Ocean Drilling Program (Legs 149 and 173) transect. Journal of Geophysical Research 104, 7443–7462. Claesson, S., Huhma, H., Kinny, P.D., Williams, I.S., 1993. Svecofennian detrital zircon ages— implications for the Precambrian evolution of the Baltic Shield. Precambrian Research 64, 109– 130. Conquéré, F., 1971. Les pyroxénolites à amphibole et les amphibolites associées aux lherzolites du gisement de Lherz (Ariège, France): un example du rôle de l’eau au cours de la cristallisation fractionnée des liquides issus de la fusion partielle de lherzolites. Contributions to Mineralogy and Petrology 33, 32–61. Cornen, G., Girardeau, J., Monnier, C., 1999. Basalts, underplated gabbros and pyroxenites record the rifting process of the West Iberian margin. Mineralogy and Petrology 67, 111–142. Dann, J.C., 1991. Early Proterozoic ophiolite, central Arizona. Geology 19, 590–593. Eckstrand, O.R., 1975. The Dumont serpentinite; a model for control of nickeliferous opaque mineral assemblages by alteration reactions in ultramafic rocks. Economic Geology 70, 183–201. Fabriés, J., Lorand, J.-P., Bodinier, J.-L., Dupuy, C., 1991. Evolution of the upper mantle beneath the Pyrenees: evidence from the orogenic spinel lherzolite massifs. Journal of Petrology, Special Lherzolites Issue, 55–76. Fisk, M.R., Bence, A.E., 1980. Experimental crystallization of chrome spinel in FAMOUS basalt 527-1-1. Earth and Planetary Science Letters 48, 111–123. Gaál, G., Gorbatschev, R., 1987. An outline of the Precambrian development of the Baltic Shield. Precambrian Research 35, 15–52. Gole, M.J., Barnes, S.J., Hill, R.E.T., 1987. The role of fluids in the metamorphism of komatiites, Agnew nickel deposit, Western Australia. Contributions to Mineralogy and Petrology 96, 151– 162. Hanski, E.S., 1997. The Nuttio serpentinite belt, central Lapland: An example of Palaeoproterozoic ophiolitic mantle rocks in Finland. Ofioliti 22, 35–46. Huhma, H., 1986. Sm-Nd, U-Pb and Pb-Pb isotopic evidence for the origin Early Proterozoic Svecofennian crust in Finland. Geological Survey of Finland Bulletin 337, 47.

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Ionov, D.A., O’Reilly, S.Y., Genshaft, M.G., Kopylova, Y.S., 1996. Carbonate-bearing mantle peridotite xenoliths from Spitsbergen: phase relations, mineral compositions and trace element residence. Contributions to Mineralogy and Petrology 125, 375–392. Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Sciences 8, 523–548. Kärki, A., Laajoki, K., 1995. An interlinked system of folds and ductile shear zones—late stage Svecokarelian deformation in the central Fennoscandian Shield, Finland. Journal of Structural Geology 17, 1233–1247. Kohonen, J., 1995. From continental rifting to collisional crustal shortening—Paleoproterozoic Kaleva metasediments of the Höytiäinen area in North Karelia, Finland. Geological Survey of Finland Bulletin 380, 79. Koistinen, T., 1981. Structural evolution of an early Proterozoic sratabound Cu-Co-Zn deposit, Outokumpu, Finland. Transactions of the Royal Society of Edinburgh, Earth Sciences 72, 115–158. Koistinen, T.J., Stephens, M.B., Bogatchev, V., Nogulen, Ø., Wennerström, M., Korhonen, J., Geological map of the Fennoscandian Shield, scale 1:2 000 000. Geological Surveys of Finland, Norway and Sweden, and the North-West Department of Natural Resources of Russia. Kontinen, 1987. An Early Proterozoic ophiolite—the Jormua mafic-ultramafic complex, northeastern Finland. Precambrian Research 35, 313–341. Kontinen, A., Sorjonen-Ward, P., 1991. Geochemistry of metagraywackes and metapelites from the Palaeoproterozoic Nuasjärvi group, Kainuu schist belt and the Savo Province, north Karelia: implications for provenance, lithostratigraphic correlation and depositional setting. Geological Survey of Finland Special Paper 12, 21–22. Kontinen, A., 1998. Geological map of the Jormua area 1:50 000. Geological Survey of Finland Special Paper 26, Appendix. Kontinen, A., Peltonen, P., 1998. Excursion to the Jormua ophiolite complex. In: Hanski, E., Vuollo, J. (Eds.), International Ophiolite Symposium and Field Excursion: Generation and Emplacement of Ophiolites through Time, August 10–15, 1998, Oulu, Finland. Geological Survey of Finland Special Paper 26, 69–89. Korsman, K., Korja, T., Pajunen, M., Virransalo, P., 1999. The GGT/SVEKA transect: structure and evolution of the continental crust in the Paleoproterozoic Svecofennian orogen in Finland. International Geology Review 41, 287–333. Kukkonen, I.T., Peltonen, P., 1999. Xenolith-controlled geotherm for the central Fennoscandian Shield—implications for the lithosphere-asthenosphere relation. Tectonophysics 304, 301–315. Kusky, T.M., Li, J.H., Tucker, R., 2001. The Archean Dongwanzi ophiolite complex, North China Craton: 2.505-billion-year-old oceanic crust and mantle. Science 292, 1142–1145. Lahtinen, R., Huhma, H., 1997. Isotopic and geochemical constraints on the evolution of the 1.93– 1.79 Ga Svecofennian crust and mantle in Finland. Precambrian Research 82, 13–34. Lemoine, M., Tricart, P., Boillot, G., 1987. Ultramafic and gabbroic ocean floor of the Ligurian Tethys (Alps, Corsica, Apennines): In search of a genetic model. Geology 15, 622–625. Louden, K., Lau, H., 2001. Insights from scientific drilling on rifted continental margins. Geoscience Canada 28, 187–195. Liipo, J., Vuollo, J., Nykänen, V., Piirainen, T., Pekkarinen, L., Tuokko, I., 1995. Chromites from the early Proterozoic Outokumpu-Jormua Ophiolite Belt: a comparison with chromites from Mesozoic ophiolites. Lithos 36, 15–27. McDonough, W.F., Frey, F.A., 1989. Rare earth elements in upper mantle rocks. In: Lipin, B.R., McKay, G.A. (Eds.), Geochemistry and Mineralogy of Rare Earth Elements. In: Reviews in Mineralogy, vol. 21, pp. 99–145.

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McDonough, W.F., Sun, S.-s., 1995. The composition of the Earth. Chemical Geology 120, 223–253. McElduff, B., Stumpfl, E.F., 1991. The chromitite deposits of the Troodos Complex, Cyprus— evidence for the role of a fluid phase accompanying chromite formation. Mineralium Deposita 26, 307–318. O’Hanley, D.S., 1996. Serpentinites. Records of Tectonic and Petrological History. Oxford Univ. Press, p. 277. Pearce, J.A., 1982. Trace element characteristics of lavas from destructive plate boundaries. In: Thorpe, R.S. (Ed.), Andesites; Orogenic Andesites and Related Rocks. John Wiley & Sons, pp. 525–548. Pearce, J.A., 1983. Role of subcontinental lithosphere in magma genesis at active continental margins. In: Hawkesworth, C.J., Norry, M.J. (Eds.), Continental Basalts and Mantle Xenoliths. Shiva Publishing, UK, pp. 230–249. Pearce, J.A., Harris, N.B.W., Tindle, A.G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology 25, 956–983. Peltonen, P., Kontinen, A., Huhma, H., 1996. Petrology and geochemistry of metabasalts from the 1.95 Ga Jormua Ophiolite, northeastern Finland. Journal of Petrology 37, 1359–1383. Peltonen, P., Kontinen, A., Huhma, H., 1998. Petrogenesis of the mantle sequence of the Jormua Ophiolite (Finland): Melt migration in the upper mantle during Palaeoproterozoic continental break-up. Journal of Petrology 39, 297–329. Peltonen, P., Huhma, H., Tyni, M., Shimizu, N., 1999. Garnet-peridotite xenoliths from kimberlites of Finland: nature of the continental mantle at Archaean craton-Proterozoic mobile belt transition. In: Gurney, J.J., Gurney, J.L., Pascoe, M.D., Richardson, S.H. (Eds.), Proceedings of the 7th International Kimberlite Conference, vol. 2, pp. 664–676. Peltonen, P., Mänttäri, I., Huhma, H., Kontinen, A., 2003. Archean zircons from the mantle: the Jormua Ophiolite revisited. Geology 31, 645–648. Perez-Gussinye, M., Reston, T.J., 2001. Rheological evolution during extension at nonvolcanic rifted margins: Onset of serpentinisation and development of detachments leading to continental breakup. Journal of Geophysical Research 106, 3961–3976. Puchtel, L.S., Arndt, N.T., Hofmann, A.W., Haase, K.M., Kröner, A., Kulikov, V.S., Kulikova, V.V., Garbe-Schönberg, C.-D., Nemchin, A.A., 1998. Petrology of mafic lavas within the Onega plateau, central Karelia: evidence for 2.0 Ga plume-related continental crustal growth in the Baltic Shield. Contributions to Mineralogy and Petrology 130, 134–153. Rampone, E., Piccardo, G.B., 2001. The ophiolite-oceanic lithosphere analogue: new insights from the Northern Apennines (Italy). In: Dilek, Y., Moores, E.M., Elthon, D., Nicholas, A. (Eds.), Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America Special Paper 349, 21–34. Rämö, O.T., Vaasjoki, M., Mänttäri, I., Elliot, B.A., Nironen, M., 2001. Petrogenesis of the postkinematic magmatism of the Central Finland Granitoid Complex. Journal of Petrology 42, 1971– 1993. Roberts, S., Neary, C., 1993. Petrogenesis of ophiolitic chromitite. In: Prichard, H.M., Alabaster, T., Harris, N.B.W., Neary, C.R. (Eds.), Magmatic Processes and Plate Tectonics. Geological Society Special Publication 76, 257–272. Scott, D.J., Helmstaedt, H., Bickle, M.J., 1992. Purtuniq ophiolite, Cape Smith Belt, northern Quebec, Canada: a reconstructed section of Early Proterozoic oceanic crust. Geology 20, 173–176. Tsuru, A., Walker, R.J., Kontinen, A., Peltonen, P., Hanski, E., 2000. Re-Os isotopic systematics of the 1.95 Ga Jormua Ophiolite Complex, northeastern Finland. Chemical Geology 164, 123–141.

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Tuisku, 1997. An introduction to Palaeoproterozoic (Svecofennian) regional metamorphism in Kainuu and Lapland, Finland. In: Evins, P., Laajoki, K. (Eds.), Archaean and Early Proterozoic (Karelian) Evolution of the Kainuu-Peräpohja Area, Northern Finland. Res Terrae, Ser. A 13, 27–31. Vuollo, J., Piirainen, T., 1989. Mineralogical evidence for an ophiolite from the Outokumpu serpentinites in North Karelia, Finland. Bulletin of the Geological Society of Finland 61, 95–112. Whitmarsh, R.B., Manatschai, G., Minshull, T.A., 2001. Evolution of magma-poor continental margins from rifting to seafloor spreading. Nature 413, 150–154. Will, T.M., Powell, R., Holland, T.J.B., 1990. A calculated petrogenetic grid for ultramafic rocks in the system CaO-FeO-MgO-Al2 O3 -SiO2 -CO2 -H2 O at low pressures. Contributions to Mineralogy and Petrology 105, 347–358. Woodland, A.B., Kornprobst, J., McPherson, E., Bodinier, J.-L., Menzies, M.A., 1996. Metasomatic interactions in the lithospheric mantle: petrologic evidence from the Lherz massif, French Pyrenees. Chemical Geology 134, 83–112. Zhang, R.Y., Liou, J.G., 1994. Significance of magnesite paragenesis in ultrahigh-pressure metamorphic rocks. American Mineralogy 79, 397–400.

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THE 1.73 GA PAYSON OPHIOLITE, ARIZONA, USA J.C. DANN 90 Old Stow Road, Concord, MA 01742, USA

The 1.73 Ga Payson Ophiolite is a shallow-dipping, layered sequence of coeval gabbro, sheeted dikes, and submarine volcanic rocks, partly disjointed by later intrusion and deformation. The sheeted dike complex is spectacularly exposed as cliffs and water-polished outcrop in many shallow canyons. Gabbro-dike mingling and mutual intrusion attest to rooting of the sheeted dike complex in the underlying gabbro. A stratigraphically continuous zone of intense alteration marks the transition from sheeted dikes to submarine volcanics. A tonalite/dacite suite occurs as rare lavas and as dikes and hypabyssal plutons, mutually intrusive with the basaltic sheeted dikes and gabbro. An older basement complex occurs as roof pendants in gabbro and screens in the sheeted dike complex. An actualistic tectonic model of an intra-arc basin formed by seafloor spreading along an arc-parallel strike-slip fault system explains the origin of the Payson Ophiolite, its emplacement within the arc, and accretion to North America during the ca. 1.70 Ga Yavapai Orogeny.

1. INTRODUCTION The 1.73 Ga Payson Ophiolite (Dann, 1991, 1992, 1997a, 1997b) is about 90 km northeast of Phoenix within the Yavapai-Mazatzal orogenic belt of central Arizona (Fig. 1A). It is one of only a few Early Proterozoic ophiolites known worldwide with a well developed sheeted dike complex and the horizontally layered structure characteristic of Phanerozoic ophiolites and modern oceanic crust (see Moores, 2002, for review). The sheeted dike complex and transitions to overlying volcanic and underlying gabbroic rocks are well exposed as water-polished outcrops in easily accessed canyons (stop one in field trip guide, Karlstrom et al., 1990). Most of the ophiolite remains gently dipping. Fold-and-thrust deformation is localized, and no penetrative fabric occurs in the ophiolite. Even the overlying turbidites have only a locally developed, spaced fracture cleavage. Although low greenschist-grade metamorphism affected the ophiolite, large areas of the gabbro are 95% igneous minerals. The quality of exposure and preservation of primary features facilitated detailed structural mapping and petrological, geochemical, and geochronological analyses of the Payson Ophiolite (PO). The purpose of this paper is summarize the work done and discuss its contribution to our understanding of the mechanics of seafloor spreading, plate tectonics, and continental assembly during the Early Proterozoic. DOI: 10.1016/S0166-2635(04)13002-8

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Fig. 1. (A) Geologic map showing the location of the Payson Ophiolite within the Mazatzal crustal block in the Early Proterozoic of central Arizona (modified from Karlstrom et al., 1990, and Anderson, 1989; ‘T’ is Tonto Creek). (B) Map showing the location of central Arizona within the 1.6–1.8 Ga orogenic belt of North America (from Hoffman, 1989). (C) Tectonostratigraphic columns comparing the Ash Creek and Mazatzal crustal blocks divided by the Moore Gulch shear zone.

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First, we need to feel confident that the complicated map pattern (Fig. 2) actually conceals the layered structure that distinguishes an ophiolite from other crustal sections. Ophiolites are defined by their pseudostratigraphy or horizontally layered structure of mantle tectonite, gabbro, sheeted dikes, and submarine volcanic and sedimentary rocks (Moores, 1982). Importantly, the transitional zones between lithologic layers indicate that the layers were forming simultaneously (unlike real stratigraphy). In addition, the distribution of tonalites, hydrothermal alteration, and chemical sediments record important processes ongoing during development of the crust. In most greenstone belts, deformation and intrusion of late granitoids has dislocated submarine volcanic sequences from their hypabyssal equivalents. What makes the PO special is that this connection is intact. In addition, we need to distinguish between (1) intrusion of a dike swarm into an older plutonic complex and (2) rooting of sheeted dikes in coeval gabbro. These questions were addressed by detailed structural mapping and analysis of the field relations (Dann, 1992). Second, the story of the origin and emplacement of the ophiolite is recorded in its relationship to the rest of the orogenic belt (Dann and Bowring, 1997). Ophiolites commonly occur within terranes or shear zone-bound crustal blocks that record tectonostratigraphic histories and original tectonic settings that are incompatible with neighboring terranes (Fig. 1C). From the comparative tectonostratigraphic histories, the geochemistry of the ophiolitic magma, the geometry of extensional and convergent tectonism, and with reference to modern examples, an actualistic tectonic model can be developed that provides a useful predictive framework for understanding the creation and assembly of Proterozoic crust in the southwestern United States.

2. REGIONAL SETTING The Payson Ophiolite is in the Mazatzal crustal block bound by the Moore Gulch shear zone on the northwest and a belt of post-assembly granites to the southeast (Fig. 1A). This 50–60 km wide crustal block occurs at the eastern end of a 600 km transect of the Proterozoic orogenic belt, exposed in the Transition Zone between the Colorado Plateau and the Basin and Range Province. In central Arizona this orogenic belt consists of submarine volcanic and volcaniclastic rocks that host massive sulfide deposits and are intruded by granitoid plutons. These lithologic associations are interpreted by most workers to represent magmatic arcs (e.g., Anderson, 1989). The magmatic arcs formed over a 40 m.y. period from ca. 1.75 to 1.71 Ga (Bowring et al., 1991), locally involving older crust. From the high estimated rate of crustal growth, the predominance of juvenile magmatic arc rocks, and the juxtaposition of distinct terranes, Karlstrom and Bowring (1988) proposed that the orogenic belt formed by accretion of arc terranes along a convergent plate boundary. Most of the deformation along the block boundaries reflects post-assembly crustal shortening and differential uplift (Bowring and Karlstrom, 1990). Consequently, the role of the block boundaries during the assembly of terranes remains speculative, but it may be particularly important in the origin and emplacement of the Payson Ophiolite.

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Fig. 2. Geologic Map and cross sections of the Payson Ophiolite (contacts outside the ophiolite from Wrucke and Conway, 1987). Sheeted dikes are best exposed in American Gulch (AG), Rattlesnake Canyon, (RC), St. John’s Creek (SJ), and along the east flank of the Mazatzal Mountains (EF). The Larson Spring Formation of the basement complex, intruded and overlain by the ophiolite, are best exposed at the East Verde River (a), Crackerjack Mine (b), Larson Spring (c), Center Creek (d), and Gisela (e).

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3. PAYSON OPHIOLITE The distribution and orientation of dikes reflects the structure of the ophiolite. Sheeted dikes underlie submarine volcanic rocks on the west side of the map (EF, Fig. 2) and overlie gabbro east of the Tertiary valley (AG, RC, Fig. 2). The sheeted dikes dip about 75 degrees to the northeast where gabbro devoid of dikes occurs (‘rv’, Fig. 2). The northeastto-southwest sequence of gabbro, sheeted dikes, and volcanic rocks, combined with the 75 degree northeast dip of the dikes, indicates that the ophiolite pseudostratigraphy dips about 15 degrees to the southwest (cross section, Fig. 2). This reconstruction, based on a perpendicular relationship between dikes and the pseudostratigraphy, is justified by the angular relationships between bedded basaltic volcanic rocks and (1) underlying sheeted dikes (EF, Fig. 2) and (2) bedded felsic volcaniclastic rocks in the basement complex (Dann, 1997a). Parallel to the ophiolite pseudo-stratigraphy, the ca. 1.70 Ga Payson granite intruded its gabbroic roof as a sheet dipping 15–25 degrees to the southwest (Conway et al., 1987). Due to the shallow dip of both the Payson Granite and the ophiolite, the gabbro-norite of Round Valley (‘rv’, Fig. 2) is the deepest level of the ophiolite exposed. The mantle section of the PO remains hidden beneath the Payson Granite. Tertiary normal faults created the sediment-filled valley (Fig. 2) and the complicated map pattern by juxtaposing different levels of the ophiolite pseudostratigraphy. Description of the ophiolite proceeds from gabbro to volcanic rocks with particular emphasis on the transitions that establish the contemporaneity of the dikes with both the gabbro and volcanic section. 3.1. Gabbro Most of the exposed ophiolite is gabbro that forms a pseudostratigraphic layer of plutons beneath the sheeted dike complex. The Round Valley gabbro (‘rv’, Fig. 2) in the northeastern part of the map area is coarse-grained, completely devoid of mafic dikes, and the least altered (1–5%) of all mafic rocks in the ophiolite. The NE-SW cross section shows that the RV solidified about 1.5–2 km below the sheeted dike complex (inset, Fig. 2). Hornblende gabbro mantles the RV, and toward the northwest-trending transition to the sheeted dike complex, it becomes increasingly fractured, altered, and intruded by mafic dikes. Just below the sheeted dike complex, distinct plutons are resolved, based on texture, flow fabrics, and mineralogy, especially the isotropic quartz diorite (‘d’, Fig. 2) that yielded a U/Pb zircon age of 1.73 Ga (Bowring et al., 1991). Gabbro also intruded the sheeted dike complex and the volcanic section as sill-like bodies (e.g., ‘sm’, Fig. 2). All the gabbroic rocks contain hornblende, locally with coarse-grained, ophitic and poikolitic textures (< 4 cm). Elongate plagioclase crystals and folded modal layering define a lineation and/or foliation in the gabbros below the sheeted dike complex. The lineation is primarily orthogonal to the dikes, indicating flow of crystal-rich magma parallel to the direction of extension (Dann, 1997a, Figs. 5, 6). However, the orientation of the layering varies and is locally parallel to the dikes. Structural analysis, based on the orientation of bedding in the basement complex, reveals that the plutonic core of the ophiolite was not tectonically rotated and that magmatic layering with locally steep dips is a primary feature (Dann, 1997a, Fig. 13).

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Rothery (1983) and Greene (1989) reached similar conclusions for the Semail and Lokken ophiolites, respectively. In general, the magmatic layering is less reliable for structural reconstructions than the orientation of sheeted dikes and distribution of the pseudostratigraphy. 3.2. Sheeted Dike Complex The sheeted dike complex is spectacularly exposed in four major canyons as continuous water-polished outcrop, cliffs of parallel slabs of dikes, and hillsides of dikes standing in relief (Fig. 3A). As a result of the overlapping and nearly parallel intrusion, most dikes have well defined chilled margins against other dikes (Figs. 3B, C), and dike splitting and low-angle cross cutting displaced segments of early dikes over tens of meters. Over 600 m of measured section permit an estimate of the proportions of mafic and felsic dikes and screens of gabbro and granitic basement and other features that distinguish one area of sheeted dikes from another (see Dann, 1997a, Fig. 8, and Dann, 1997b, Fig. 3, for examples of sections). The most pronounced difference between the top and bottom of the sheeted dike complex is the 3-fold increase in the average width of dikes (e.g., compare Fig. 4A and Fig. 5E). The depth-thickness relationship inspired a new model for the vertical development of sheeted dike complexes (Dann, 1997b). No complete section from top to bottom of the sheeted dike complex is exposed. The base of the sheeted dike complex and transition to the underlying gabbro is best exposed in the Rattlesnake Canyon area (RC, Fig. 2). The top of the sheeted dike complex and transition to submarine volcanic rocks is only exposed on the west side of the Tertiary valley (EF, Fig. 2). American Gulch (AG, Fig. 2) provides the best display of the intrusive sequence of mafic and dacitic dikes into the basement screens (Figs. 3B–D). Analysis of dike widths places AG at an intermediate level within the sheeted dike complex. These three areas of sheeted dikes form the corners of a triangle, giving the appearance that the sheeted dike complex is not a continuous layer. However, the granitic pluton in the middle of the triangle (SJ, Fig. 2) contains a large roof pendant of 100% sheeted dikes that testify to the original continuity of the layer of sheeted dikes that was eroded off the top of the pluton. 3.3. Transition from Gabbro to Sheeted Dikes Best exposed in the Rattlesnake Canyon area (RC, Fig. 2), the transition from gabbro to sheeted dikes trends northwest, parallel to the trend of the dikes. Over about 100 m, dikes increase from < 50% in gabbro to > 90% with gabbroic screens. The dikes are generally thick and coarse-grained (Figs. 4A, B) with tonalitic dikes up to 15 m thick. Early thick dikes have weakly chilled contacts that locally mingle with gabbroic screens (Fig. 4B). The transition zone contains small dioritic intrusions and locally developed intrusion breccias. Some screens contain a mixture of coarse-grained gabbroic material and finer-grained dike material that is interlayered parallel to a flow foliation/lineation (Fig. 4C). Xenoliths of porphyritic basalt aligned in the flow foliation are identical to dikes cutting the gabbro

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Fig. 3. Sheeted dikes of American Gulch (AG, Fig. 2). (A) View looking northwest down into the gulch. Paler dike (V) is a 3 m thick dacitic dike that is also on outcrop map (D). (B) Water-polished outcrop of sheeted dikes. Basaltic dikes (b) intrude a dacitic dike (x) that intrudes a granitic basement screen (g). (C) One-way chilling (arrows) of basaltic dikes against porphyritic dike (scale bar is 1 cm). (D) Outcrop map of sheeted dikes showing 3 dacitic dikes and 18 basaltic dikes.

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(Fig. 4D). In the gabbro below the sheeted dikes, mafic dikes are pulled apart parallel to the flow foliation, forming trains of enclaves with deformed shapes and mingled contacts. In addition, the gabbro contains conjugate shear bands locally cut by dikes, indicating that the gabbro underwent hot sub-solidus extension consistent with the dikes (see Dann, 1997a, Fig. 6). Gabbro-dike mingling, mutual intrusion, and stretching of xenoliths parallel to the flow foliation in the gabbro indicate that the dikes are rooted in coeval gabbro (see Dann, 1997b, Fig. 6, for summary figure). Flowing gabbroic crystal mush cannibalized the base of the dikes, forming trains of deformed enclaves and mixed dike-gabbro layers along the direction of flow. The transition represents a fluctuating rheological boundary between brittle fracture with dike intrusion and fluid flow of crystal-rich gabbroic magma. This process caused a stepwise decrease in the abundance of dikes with depth, such that the sheeted dike complex bottoms out and dikes are completely absent 1–1.5 km below the sheeted dike complex. Similar features that attest to the dynamic process of dikes rooting in the underlying gabbro occur in many ophiolites (Pederson, 1986; Furnes et al., 1988; Nicholas and Boudier, 1991; Skjerlie and Furnes, 1996). 3.4. Volcanic Section The volcanic section is a thin 400–500 m thick sequence that lies between overlying turbidites (Fig. 5A) and underlying sheeted dikes (Fig. 5E) along the west side of the map area (EF, section Y–Y , Fig. 2). The section consists mostly of basaltic sheet flows with rare outcrops of pillowed flows. The pillows have a 2–3 cm thick band of vesicles inside well-defined selvages. Flow tops are readily recognizable by large amygdules, breccia, jasper infillings, and interflow sediments (Fig. 5B). Besides basaltic flows, the volcanic section contains an auto-brecciated dacitic flow and associated debris-flow deposits. Clastic interflow sediments include volcanic conglomerates, greywacke, and tuffs. Graded beds and scour-and-fill structures indicate consistent northwest younging (Fig. 5B). Chemical interflow sediments include thick lenses of magnetite-rich banded iron formation, jasper, and chert. Sediments are locally deformed by the weight of overlying flows (Fig. 5B). The base of the volcanic section is intruded by thin basaltic sills and dikes that are distinguished by their cross-cutting relations or vesicle layering parallel to chilled margins. Although the paucity of pillowed flows is unusual for a submarine volcanic section, the thickness of the volcanic section (400–500 m) is typical of many ophiolites (usually < 1 km;

Fig. 4. Base of sheeted dike complex. (A) Two thick mafic dikes (arrows, 4 m wide; map folder at top for scale) in the sheeted dike complex of Rattlesnake Canyon (RC, Fig. 2). (B) Mingled and poorly chilled contact (arrow) between mafic dike and gabbroic screen (gb; pen for scale). (C) Gabbroic screen (gb) in sheeted dikes (arrows point to chills), showing early dike material (dark and fine-grained, between arrows) mingled with, and drawn out along the flow foliation in, the gabbro (1 m across field of view). (D) Porphyritic dike xenoliths elongate parallel to the flow foliation of the gabbro are identical to porphyritic dikes cutting the gabbro (pen for scale).

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Fig. 5. (A) Graded beds of the turbidite sequence overlying the ophiolite (arrow points in younging direction). (B) Interflow clastic sediment, graded from course sand above the amygdaloidal flow top (arrow base) to pale tuff that penetrates crack in overlying flow (arrow tip). (C) Silicified mafic rock of the altered transition from sheeted dikes to volcanic flows (coin for scale). Silica veining (‘s’) outlines pseudo-breccia texture (‘p’). (D) Schematic column showing increasing alteration (white) at the top of the sheeted dike complex. (E) Relatively unaltered sheeted dikes, several 100 m below the volcanic section, with an average width of about 80 cm.

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Moores, 1982) and sections of oceanic crust exposed on the ocean floor (< 500 m; Francheteau et al., 1992). 3.5. Transition from Sheeted Dikes to Submarine Basalt The transition from sheeted dikes to volcanic rocks is marked by a cliff-forming, stratigraphically continuous, siliceous zone. This transitional zone and the overlying volcanic rocks are steeply dipping to overturned along the east flank of the Mazatzal Mountains (EF, section Y–Y , Fig. 2), where canyons provide cross-sectional exposures. The volcanic section has a sharp lower contact with the massive siliceous zone. In contrast, the sheeted dike complex has a gradational upper contact with the siliceous zone (Fig. 5D). With increasing degrees of alteration, the sheeted dikes merge upwards with the siliceous zone where all primary features are obliterated and the rock takes on a pseudo-breccia texture (Fig. 5C). A few dikes cut the base of the siliceous zone and some bedded, tuffaceous, cherty sediments occur near the top, indicating that this transition zone consists of a mixture of dikes and flows. Although the siliceous transition is internally disjointed by conjugate fractures and shear bands and is the only ophiolitic unit with a crude, nonpenetrative cleavage, the magmatic transition from sheeted dikes to volcanic rocks is intact. The magmatic connection between the dikes and overlying volcanic flows is also indicated by their co-varying geochemistry and phenocryst types (e.g., plag-phyric flows overlie areas of plag-phyric dikes, etc). Although some sills and dikes occur within the volcanic section, the transition from 100% dikes to the volcanic section occurs within about 100 m. The abrupt transition from sheeted dikes to volcanic rocks is characteristic of many well-studied ophiolites (Pallister, 1981; Moores, 1982; Rosencrantz, 1983; Harper, 1984; Baragar et al., 1987) and distinguishes ophiolites from other types of volcanic centers. 3.6. Tonalites and Dacites A suite of tonalitic and dacitic rocks occurs in all three exposed layers of the ophiolitegabbro, sheeted dikes, and volcanic rocks. Rare dacitic flows, breccias, and tuffs are interbedded with the basaltic flows. Tonalitic and dacitic dikes make up about 10% of the sheeted dike complex and locally up to 38%. They are mutually intrusive with, and parallel to, the mafic dikes (Figs. 3B, D). The PO is unusual for the high proportion of dacitic dikes in the sheeted dike complex. These dikes are white to pink in color, and the contrast with the dark green and gray basaltic dikes shows off the cross cutting contacts and makes it easier to reconstruct the intrusive sequence. Besides the mutually crosscutting relations in the sheeted dike complex, composite dikes in the gabbro have co-mingled dacitic and basaltic phases that testify to the coeval intrusion of these two distinct magma types (see Dann, 1997a, Fig. 13). In addition, the tonalite/dacite suite includes small sub-spherical mafic enclaves that indicate mingling and incomplete mixing of basalt in the tonalitic magma. Locally, granitic screens melted and intruded the mafic dikes, suggesting that the abundance of dacitic magma may have been generated from felsic basement within the ophiolite. On

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the other hand, the geochemistry of the upper gabbro and tonalite are complimentary, suggesting that the tonalite represents a residual liquid, filter pressed from the gabbro (Dann, 1992), similar to tonalites in the Karmoy Ophiolite (Pederson and Malpas, 1984). A more detailed petrological and geochemical study is needed to determine the petrogenesis of the variety of tonalites within the PO.

4. BASEMENT COMPLEX The basement complex consists of (1) coarse-grained granitoids, and (2) hypabyssal granite overlain by (3) submarine felsic volcaniclastic rocks of the Larson Spring Formation (Fig. 2). The basement complex is intruded by gabbro and cut by mafic dikes. Few ophiolites contain such a well-defined suite of older felsic rocks, but most of the ophiolite is devoid of roof pendants or screens. The basement complex indicates that the ophiolite developed from extension of older arc crust. It provides a valuable reference frame within the ophiolite, recording a tectonic event that preceded development of the ophiolite as well as the degree of magmatic extension. A northeast-trending belt of coarse-grained, foliated, tonalite and quartz monzodiorite (Fig. 2) occurs as roof pendants in gabbro and is intruded by a swarm of basaltic dikes and small plutons of isotropic porphyritic diorite. The presence of alkali feldspar, low An content of the plagioclase, darker hornblende, and more quartz and biotite distinguishes this suite of rocks from the gabbro and quartz diorite of the ophiolite. Abundant zircons yield a U/Pb age of 1.75 Ga (Dann et al., 1989), making it 20 m.y. older than the ophiolite and one of the oldest rocks in central Arizona. What this pluton intruded is not exposed. Screens in the sheeted dike complex and roof pendants in gabbro define a northeasttrending belt of felsic volcanic rocks, the Larson Spring Formation, and underlying, finegrained, isotropic granitoids (‘a–d’, Fig. 2). Near Larson Spring (‘c’, Fig. 2), the most intact felsic section underlain by hypabyssal granite sits as a block in gabbro, intruded by basaltic dikes (< 20%). Felsic volcaniclastic breccia, beds graded from breccia to porcelainite, and plagio-arenites with scour-and-fill structures indicate that the bedding dips consistently to the northwest. Massive, aphyric, felsic flows, mafic sediments, and chert also occur locally. How much older than the ophiolite these rocks are is not known.

5. HYDROTHERMAL ALTERATION Extensional tectonism generates hydrothermal circulation by facilitating the shallow emplacement of hot magma (dikes) and increasing the fracture porosity. Seafloor hydrothermal alteration and later burial metamorphism produce similar greenschist assemblages in mafic rocks. The effects of seafloor alteration can be distinguished by locating alteration or hydrothermal products that are unambiguously associated with intrusion or eruption of magma. In addition, the location of the most severe alteration is diagnostic of a narrow axial zone of intrusion, the hallmark of seafloor spreading.

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The sheeted dike complex of the PO displays patterns of alteration that are unambiguously associated with intrusion of the dikes. First, the interiors of some dikes with identifiable chilled margins are completely replaced by epidote and quartz. These green epidosites are locally intruded by grey greenschist dikes that retain their igneous textures and whole rock compositions. Second, dikes with a reddish-brown, hematitic alteration near the top of the sheeted dike complex are split by dikes with the more usual greenschist alteration. Third, veins of quartz (with or without sulfides), especially common along the margins of dikes, are also cut by late dikes. Based on the cross cutting relations, the epidosite, hematitic alteration, and quartz veins are temporally associated with intrusion of the dikes. The epidosites are typically found in sheeted dike complexes of ophiolites and are interpreted to require strongly localized, high temperature, fluid fluxes (Gillis and Banerjee, 2000). The most intensely altered unit in the Payson Ophiolite is the siliceous zone that marks the sheeted dike-volcanic transition. Except a few patches of sediment near the top and late dikes near the base of this zone, all original outcrop-scale features are obliterated by quartz-chlorite alteration that locally renders a pseudo-breccia texture (Fig. 5C). Disseminated sulfides are common, and weathering of small patches of gossan creates orange iron straining on some cliff exposures. Silicification and mineralization are concentrated at the transition from sheeted dikes to volcanic rocks in both modern oceanic crust (Alt et al., 1986) and in other ophiolites (e.g., Josephine Ophiolite; Harper et al., 1988). The transition from sheeted dikes to volcanic rocks is the site of eruption on the seafloor. Therefore, the alteration represented by the siliceous zone occurred around vents at the axial zone of seafloor spreading, where fractures open as fissures and permeability is highest. What distinguishes alteration of the PO from other ophiolites is the high degree of silicification to form an erosion-resistant, stratigraphically continuous zone at the dike-volcanic transition. Chemical sediments are interlayered with the volcanic flows and represent the exhalative products of a hydrothermal system that was active during seafloor spreading. A Cu-Pb massive sulfide deposit occurs just above the transition from sheeted dikes to submarine basalts in the southernmost exposure of the PO (Wessels and Karlstrom, 1991) in Tonto Creek (‘T’ on Fig. 1A). Massive sulfide deposits are common components of ophiolites (Gillis and Banerjee, 2000).

6. GEOCHEMISTRY All components of the 1.73 Ga Payson Ophiolite—submarine basalts, sheeted dikes, gabbro, and tonalite—as well as the 1.75 Ga basement complex have geochemical signatures of magmatic arc rocks (Dann, 1991, 1992). These include light rare earth element (LREE) and large-ion lithophile element (LIL) enrichment and relative high-field strength element (HFSE) depletion, typical of arc rocks (Pearce et al., 1984). The basaltic rocks plot in the ‘arc’ or ‘suprasubduction zone’ field in all tectonic discrimination diagrams (Th-HfTa, Ti-Cr, Cr-Y, etc.). Analyses of mafic dikes define a tholeiitic fractionation trend of

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increasing FeO, TiO2 , P2 O3 , and V with decreasing MgO, typical of arc tholeiites. Hydrous igneous minerals (hornblende, biotite) in the gabbros and clinopyroxene-controlled fractionation of the dikes indicate that the parental magma was hydrous, consistent with generation above a subduction zone. Nd isotopic analyses indicate the influence of an older LREE-enriched component (Dann et al., 1993). The ophiolite has the geochemical and isotopic features of the high Ce/Yb suite of arc magmas (cf. Hawkesworth et al., 1993). Sheeted dikes may form during rifting of arcs, rifting of continental crust, or rifting of volcanic islands (e.g., Hawaii (Walker, 1987) or Canary Islands (Stillman, 1987)). However, the geochemistry alone indicates that the PO formed during an extensional phase in the evolution of a magmatic arc.

7. DEFORMATION OF THE OPHIOLITE Four stages of deformation are recorded by structures of the PO. First, rotation of the basement complex occurred prior to development of the ophiolite. Then, crustal extension guided the intrusion of basaltic magma, culminating in seafloor spreading. Structures that formed during intrusion of the basaltic magma provide important clues about the tectonic setting. Third, the ophiolite and overlying rocks are affected by two Early Proterozoic episodes of coaxial convergent deformation, the ca. 1.70 Ga Yavapai orogeny (D2 ) and the ca. 1.67 Ga Mazatzal Orogeny (D3 ). An earlier period of deformation (D1 ) only affected terranes west of the Moore Gulch Fault (Figs. 1A, C). Finally, Tertiary extensional faulting created the sediment-filled valley that cuts through the middle of the map area (Fig. 2). The bedded rocks of the Larson Spring Formation occur as narrow screens in the sheeted dike complex in Center Creek (‘d’, Fig. 2). The orientation of bedding defines an angular unconformity beneath the volcanic section of the ophiolite, which exposed granite on the pre-ophiolite paleosurface (see Dann, 1997a, Fig. 12). This angular relationship, the consistent orientation of bedding in the roof pendants, and repetition of the same lithologies from one roof pendant to another suggests that blocks of the basement complex rotated along listric normal faults prior to development of the ophiolite. Fault-bound domains of sheeted dikes with dips diverging 30 degrees from adjacent domains indicate block rotations along normal faults in the RC area. Outcrops of the overlying Cambrian Tapeats Formation are only rotated 5–10 degrees by Tertiary normal faults. As a result, the larger block rotations may reflect normal faults active during development of the ophiolite. Normal faults and rotated blocks occur along the mid-ocean ridges and in well-exposed ophiolites. Rotated dikes define grabens in the Troodos Ophiolite (Varga and Moores, 1985), and in the Josephine Ophiolite, entire crustal sections were rotated as much as 50 degrees relative to the overlying sediments and underlying Moho (Harper, 1984). Proving their syn-ophiolite origin, vertical dikes cut the rotated sections of sheeted dikes in these examples. This relationship has not yet been found in the PO. The most prominent fold of the sheeted dike complex occurs in RC area (‘s’, Fig. 2). The dikes and foliation in gabbroic screens rotate about 75 degrees clockwise at they approach, and end at, a poorly exposed, northeast-trending boundary with a younger granite. On

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an outcrop scale, dextral shear zones and faults are intruded by dikes. Consequently, this boundary may have been a dextral strike-slip shear zone active during development of the ophiolite. A system of northeast-trending strike-slip shear zones, including the boundaries of the Mazatzal block, is consistent with arc-parallel strike-slip faults that commonly occur within island arcs. Fold-and-thrust structures in the Mazatzal Mountains east of the Tertiary valley (Fig. 2) record two episodes of convergent deformation, D2 and D3 , separated by the unconformity at the base of the Mazatzal Group (section Y–Y , Fig. 2). Overall NE-trending structures, in addition to the northeasterly trend of Yavapai-Mazatzal orogenic belts across North America (Fig. 1C), indicate the presence of a NE-trending convergent margin, subduction zones, and island arcs during the main phase of crustal assembly. In the Mazatzal block, the ca. 1.70 Ga Yavapai Orogeny produced the northeast-trending syncline in the turbidites overlying the ophiolite (‘f’ and section Y–Y , Fig. 2). Crustal thickening drove uplift, subaerial exposure, and erosion, marked by an unconformity that truncated the fold in the turbidites and, to the south, cut down into the sheeted dikes. The unconformity is closely associated with eruption of the ca. 1.70 Ga, subaerial, Red Rock Rhyolite and intrusion of hypabyssal equivalents and large sheets of granite (i.e., Payson Granite, Fig. 2). Only subvolcanic facies are preserved in the map area (Fig. 2). Siliciclastic sedimentation of the Mazatzal Group covered the unconformity. Deposition of these rocks in a foreland setting suggests that the earlier Yavapai Orogeny involved accretion of the Mazatzal block to North America. Coaxial with the Yavapai structures, the Mazatzal Orogeny produced a foreland system of thrust faults and related folds in the Mazatzal Group, defining the D3 -phase of convergent deformation. Despite the deformation of the bedded sequences east of the Tertiary valley, the plutonic core of the exposed ophiolite was not folded.

8. TECTONOSTRATIGRAPHIC ANALYSIS The tectonostratigraphic history of the Mazatzal block is recorded in three distinct units: (1) the 1.75 Ga basement complex, (2) the 1.73 Ga PO and overlying basin-filling sedimentary and submarine volcanic sequences (1.73–1.71 Ga), and (3) 1.70 Ga subaerial rhyolite and related granite and later fluvial to shallow-marine siliciclastic sediments (Figs. 1C, 2). The 1.75 Ga granitoids of the basement complex and the 1.73 Ga PO are incompatible with the tectonostratigraphic history (Fig. 1C) of the adjacent Ash Creek block and other blocks to the northwest (Fig. 1A). The Ash Creek block records submarine arc volcanism and northwest-trending D1 deformation prior to intrusion of the 1.735 Ga Cherry Batholith (Fig. 1C; Karlstrom and Bowring, 1991). Northwest-trending dikes or other evidence of rifting while the PO was forming are lacking. Likewise, the basement complex shows no indication of the northwest-trending D1 deformation. Juvenile Nd isotopic signatures of the Ash Creek block contrast with evidence for a LREE-enriched component in all rocks of the Mazatzal block (Dann et al., 1993). This period of incompatibility between adjacent blocks suggests that movement along Moore Gulch shear zone juxtaposed the Ash Creek and Mazatzal crustal blocks. The lack of evidence for a subduction zone or low angle fault

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that juxtaposed the two crustal blocks is best explained by a strike-slip boundary. Late dip-slip movement along Moore Gulch shear zone juxtaposes rocks from different crustal depths and obscures early fabrics. The PO is overlain by a volcano-sedimentary sequence, characterized by a lack of stratigraphic continuity across the Mazatzal block due to both facies changes and structural imbrication. In the map area, the basalts of the PO are directly overlain by dacitic volcaniclastic breccia with lenses of jasper and then a thick sequence of turbidites (Fig. 5A). Ash beds within the turbidites have U/Pb zircons ages of ca. 1.72 Ga (Dann et al., 1989). West of the map area (‘H’, Fig. 2), andesitic flows and coarse volcaniclastic breccias are interbedded with, and overlain by, turbidites and pelitic sediments with ash beds. These relations indicate that andesitic arc volcanoes erupted before and during turbidite deposition (Anderson, 1989) and on, or adjacent to, the PO. Finally, three granodiorite plutons (Fig. 2) intruded the PO at ca. 1.71 Ga (Conway et al., 1987). Anderson (1989) mapped a sequence of slates across the Moore Gulch shear zone, suggesting an overlap and contiguity of the two crustal blocks. In addition, ca. 1.70 Ga Yavapai deformation is recorded on both sides of the shear zone. Consequently, the Ash Creek and Mazatzal blocks must have been juxtaposed prior to the main phase of convergent deformation. After Yavapai deformation, the Ash Creek and Mazatzal blocks underwent differential uplift, probably reflecting different crustal profiles established early in their development. This difference is best appreciated by noting that the Mazatzal crustal block is unique, not only for the PO but also for preserving at least 4 unconformities and 3 transitions from plutonic to coeval volcanic rocks. Apparently, the Mazatzal block was uplifted less than adjacent blocks, suggesting less over-thickening during convergent deformation. The unique character of the crustal section probably began with the origin and emplacement of the ophiolite.

9. ORIGIN AND EMPLACEMENT OF THE PAYSON OPHIOLITE Ophiolites commonly originate in marginal basins above subduction zones because this tectonic setting predisposes them to be incorporated into continental crust during accretion of arcs or collision of continents. The formative tectonic setting and mechanics of emplacement are closely related, a theme that is important to any tectonic model of ophiolites. Seafloor spreading produced the horizontal layered structure of the PO as indicated by (1) the laterally extensive sheeted dike complex, (2) rooting of the sheeted dikes in coeval gabbro, (3) the abrupt transition to submarine lavas, and (4) the distribution of hydrothermal alteration and its exhalative products. The supra-subduction zone signature of the mafic rocks and older and younger arc lithologies implies that seafloor spreading took place within a magmatic arc. Within a 40 m.y. period, the basement complex developed within an arc, extensional tectonics rifted the arc and opened an intra-arc basin by seafloor spreading, arc volcanics and derived turbidites filled the basin, and arc granitoids intruded the basin floor. In modern arc settings, seafloor spreading creates both intra-arc and backarc submarine basins.

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The D2 and D3 convergent deformations affecting the Mazatzal block reflect a northeast-trending convergent margin that controlled the assembly of arc terranes and/or collisional orogenesis. Northwest-trending dikes of the PO formed by northeast-southwest extension, parallel to the convergent boundary prior to continental assembly. The schematic block diagram in Fig. 6A shows the Mazatzal and Ash Creek blocks prior to the ca. 1.70 Ga D2 Yavapai orogeny. In a Cretaceous back-arc basin represented by the Rocas Verdes ophiolite (de Wit and Stern, 1981), dikes are parallel to, and extension was perpendicular to, the convergent boundary. So, if the PO formed above a northeast-trending subduction zone, a simple back-arc model does not fit. Alternatively, arc-parallel extension occurs in modern arcs along arc-parallel strike-slip faults. For example, the Marinduque intra-arc basin in the Philippines developed from a pull-apart structure to seafloor spreading along the arc-parallel, strike-slip Philippine fault zone (Sarewitz and Lewis, 1991). This small basin, about the size of the Mazatzal block, contains a fossil axial-spreading center orthogonal to the strike-slip faults (Fig. 6B). Volcanoes rise from the basin floor. Turbidites are pouring in and interfingering with the volcanic debris. Strike-slip faults were active during seafloor spreading. When the basin crust is finally uplifted and exposed, we probably would see gabbro, sheeted dikes, and submarine volcanics overlain by turbidites and, like the PO, screens and roof pendants of older arc crust. In addition, we might see evidence for strike-slip shear zones within the ophiolite. Since the ophiolite remains above the subduction zone after its formation, we might expect to see arc plutons intruding the basin crust. As the basin is transported along the fault, the ophiolitic crust will be juxtaposed with arc crust that formed 100’s of km away in the same arc (Sarewitz and Lewis, 1991), probably with a contrasting tectonostratigraphic history. The Philippine arc is a collage of terranes including fragments of older deformed crust and several ophiolites, juxtaposed along major faults with up to 1000 km of displacement. An important feature of this model is that this juxtaposition occurs prior to the main phase of convergent deformation that accretes the arc to the continent. The Payson Ophiolite may be the oldest example of this mode of ophiolite generation and emplacement, attesting to the complex evolution of Early Proterozoic magmatic arcs leading up to the assembly of continental crust. High estimated crustal growth rates, indicated by large areas of juvenile crust like Yavapai-Mazatzal orogenic belt, suggest that continental assembly is episodic. Patchett and Chase (2002) estimate a 16% probability for margin-parallel strike-slip movement > 400 km that could effectively concentrate juvenile crust in small regions, giving the impression of higher than actual crustal growth rates. Direct evidence for early strike-slip movement along terrane boundaries that are reactivated during convergent deformation and post-assembly differential uplift is inherently difficult to recover. As a result, only by piecing together the tectonostratigraphic histories of terranes can the role of strike-slip tectonics in the assembly of continental crust be appreciated. Further analysis of the Payson Ophiolite and associated rocks is needed to better understand plate tectonics and the timing and mechanics of continental assembly during the Early Proterozoic.

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Fig. 6. (A) Schematic model of Mazatzal and Ash Creek crustal blocks at the time of formation of the Payson Ophiolite (ca. 1.73 Ga) and prior to D2 deformation of the ca. 1.70 Ga Yavapai Orogeny. The contrast in tectonostratigraphies requires an active boundary to juxtapose the distinct terranes. Arc-parallel extension in a step-over zone within a system of arc-parallel strike-slip faults rifted older arc crust and culminated in seafloor spreading and formation of an intra-arc basin. The basin was a locus of deposition within the arc. Transported along the arc by strike-slip movement, the basin was juxtaposed with the Ash Creek block prior to convergent deformation. (B) Similar in size to the Mazatzal block, the Marinduque intra-arc basin formed by seafloor spreading along the Philippine strike-slip fault zone in the Philippine arc (modified from Sarewitz and Lewis, 1991).

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Greene, T., 1989. Magmatic evolution of the Lokken SSZ ophiolite, Norwegian Caledonides: relationships between anomalous lavas and high-level intrusions. Geological Journal 24, 251–274. Harper, G.D., 1984. The Josephine ophiolite, northwestern California. Geological Society of America Bulletin 95, 1009–1026. Harper, G.D., Bowman, J.R., Kuhns, R., 1988. A field, chemical, and stable isotope study of subseafloor metamorphism of the Josephine ophiolite, California-Oregon. Journal of Geophysical Research 93, 4625–4656. Hawkesworth, C.J., Gallagher, K., Hergt, J.M., McDermott, F., 1993. Mantle and slab contributions in arc magmas. Annual Review of Earth and Planetary Sciences 21, 175–204. Hoffman, P.E., 1989. Precambrian geology and tectonic history of North America. In: Bally, A.W., Palmer, A.R. (Eds.), The Geology of North America—An Overview. Geological Society of America, The Geology of North America A, pp. 447–512. Karlstrom, K., Bowring, S.A., 1988. Early Proterozoic assembly of tectonostratigraphic terranes in southwestern North America. Journal of Geology 96, 561–576. Karlstrom, K.E., Bowring, S.A., 1991. Styles and timing of Early Proterozoic deformation in Arizona: constraints on tectonic models. In: Karlstrom, K.E. (Ed.), Proterozoic Geology and Ore Deposits of Arizona. Arizona Geological Society Digest 19, 1–10. Karlstrom, K.E., Doe, M.F., Wessels, R.L., Bowring, S.A., Dann, J.C., Williams, M.L., 1990. Juxtaposition of Proterozoic crustal blocks: 1.65–1.60 Ga Mazatzal orogeny. In: Gehrels, G.E., Spencer, J.E. (Eds.), Geologic Excursions through the Sonoran Desert Region, Arizona, and Sonora. Arizona Geological Survey Special Paper 7. Moores, E.M., 1982. Origin and emplacement of ophiolites. Reviews of Geophysics and Space Physics 20, 735–760. Moores, E.M., 2002. Pre-1 Ga (pre-Rodinian) Ophiolites: Their tectonic and environmental implications. Bulletin of the Geological Society of America 114, 80–95. Nicholas, A., Boudier, F., 1991. Rooting of the sheeted dike complex in the Oman ophiolite. In: Peters, T., Nicholas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of Oceanic Lithosphere. Kluwer Academic, Dordrecht, The Netherlands, pp. 39–54. Pallister, J.S., 1981. Structure of the sheeted dike complex of the Samail ophiolite near Ibra, Oman. Journal of Geophysical Research 86, 2661–2672. Patchett, P.J., Chase, C.G., 2002. Role of transform continental margins in major crustal growth episodes. Geology 30, 39–42. Pearce, J.L., Lippard, S., Roberts, S., 1984. Characteristics and tectonic significance of suprasubduction zone ophiolites. In: Kokelaar, B.P., Howells, M.F. (Eds.), Marginal Basin Geology: Volcanic and Associated Sedimentary and Tectonic Processes in Modern and Ancient Marginal Basin. Blackwell Scientific Publications, Oxford, England, pp. 77–94. Pederson, R.B., 1986. The nature and significance of magma chamber margins in ophiolites: examples from the Norwegian Caledonides. Earth and Planetary Science Letters 77, 100–112. Pederson, R.B., Malpas, C.A., 1984. The origin of oceanic plagiogranites from the Karmoy ophiolite, Western Norway. Contributions to Mineralogy and Petrology 88, 36–52. Rosencrantz, E., 1983. The structure of sheeted dikes and associated rocks in the North Arm massif, Bay of Islands ophiolite complex, and the intrusive process at oceanic spreading centers. Canadian Journal of Earth Sciences 20, 787–801. Rothery, D.A., 1983. The base of a sheeted dyke complex, Oman ophiolite: implications for magma chambers at oceanic spreading axes. Journal of the Geological Society of London 140, 287–296.

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Sarewitz, D.R., Lewis, S.D., 1991. The Marinduque intra-arc basin, Philippines: basin genesis and in situ ophiolite development in a strike-slip setting. Geological Society of America Bulletin 103, 597–614. Skjerlie, K.P., Furnes, H., 1996. The gabbro-dyke transition zone demonstrated on Tviberg, SolundStavfjord ophiolite complex. Geological Magazine 133, 573–582. Stillman, C.J., 1987. A Canary island dyke swarm: Implications for the formation of oceanic islands by extensional fissural volcanism. In: Halls, H.C., Fahrig, W.F. (Eds.), Mafic Dike Swarms. Geological Association of Canada Special Paper 34, 243–255. Varga, R.J., Moores, E.M., 1985. Spreading structure of the Troodos ophiolite, Cyprus. Geology 13, 846–850. Walker, G.P.L., 1987. The dike complex of Koolau volcano, Oahu: Internal structure of a Hawaiian rift zone. U.S. Geological Survey Professional Paper 1350, 961–993. Wessels, R.L., Karlstrom, K.E., 1991. Evaluation of the tectonic significance of the Proterozoic Slate Creek shear zone in the Tonto Basin area. In: Karlstrom, K.E. (Ed.), Proterozoic Geology and Ore Deposits of Arizona. Arizona Geological Society Digest 19, 193–210. Wrucke, C.T., Conway, C.M., 1987. Geologic map of the Mazatzal Wilderness and contiguous roadless areas, Gila, Maricopa, and Yavapai counties, Arizona. United States Geological Survey OpenFile Report 87-664, scale 1:48,000.

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Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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NEOPROTEROZOIC OPHIOLITES OF THE ARABIAN-NUBIAN SHIELD ROBERT J. STERNa , PETER R. JOHNSONb , ALFRED KRÖNERc AND BISRAT YIBASd a Geosciences

Department, University of Texas at Dallas, Box 830688, Richardson, TX 57083-0688, USA b Saudi Geological Survey, P.O. Box 54141, Jiddah 21514, Saudi Arabia c Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, Germany d Pulles Howard and de Lange, Environmental and Water Quality Management, P.O. Box 861, Auckland Park 2006, South Africa

Ophiolites of mid-Neoproterozoic age are abundant in the Arabian-Nubian Shield (ANS) of NE Africa and Arabia. ANS ophiolites range in age from 690 to 890 Ma and litter a region that is 3000 km N-S and > 1000 km E-W. In the northern ANS, ophiolites occur as nappe complexes marking suture zones between terranes. Although dismembered and altered, all of the diagnostic components of ophiolites can be found: harzburgite, cumulate ultramafics, layered as well as higher level gabbro and plagiogranite, sheeted dikes, and pillowed basalt. Allochthonous mafic-ultramafic complexes in the southern ANS, in Ethiopia and Eritrea, are interpreted as ophiolites, but are more deformed and metamorphosed than those in the north. Reconstructed ophiolitic successions have crustal thicknesses of 2.5 to 5 km. The ANS ophiolitic mantle was mostly harzburgitic, containing magnesian olivines and spinels that have compositions consistent with extensive melting. Cr# for spinels in ANS harzburgites are mostly > 60, comparable to spinels from modern forearcs and distinctly higher than spinels from mid-ocean ridges and backarc basin peridotites. ANS ophiolites are often associated with a thick (1–3 km) sequence of cumulate ultramafic rocks, which define a transition zone between seismic and petrologic Mohos. These cumulates are dominated by dunite, with subordinate pyroxene-rich lithologies. Cumulate ultramafics transition upwards into layered gabbro. Several crystallization sequences are inferred from ANS transition zones and cumulate gabbro sections. In all samples studied, olivine and spinel crystallized first, followed (in order of decreasing abundance) by cpx-plag, cpxopx-plag, and opx-cpx-plag. ANS ophiolitic lavas mostly define a subalkaline suite characterized by low K and moderate Ti contents, that has both tholeiitic and calc-alkaline affinities and includes a significant, although subordinate, amount of boninites. The lavas are fractionated (mean Mg# = 55) but have higher abundances of Cr (mean = 380 ppm) and Ni (mean = 135 ppm) than would be expected for such a low Mg#. The ANS ophiolitic lavas include both LREE-depleted and LREE-enriched varieties, but as a group are slightly DOI: 10.1016/S0166-2635(04)13003-X

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LREE-enriched: mean (Ce/Yb)n ∼ 2.2 and Sm/Nd ∼ 0.30. On a variety of discrimination diagrams, the lavas plot in fields for MORB, BABB, arc tholeiite, and boninite. ANS lavas cluster around Ti/Zr = 97, indicating that Ti-bearing phases did not precipitate early. Nd isotopic compositions indicate derivation from a long-depleted mantle source, with εNd(t) ∼ +5 to +8. Mineral and lava compositions are consistent with the hypothesis that most ANS ophiolites formed in ‘suprasubduction zone’ (SSZ) settings, and the high Cr# of ANS ophiolitic harzburgites suggests a forearc environment. Geochemical studies of deep water sediments deposited on ANS ophiolites are needed to better characterize and understand the Neoproterozoic ocean where ANS ophiolites formed.

1. INTRODUCTION The Arabian-Nubian Shield (ANS) in NE Africa and W. Arabia is the largest tract of juvenile continental crust of Neoproterozoic age on Earth (Patchett and Chase, 2002). This crust was generated when smaller terranes of arc and back arc crust were generated within and around the margins of a large oceanic tract known as the Mozambique Ocean, which formed in association with the breakup of Rodinia ∼ 800–900 Ma (Stern, 1994). Oceanic plateaus may also have been accreted (Stein and Goldstein, 1996). These crustal fragments collided as the Mozambique Ocean closed, forming arc-arc sutures, composite terranes, the Arabian-Nubian Shield (ANS; Fig. 1), and the larger collisional belt known as the East African Orogen (Stern, 1994; Kusky et al., 2003). The Arabian-Nubian Shield was caught between fragments of East and West Gondwanaland as these collided at about 600 Ma (Meert, 2003) to form a supercontinent variously referred to as Greater Gondwanaland (Stern, 1994), Pannotia (Dalziel, 1997) or just Gondwanaland. The ANS was subsequently buried by Phanerozoic sediments but has been exposed by uplift and erosion on the flanks of the Red Sea in Oligocene and younger times. Several lines of evidence support the idea that the ANS is juvenile Neoproterozoic crust, including non-radiogenic initial Sr and radiogenic initial Nd isotopic compositions for a

Fig. 1. Location of the Arabian-Nubian Shield and location of ophiolites and related rocks within it. (A) Political and modern geographic features of the region. B = Bahrain, Dj = Djibouti, Is = Israel, Jrdn = Jordan, K = Kuwait. Location of Figs. 2A and B shown in dashed rectangles. (B) Location of Precambrian basement exposures, crustal types, and ophiolites and ophiolitic rocks (shown in black). Abbreviations for some of the better studied ophiolites follow. Saudi Arabia: H = Halaban, JT = Jebel Tays; JU = Jebel Uwayjah; JE = Jebel Ess; AA = Al ‘Ays (Wask); BT = Bi’r Tuluhah; A = Arjah; DZ = Darb Zubaydah; BU = Bi’r Umq; Th = Thurwah; JN = Jebel Nabitah; T = Tathlith. Egypt: F = Fawkhir; Br = Barramiya; Gh = Ghadir; AH = Allaqi-Heiani; G = Gerf. Sudan: OSH = Onib-Sol Hamed; Hs = Hamisana; AD = Atmur-Delgo; K = Keraf; R = Rahib; M = Meritri; Os = Oshib; Kb = Kabus (Nuba Mts); I = Ingessana; Kk = Kurmuk. Eritrea: Hg = Hagar Terrane. Ethiopia: Zg = Zager Belt; DT = Daro Tekli Belt; Bd = Baruda; TD = Tulu Dimtu; Y = Yubdo; A = Adola; My = Moyale. Kenya: S = Sekerr, B = Baragoi. The Bi’r Umq-Nakasib suture is defined by the Bi’r Umq-Thurwah-Meritri-Oshib ophiolite belt.

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wide range of igneous rocks. The ANS can be isotopically defined as the region in NE Africa and Arabia where Nd-model ages approximate crystallization ages (Stern, 2002). These indications that abundant juvenile continental crust and mantle lithosphere were generated during Neoproterozoic time in the region has been confirmed by Nd and Sr isotopic studies of samples of mafic lower crust and mantle lithosphere brought up as xenoliths in Tertiary lavas from Saudi Arabia, which also indicate that the lower crust and lithospheric mantle of the region formed during Neoproterozoic time (Henjes-Kunst et al., 1994; McGuire and Stern, 1993). Ophiolites and ophiolitic rocks are remarkably abundant in the Arabian-Nubian Shield. They are scattered across most of the ANS, over a distance of ∼ 3000 km from the farthest north (Jebel Ess) almost to the equator, and from Rahib in the west to Jabal Uwayjah (45 ◦ E) in the east, encompassing an area of about two million square kilometers (Fig. 1). Ophiolites are particularly well studied in Arabia (see companion paper by Johnson et al. (2004). If ophiolites are the remains of oceanic lithosphere, then the ANS is a massive graveyard of Neoproterozoic oceanic lithosphere. The abundance of ophiolites is a further indication that ANS crust and lithosphere were produced by processes similar to those of modern plate tectonics. Their abundance has made it difficult to define the orientation of sutures solely from the distribution of ophiolitic rocks (Church, 1988; Stern et al., 1990). This is further complicated by the fact that not all ANS mafic-ultramafic complexes formed in a seafloor spreading environment—some appear to be roots of island arcs, such as Darb Zubaydah in Arabia (Quick and Bosch, 1989)—and others are autochthonous layered intrusions, such as Dahanib in Egypt (Dixon, 1981). Care must be exerted to avoid misidentification, but even the most conservative estimates indicate that there is a remarkable abundance of ophiolites in the ANS. Ophiolites were first recognized in the region by Rittmann (1958), but were otherwise ignored until the pioneering study of Bakor et al. (1976). This was followed by a flurry of field, petrological, and geochemical studies throughout the 1980’s. This level of activity has decreased significantly through the 1990’s and into the 21st century. This review has three objectives. First, it is intended to summarize the most important observations of this first phase of studying ANS ophiolites. Second, because the ophiolites of the Arabian-Nubian Shield are so common and so well-exposed, it is hoped that this example will provide a basis of what is expected to be preserved when a major episode of juvenile crust formation associated with modern-style plate tectonics occurs. Finally, it is hoped that this overview will stimulate a resurgence of ANS ophiolite studies. Note that the Arabic word for mountain is variously spelled ‘Jebel’ (Egypt), ‘Gebel’ (Sudan), or ‘Jabal’ (Arabia). For simplicity, we use the ‘Jebel’ spelling throughout this contribution.

2. OUTCROP PATTERNS We define the ANS as the northern, juvenile part of the Neoproterozoic EAO. More restrictive definitions have the shield ending in the south with the southernmost contiguous

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outcrops of basement exposed around the Red Sea. The Arabian-Nubian Shield as defined here includes basement outliers well to the south in Ethiopia and Kenya. For the purposes of this presentation, the ANS is that part of the East African Orogen characterized by juvenile Neoproterozoic crust and is also where ophiolites and ophiolitic rocks are encountered (Fig. 1). The ANS can be usefully subdivided into northern and southern halves. North of the Bi’r Umq-Nakasib suture, which extends NE from the Oshib and Meritri ophiolites in Sudan and continues across the Red Sea in Arabia through the Thurwah and Bi’r Umq ophiolites, the ophiolite belts trend approximately E-W. It is relatively easy to identify structures developed during ophiolite obduction in this region. Greenschist-facies metamorphism is characteristic for these ophiolites, and diagnostic features, such as pillowed basalts, are well preserved. This is shown in Table 1, which lists ophiolites with a Penrosetype succession; these are from the northern ANS or from the Bi’r Umq-Nakasib suture. Such excellent preservation largely reflects the fact that these northern ophiolites are relatively undisturbed by steep N-S structures developed during terminal collision between E and W Gondwanaland, which are pervasive farther south (Fig. 2A). Many ophiolites in the northern ANS are, however, disrupted by NW-trending strike-slip faults and shear zones of the Najd fault system (Sultan et al., 1988). Remote sensing has proven to be an effective way to map the distribution of spectrally distinct lithologies, especially serpentinites and amphibole-bearing mafic rocks, providing a quantitative, if indirect, assessment of the distribution and abundance of disrupted ophiolites in the basement of Egypt (Sultan et al., 1986). A more detailed presentation of the field relations and structure of Arabian Shield ophiolites is presented by Johnson et al. (2004). It is a significantly greater challenge to identify ophiolites to the south of the Bi’r Umq-Nakasib suture. This area was closer to and thus more intensely affected by the endNeoproterozoic terminal collision, such that structures related to ophiolite obduction are transposed or obliterated (Abdelsalam and Stern, 1996). Basement structures dip steeply (Fig. 2B), units are intensely deformed and shuffled by high-angle thrusting and subhorizontal shearing, and metamorphism is typically amphibolite-facies (Yihunie, 2002). Diagnostic features and emplacement fabrics for units that might originally have been ophiolites are not common. Purists would hesitate to identify the linear mafic-ultramafic complexes of the southern ANS as ophiolites, but the association of harzburgitic ultramafics in association with MORB-like and even boninitic mafic units as well as the regional association of southern ANS mafic-ultramafic complexes to the abundant and unequivocal ophiolites of the northern ANS makes it likely that these mostly represent ophiolites in different stages of preservation.

3. CRUSTAL STRUCTURE The best preserved ophiolites in the northern ANS contain all or most of the components of complete ‘Penrose’ ophiolites (Table 1), including pillowed basalts, gabbros,

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Table 1. ANS ophiolites with a “Penrose-type” assemblage Name Darb Zubaydah Bi’r Tuluhah Wadi Khadra Halaban Ess Al ‘Ays (Wask) Tharwah Bi’r Umq Tathlith Oshib-Ariab belt Arbaat Atmur-Delgo

Country Saudi Arabia Saudi Arabia Saudi Arabia Saudi Arabia Saudi Arabia Saudi Arabia Saudi Arabia Saudi Arabia Saudi Arabia Sudan Sudan Sudan

Sol Hamed Wadi Onib Gebel Gerf Wadi Ghadir Fawkhir

Sudan Sudan Egypt Egypt Egypt

Reference Quick (1990) Pallister et al. (1988) Quick (1991) Al-Saleh et al. (1998) Pallister et al. (1988) Bakor et al. (1976) Nassief et al. (1984) Shanti (1983) Pallister et al. (1988) Abdel-Rahman (1993) Abdelsalam and Stern (1993) Harms et al. (1994), Schandelmeier et al. (1994) Fitches et al. (1983) Hussein et al. (1983) Zimmer et al. (1995) El-Bayoumi (1983) El-Sayed et al. (1999)

and tectonized harzburgites. Several ANS ophiolites have sheeted dike complexes (such as Ghadir, Onib, and Ess) but these are not always reported. Even the well-preserved ophiolites are faulted, folded, and otherwise disrupted, so that reconstructing a complete ophiolite pseudostratigraphy is difficult and equivocal. Nevertheless, three such reconstructions of ANS ophiolite crustal structure are shown on Fig. 3. These reconstructions differ in the relative abundances of volcanics, pillowed lavas, sheeted dikes, and gabbro but all suggest that the oceanic crust represented by ANS ophiolites was generally in the range of 2.5 to 6 km thick. As discussed in the next section, ANS ophiolites were generated and em-

Fig. 2. Remote sensing images of ANS ophiolites, showing the different outcrop patterns of ophiolites in the northern (A) and southern (B) parts of the Arabian-Nubian Shield. Locations shown in Fig. 1. (A) Allaqi-Heiani Suture along the Egypt-Sudan border. Image is approximately 300 km across and N is towards the top of image. Dashed line approximates trace of Allaqi-Heiani Suture. Note that the general E-W structure of the ophiolite belt, which formed during suturing of the SE Desert and Gabgaba terranes is only disrupted by younger N-S structures (developed during terminal collision between East and West Gondwanaland) of the N-S Hamisana Shear Zone. This outcrop pattern indicates the ophiolites and associated accretionary structures are subhorizontal. NASA astronaut photograph (S32-74-100). (B) Landsat TM image of basement units in N. Eritrea and E. Sudan, showing dominance of complex deformation related to terminal collision between East and West Gondwanaland, resulting in ∼ N-S structures. Terrane names are modified after Drury and Filho (1998). Scene is about 90 km across.

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Fig. 3. Reconstructed crustal sections of ANS ophiolites. Al ‘Ays (Al Wask), Saudi Arabia (Bakor et al., 1976); Bi’r Umq, Saudi Arabia (Al-Rehaili and Warden, 1980); Onib, Sudan (Kröner et al., 1987).

placed relatively early in the history of the ANS and EAO. Fragments from dismembered ophiolites are common in the ANS, and it becomes more difficult to interpret these as once being allochthonous pieces of oceanic crust as these fragments become more deformed and metamorphosed. Suffice it to say that not all ultramafic rocks in the shield are—or were— parts of ophiolites. There are many layered igneous intrusions containing non-ophiolitic ultramafics and gabbros and there are many examples of non-ophiolitic pillowed lavas. Nevertheless, the association of harzburgitic ultramafics and low-K tholeiitic metabasalt argues strongly that these once belonged to a coherent ophiolite.

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Fig. 4. Ages of ophiolites in the Arabian-Nubian Shield, binned at 25 million year intervals. Neoproterozoic time (544 to 1000 Ma) is subdivided into Tonian (1000–850 Ma), Cryogenic (850 to ∼ 600 Ma) and Neoproterozoic III (∼ 600–544 Ma), after Knoll (2000). Age range when Arabian-Nubian Shield was tectonically and magmatically active (870 Ma to end of Neoproterozoic) is also given. Ophiolite ages are U-Pb and Pb-Pb zircon ages and Sm-Nd ages from (Stacey et al., 1984; Claesson et al., 1984; Pallister et al., 1988; Kröner et al., 1992; Zimmer et al., 1995; Worku, 1996). Mean age for ophiolites ±1 standard deviation is given.

4. AGE ANS ophiolites have been reliably dated using U-Pb zircon techniques (Stacey et al., 1984; Pallister et al., 1988) and Pb-Pb zircon evaporation techniques (Kröner et al., 1992; Zimmer et al., 1995) on zircons separated from gabbros and plagiogranites. Other ages have been generated using Sm-Nd mineral and whole-rock techniques (Claesson et al., 1984; Zimmer et al., 1995; Worku, 1996). These results give age ranges of 694 ± 8 Ma for the youngest ANS ophiolite (Urd/Halaban; Stacey et al., 1984) to 870 ± 11 Ma for the oldest (Thurwah; Pallister et al., 1988). A mean of 781 Ma (1 standard deviation = 47 Ma) is obtained for 16 robust ophiolite ages (Fig. 4). Ophiolites formed during the first half of the time period encompassed by tectonic and magmatic activity of the Arabian-Nubian Shield. There is no obviously systematic geographic variation in the distribution of ANS

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ophiolite ages; the oldest ophiolite is in the middle of the ANS and the youngest is near the eastern edge of ANS exposures.

5. OPHIOLITE COMPONENTS 5.1. Harzburgites Where the original protolith can be identified, the mantle peridotites associated with ANS ophiolites are predominantly tectonized harzburgites, lherzolite being rarely reported. Studies of ANS ophiolite fabrics are at a very early stage, but harzburgites and dunites of the Thurwah ophiolite have been tectonized at high temperatures (Nassief et al., 1984), and a similar high-temperature ductile fabric is reported from the Sol Hamed ophiolite (Fitches et al., 1983). The harzburgites represent residual mantle after extensive melting, whereas the dunites and wehrlites are cumulates or reflect melt-wallrock interactions. Harzburgites are mostly altered (Figs. 5A, B), but relict olivines and pyroxenes have been analyzed by electron microprobe for 5 ophiolites. Harzburgite associated with the Al Ays ophiolite contains olivine of Fo91–94, orthopyroxene of En89–91, and clinopyroxene of En48–53, Fs2.2–3.0, Wo44–49.2 (Chevremont and Johan, 1982a; Ledru and Auge, 1984). Harzburgite from the nearby Hwanet ophiolite has orthopyroxene of En88.9–91 and clinopyroxene of En48.9–50.3, Fs2.3–3.4, Wo46.3–48.9 (Chevremont and Johan, 1982a). Harzburgite associated with the Ess ophiolite contains olivine of Fo91–93, orthopyroxene of En88–92, and diopsidic clinopyroxene of En49–52, Fs2.4–3.1, Wo44.6–48.3 (Al-Shanti, 1982). Nassief et al. (1984) identified a mantle sequence that is up to 20 km thick for the Thurwah ophiolite, and harzburgite from this consists of 70–90% olivine (Fo89.5–93.4), 15–30% orthopyroxene (En90–92) and < 1% each of chromite and clinopyroxene. Mouhamed (1995) argued on the basis of CIPW normative compositions of Muqsim serpentinites along the AllaqiHeiani suture (Fig. 2) that these were originally harzburgites. Olivines from the Ingessana ophiolite are Fo91–97 (Price, 1984). These compositions are at the magnesium-rich end of peridotites, as shown on Fig. 6. Olivine compositions provide insights into the tectonic setting of ophiolites because magmagenesis in different tectonic settings reflects differing extents of melting. Because residual olivines become increasingly magnesian as melting progresses, residual mantle should have olivines that are more magnesian than the Fo88 of undepleted ‘pyrolitic’ upper mantle (Fig. 6). Bonatti and Michael (1989) suggest that mantle melting ranges from nearly zero for undepleted continental peridotites to about 10–15% melting for rifted margins to 10–25% melting associated with mid-ocean ridge (MOR) peridotites to 30% for peridotites recovered from forearcs, which generally form during the early stages in the evolution of the associated subduction zone (Bloomer et al., 1995). Mantle peridotites from back-arc basins were not available when this diagram was originally generated, but since that time mantle peridotites from the Mariana Trough active back-arc basin in the western Pacific have been studied (Ohara et al., 2002). These harzburgites have olivine compositions that are indistinguishable from MOR harzburgites and are distinctly less magnesian

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Fig. 5. Outcrop photographs of ANS ophiolites. (A) Barramiya, Egypt. Light areas are talc-carbonate alteration of harzburgitic ultramafics, darker areas are serpentinized ultramafics. White house for scale. (B) Thurwah, Saudi Arabia. Light areas are talc-carbonate alteration of harzburgitic ultramafics, darker areas are serpentinized ultramafics. Seated geologist (far left) for scale. (C) Carbonated ultramafics of the Bi’r Umq ophiolite (Saudi Arabia) thrust south over metasediments of the Mahd Group. Thrust contact is dashed. Three geologists climbing ridge for scale. (D) N-dipping layered gabbros and cumulate ultramafics on the south side of the Gerf ophiolite, SE Egypt/NE Sudan. Ridge is about 100 m tall. (E) Layered gabbros of the Onib ophiolite, NE Sudan. Light layers are anorthositic, darker layers are richer in mafic minerals. Card (∼ 15 cm long) for scale. (F) Pillowed basalts, Wadi Zeidun (near Fawkhir ophiolite), Egypt. Hammer for scale.

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Fig. 5. (Continued.)

than those of forearc peridotites. In this context, the Mg-rich nature of ANS ophiolite peridotite olivines is best interpreted to indicate that these are residual after extensive melting, similar to that observed for forearc peridotites. Inferences about the extent of melting based on olivine compositions are supported by abundant compositional data for spinels. Spinel resists alteration better than olivine, may be economically important, and thus are relatively well studied for ANS ophiolites (Fig. 7). Progressive melting of peridotites depletes Al relative to more refractory Cr in residual spinels, such that the Cr# of spinels (= 100Cr/Cr + Al) increases with melting

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Fig. 6. Composition of olivines in peridotites from various tectonic settings, modified after Bonatti and Michael (1989). The tectonic settings are arranged in an order in which melt depletion increases to the right. Note that the field for olivines from peridotites of 4 ANS ophiolites are Fo91 or are more magnesian. These compositions are most consistent with formation in a forearc setting.

(Dick and Bullen, 1984). As is the case for Fo content of harzburgitic olivines, the increase in Cr# of harzburgitic spinels reflects increasing degree of melting. Degree of melting is related in a general way to tectonic setting, at least for Cenozoic peridotites (Fig. 7A). Spinels from MOR peridotites generally have Cr# < 50 (although Barnes and Roeder, 2001 report that a subordinate proportion of MOR peridotites have Cr# up to 80). A limited dataset for spinels from back-arc basin peridotites indicates that these experienced extents of melting similar to MOR and thus have spinel Cr# that are similar to those of MOR peridotites (Ohara et al., 2002). Spinels in forearc harzburgites generally have higher Cr#

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Fig. 7. Composition of spinels in ANS ophiolitic peridotites compared with those in modern peridotites. Data are plotted on 100Cr/Cr + Al (Cr#) vs. 100Mg/Mg + Fe (Mg#) diagram, modified after Dick and Bullen (1984). (A) Spinels from ANS ophiolitic peridotites, location and size of rectangle determined by means and standard deviations listed in Table 1. (B) Fields defined by spinel compositions for likely ophiolitic analogues of Cenozoic age, modified after Bloomer et al. (1995). Note that spinels from peridotites from mid-ocean ridges and back-arc basin spreading axes characteristically have Cr# < 60, whereas spinels from forearc peridotites and boninites Cr# > 40. The vast majority of ANS ophiolitic spinels have high Cr#, suggesting these formed in a forearc setting.

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Table 2. Mean compositions of spinels in ANS harzburgites Locality Sudan Onib/Sol Hamed Oshib Rahib Ingessana

Mean Cr#

1 Std. Dev.

Mean Mg#

1 Std. Dev.

70.4 71.9 73.4 77.5

17.6 9.2 11.7 7.0

56.4 65.1 56.5 57.6

11.9 5.4 22.3 14.7

Egypt Various

52.3

5.8

56.6

4.9

Saudi Arabia Ess/Al ‘Ays Nabitah Al Amar Thurwah Bi’r Umq Tuluhah/Arjah Halaban

67.2 78.3 62.9 79.8 52.6 72.4 76

10.0 4.9 14.2 6.4 7.8 4.4 6.8

60.6 64.7 52.4 56.0 61.0 70.2 67.4

9.9 4.5 5.4 2.8 4.1 8.0 2.9

Ethiopia Adola Moyale

91.4 58.2

2.5 2.8

63.1 64.4

7.6 2.6

Kenya Sekerr Baragoi

88.1 82.6

1.3 1.1

68.7 56.7

7.9 5.1

Data sources: Onib-Sol Hamed, Sudan (N = 60: Abdel-Rahman, 1993; Hussein, 2000; Price, 1984); Oshib, Sudan (N = 15: Abdel-Rahman, 1993); Rahib, Sudan (N = 15: Abdel-Rahman, 1993); Ingessana-Kurmuk, Sudan (N = 97: Abdel-Rahman, 1993; Price, 1984); Egypt (various; 4 averages for spinels in harzburgite: Ahmed et al., 2001); Ess-Al ‘Ays, Saudi Arabia (N = 81: Al-Shanti and El-Mahdy, 1988; Al-Shanti, 1982; Chevremont and Johan, 1982a; Chevremont and Johan, 1982b; Ledru and Auge, 1984); Nabitah, Saudi Arabia (N = 9: Al-Shanti and El-Mahdy, 1988); Bi’r Umq, Saudi Arabia (N = 20: LeMetour et al., 1982); Thurwah, Saudi Arabia (N = 9: Al-Shanti and El-Mahdy, 1988); Tuluhah/Arjah, Saudi Arabia (N = 14: Al-Shanti and El-Mahdy, 1988); Halaban, Saudi Arabia (N = 8: Al-Shanti and El-Mahdy, 1988); Adola, Ethiopia (N = 10: Bonavia et al., 1993); Moyale, Ethiopia (N = 32: Berhe, 1988); Sekerr, Kenya (N = 123: Price, 1984); Baragoi, Kenya (N = 50: Berhe, 1988).

(up to 80) and spinels from boninites typically have Cr# of 70–90 (Fig. 7B), consistent inferences from olivine compositions that these are manifest residues and products of the highest degree of melting found for post-Archean igneous rocks. Table 2 summarizes spinel compositions from peridotites (mostly harzburgites) associated with 16 ANS ophiolites. These are mostly for spinels in harzburgites, but where this data was unavailable, compositions of podiform chromite were used. The data indicate that ANS peridotitic spinels have Cr# that are mostly > 60 (Egypt, Bi’r Umq, and Moyale being notable exceptions). Spinels from the Sekerr ophiolite (mean Cr# = 88) and the Adola ultramafic complex (mean Cr# = 91) are remarkably rich in chromium. Overall, the Cr#

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of ANS ophiolitic peridotites appear most similar to those of modern forearc peridotites (Fig. 7B), although there is a hint of a regional gradient, from higher Cr# in the south to lower Cr# in the north. Podiform chromites from the Adola complex in southern Ethiopia are associated with harzburgites; Bonavia et al. (1993) did not explicitly identify this as part of an ophiolite, but they argued from the absence of phase layering and low Rh/Ir that the rocks did not form as part of a layered intrusion. These data supports an interpretation that the Adola mafic-ultramafic complex represents a highly metamorphosed, deformed ophiolite. This inference is supported by the interpretation of Yibas et al. (2003) that many mafic-ultramafic complexes in southern Ethiopia are ophiolites. 5.2. Transition Zone and Gabbros The transition zone lies between the petrologic Moho (defined at the base of the cumulate section and the top of the tectonized peridotites) and the seismic Moho (defined as the top of cumulate ultramafic section and the base of gabbros). Several ANS ophiolites have well-preserved transition zones characterized by interlayered pyroxenite, wehrlite, lherzolite, dunite, and/or chromite at the base (Fig. 5D) grading upwards into sections that are increasingly dominated by gabbro. Other ophiolites have thin to non-existent transition zones (e.g., Fawkhir; El-Sayed et al., 1999). The Onib ophiolite is characterized by an unusually thick (2–3 km) transition zone of interlayered cumulate ultramafics, podiform chromites, and layered gabbros (Fig. 3; Hussein et al., this volume; Kröner et al., 1987). The ultramafic rocks of the Sol Hamed ophiolite, NW of Onib, also seem to represent a similar crust-mantle transition zone, and are 80% dunite (with interbedded chromitites), with lesser proportions of wehrlite, harzburgite, and pyroxenite (Fitches et al., 1983; Price, 1984). Similarly, the Oshib ophiolite has a 2-km thick transition zone, which grades upwards from harzburgite tectonite through dunite and wehrlite followed by pyroxenite and cumulate gabbro (Abdel-Rahman, 1993). The transition zone of the Thurwah ophiolite contains about 1 km of cumulate dunite, lherzolite, and pyroxenite, in similar proportions (Nassief et al., 1984). Each rock layer is 1–10 m thick and is developed in the upward sequence dunite-lherzolite-pyroxenite. The Ess ophiolite contains a cumulate peridotite section that is about 400 m thick and consists of wehrlite and dunite (Shanti and Roobol, 1979). Significant bodies of dunite, often associated with podiform chromite, exist at the base of some gabbro sections (El-Bayoumi, 1983). Olivines in dunites from the Thurwah ophiolite are significantly more Fe-rich (Fo86.5–89.5; Nassief et al., 1984) than olivines of typical ANS harzburgites. The Al Ays ophiolite contains over 350 lenses of podiform chromitite within the dunite unit (Bakor et al., 1976). Chromites associated with dunites are often significantly more Cr-rich (Cr# = 65–85) than those associated with harzburgite (Cr# ∼ 50; Ahmed et al., 2001); the reason for this is not understood. Some chromites contain inclusions of olivine (Fo97–99) and diopside, along with hydrous phases, such as amphibole (edenite-tremolite) and phlogopite (Ahmed et al., 2001). The primary mineralogy of these inclusions is remarkably preserved, and studying inclusions in chromites and other resistant minerals is a promising avenue for understanding ANS ophiolites.

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Modal abundances of olivine decrease upsection as the abundance of especially clinopyroxene increases; correspondingly dunites are upwardly succeeded by wehrlites or lherzolites and then pyroxenites. Cumulate pyroxenites are dominated by diopside. Cumulate pyroxenites of the Onib transition zone are low in TiO2 (0.01–0.06%) and rich in Cr (1500– 3000 ppm; Kröner et al., 1987). Gabbro bodies up to 1 km thick are interleaved with the ultramafic cumulates at Igariri, part of the Oshib ophiolite (Abdel-Rahman, 1993). ANS ophiolites with well-developed transition zones may have formed by fast seafloor spreading, whereas those that do not have thick transition zones may represent seafloor produced by slow spreading (Dilek et al., 1998). Gabbros are ubiquitous and important components of ANS ophiolites. Original igneous textures are common but metamorphic recrystallization under greenschist to amphibolite facies conditions is ubiquitous. ANS ophiolitic gabbros are mostly pyroxene gabbro; olivine gabbro is much less common. Clinopyroxene generally dominates over orthopyroxene. Where ophiolitic gabbros are well preserved and studied, they are commonly layered, at least in part. Phase layering, evidenced by melanogabbro and anorthositic gabbro couplets are typically ∼ 5 cm thick and can be traced for meters. Layered metagabbros made up of alternating plagioclase (An63–83)- and amphibole-rich layers are part of the Ess ophiolite (Shanti, 1983). Plagioclase compositions in Sol Hamed gabbros change from An70–85 at the base to more sodic compositions upsection (Fitches et al., 1983). Amphibole may be uralitized pyroxene (El-Bayoumi, 1983). Similar uralitized clinopyroxene gabbros are reported for the Al Ays (al Wask) ophiolite. Upwards through the cumulate sequences, from the base of the transition zone into the layered gabbros, the sequence of rocks indicates a crystallization sequence of olivine ± chromite-clinopyroxene-plagioclase (Price, 1984), olivine ± chromite-clinopyroxene-orthopyroxene-plagioclase (Nassief et al., 1984), or, less commonly, olivine ± chromite-orthopyroxene-clinopyroxene-plagioclase (AbdelRahman, 1993). High level gabbros are more massive, lack layering and other evidence of crystal accumulation, and commonly include pegmatitic gabbro and isolated bodies of plagiogranite. ANS plagiogranites are high in SiO2 (70–77%) and low in K2 O (0.04–1.9%: Shanti, 1983; Abdel-Rahman, 1993). These mostly plot in the field of trondhjemite and tonalite on a normative Ab-An-Or diagram (Shanti, 1983). The high level gabbros are often intruded by mafic dikes, which represent the base of the sheeted dike complex. 5.3. Sheeted Dykes Sheeted diabase dikes are common components of ANS ophiolites. Where observed they typically transition downwards into the high-level gabbros and grade upwards into pillowed basalts. Sheeted dikes are reported from the following ophiolites: Rahib (Abdel-Rahman et al., 1990), Ess (Shanti and Roobol, 1979), Sol Hamed (Fitches et al., 1983), Ingessana (Price, 1984), Ghadir (El-Bayoumi, 1983), Gerf (Zimmer et al., 1995), and Thurwah (Nassief et al., 1984). Sheeted dykes for many other ANS ophiolites are not identified or are poorly developed (e.g., Al Ays and Fawkhir: Bakor et al., 1976; El-Sayed et al., 1999). It

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is not yet clear whether this represents differential preservation of diagnostic dike-on-dike textures or represents original differences in crustal sections. 5.4. Pillowed Basalts and Other Volcanics Lavas with well-preserved pillow structures are diagnostic parts of ANS ophiolites (Fig. 5F). These are especially important because lava compositions provide valuable clues about tectonic setting and the nature of melt generation. This section thus concentrates on the chemical composition of ANS ophiolitic lavas. We recognize that the greenschist-facies metamorphism that these basalts have suffered probably disrupted their chemical composition, but it is likely that these effects may cancel each other out (e.g., Mg gain in one lava, Mg loss in another). For the purpose of understanding the compositional variability of ANS ophiolitic lavas, we compiled available chemical data for these. A total of 200 samples were used in the compilation, including 37 analyses from Egypt (El-Sayed et al., 1999; Stern, 1981; Zimmer et al., 1995), 52 from Sudan (Price, 1984; AbdelRahman, 1990, 1993; Harms et al., 1994; Hussein, 2000), 32 from Arabia (excluding Darb Zubaydah; Al-Shanti, 1982; Bakor et al., 1976; Kattan, 1983; Nassief et al., 1984; Shanti, 1983), and 29 from the Sekerr ophiolite, Kenya (Price, 1984). An additional 50 samples from Eritrea (Woldehaimanot, 2000) and Ethiopia (Wolde et al., 1996; Woldehaimanot and Behrmann, 1995) were used to compare the metavolcanic sequences of suspected ophiolites in these countries. These data are summarized in Table 3. Table 3. Mean composition of ANS ophiolitic pillowed basalts Egypt

Arabia

Sudan

Ethiopia & Eritrea N 37 32 52 50 50 ± 3 49 ± 3 49 ± 4 SiO2 (%) 51 ± 4 1.3 ± 0.4 1.0 ± 0.5 1.1 ± 0.6 0.7 ± 0.7 TiO2 9.8 ± 2.4 9.4 ± 2.2 10.0 ± 1.8 11.2 ± 3.0 FeO∗ MgO 6.2 ± 2.0 6.9 ± 1.7 7.5 ± 2.2 9.7 ± 5.9 0.18 ± 0.13 0.28 ± 0.31 0.15 ± 0.21 0.43 ± 0.8 K2 O Mg# 52 ± 8 56 ± 7 56 ± 9 58 ± 14 Sr (ppm) 147 ± 144 143 ± 47 191 ± 114 192 ± 205 Ba 84 ± 81 25 ± 21 70 ± 44 72 ± 102 Y 33 ± 13 21 ± 9 24 ± 7 13 ± 10 Zr 87 ± 36 65 ± 44 70 ± 37 42 ± 42 V 320 ± 90 230 ± 75 287 ± 75 84 ± 67 Cr 295 ± 150 326 ± 212 356 ± 238 467 ± 591 Ni 110 ± 56 99 ± 76 113 ± 84 155 ± 246

Sekerr DPL 16 46 ± 2 1.4 ± 0.3 9.9 ± 1.0 8.4 ± 3.7 0.14 ± 0.09 58 ± 8 257 ± 135 99 ± 86 25 ± 7 111 ± 20 222 ± 40 621 ± 535 304 ± 305

Sekerr UPL 13 47 ± 2 2.7 ± 0.7 11.2 ± 0.9 4.6 ± 1.2 0.48 ± 0.35 42 ± 7 466 ± 85 259 ± 241 33 ± 7 171 ± 15 227 ± 19 160 ± 154 69 ± 81

ANS mean 200 49 ± 4 1.2 ± 0.7 10.2 ± 2.3 7.6 ± 3.8 0.26 ± 0.46 55 ± 11 199 ± 161 85 ± 112 23 ± 12 76 ± 49 234 ± 108 380 ± 392 135 ± 174

Data sources: Egypt (El-Sayed et al., 1999; Stern, 1981; Zimmer et al., 1995), 52 from Sudan (Abdel-Rahman et al., 1990; Price, 1984; Abdel-Rahman, 1993; Harms et al., 1994; Hussein, 2000); Arabia (excluding Darb Zubaydah; Al-Shanti, 1982; Bakor et al., 1976; Kattan, 1983; Nassief et al., 1984; Shanti, 1983); Kenya (Price, 1984); Eritrea (Woldehaimanot, 2000) and Ethiopia (Wolde et al., 1996; Woldehaimanot and Behrmann, 1995).

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ANS ophiolitic lavas are dominantly basalts, with a mean of 49% SiO2 (1 standard deviation = 4% SiO2 ). There are some andesitic and even a few dacitic samples identified as part of the ophiolitic sequence (e.g., Fawkhir, Egypt). These lavas average 1.2% (±0.7%) TiO2 . ANS ophiolitic pillow basalts define a low-K suite, with mean K2 O = 0.26%. These major element characteristics are similar to a wide range of oceanic lavas, including MORB, some intra-oceanic arc lavas, and back-arc basin basalt (BABB). Ophiolitic ANS lavas are often fractionated, with Mg# (= 100Mg/Mg + Fe) = 55 ± 11. Mg# for basaltic melt in equilibrium with mantle peridotite is expected to be in the range 65–70. The mean of 135 ppm Ni (±174 ppm) and 380 ppm Cr (±392 ppm) is higher than would be expected for a mean Mg# = 55. Eritrean and Ethiopia suspected ophiolitic lavas typically contain less TiO2 , Y, and Zr and have higher K2 O and Mg#, Cr and Ni than most other ANS ophiolitic lavas. Upper pillow lavas from the Sekerr ophiolite are also distinct, with much higher TiO2 , Y, and Zr along with lower Mg#, Cr, and Ni than most other ANS ophiolitic sequences. The Sekerr UPL sequence may represent an alkalic succession, something that is rarely found for other ANS ophiolitics. ANS ophiolitic lavas are mostly tholeiitic but also include some calc-alkaline examples (Fig. 8). There is no obvious geographic variation. Samples from Kenya and Eritrea are dominated by tholeiites whereas those from Ethiopia are predominantly calc-alkaline. Most samples from Sudan, Egypt, and Arabia are tholeiitic, but with a significant proportion of calc-alkaline representatives as well. REE data for 67 samples from the ANS (including several from suspected ophiolites in Ethiopia and Eritrea) indicate that both LREE-enriched and LREE-depleted varieties exist. A simple way to see the extent to which samples are LREE-enriched or LREEdepleted is by use of the chondrite normalized Ce/Yb ratio, or (Ce/Yb)n . If (Ce/Yb)n > 1, the REE pattern is generally LREE-enriched (and the higher the ratio, the greater the LREE-enrichment). Similarly, if (Ce/Yb)n < 1, the sample is LREE-depleted. The mean (Ce/Yb)n for ANS ophiolitic lavas is 2.18, but with a very large standard deviation of 3.07. This largely results from one unusual sample or analysis (MV33 with (Ce/Yb)n = 21.9; El-Sayed et al., 1999), and omitting this sample lowers the mean (Ce/Yb)n to 1.89 ± 1.88. Another way to summarize the REE patterns is to look at Sm/Nd. The chondritic value, taken to approximate the bulk Earth, is ∼ 0.325, so values greater than this are LREEdepleted and values less than this are LREE-enriched. The mean for ANS ophiolitic lavas (0.30 ± 0.13) again indicates modest LREE-enrichment. Including data for Eritrea and Ethiopia lowers the mean Sm/Nd to 0.29 ± 0.10. A variety of discriminant diagrams can be applied to examine the tectonic affinities of ANS ophiolitic lavas. A plot of V vs. Ti (Fig. 9) shows that most samples have Ti/V between 20 and 50, and plot in fields defined by mid-ocean ridge and back-arc basin basalts. This includes most of the samples from Egypt, Sudan, and Sekerr, Kenya, downstream pillow basalts. A substantial portion of the ophiolitic basalts have Ti/V < 20 and so plot in the field of island arc tholeiites or calc-alkaline basalts. This includes a subordinate proportion of samples from Egypt and Sudan. No field for boninitic rocks is shown on these diagrams, but boninites have very low Ti and V and plot near the origin. Eritrea and Ethiopia samples

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Fig. 8. Magmatic affinities of ANS ophiolitic pillow lavas and suspected ophiolitic metavolcanics from Eritrea and Ethiopia. (A) TiO2 vs. FeO∗ /MgO. (B) SiO2 vs. FeO∗ /MgO. Field boundaries after Miyashiro (1975).

have low Ti/V and some have very low Ti contents, supporting suggestions that these are boninitic. Only a subset of the upstream pillow lavas from the Sekerr ophiolite plot in the field of within-plate basalts.

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Fig. 9. Ti-V discrimination diagram for basalts, after Shervais (1982). IAT = arc tholeiite, MORB & BABB = mid-ocean ridge basalt and back-arc basin basalt, WPB = within-plate basalt. Calc-alkaline basalts have low Ti concentrations and a wide range of Ti/V and plot in the grey field.

A plot of Cr vs. Y shows similar mixed tectonic affinities for ANS ophiolitic lavas (Fig. 10). Samples from Egypt, Sudan, and downstream pillow lavas from the Sekerr ophiolite mostly plot in the MORB field (note that there is no field for back-arc basin basalts on this diagram). Samples from Arabia plot in the MORB field and in the field for arc basalts or even on the low-Y side of the field for arc basalts. Samples from Eritrea and Ethiopia plot within the field for arc basalts and to the low-Y side of the arc field. Again, only a subset of the upstream pillow lavas from the Sekerr ophiolite plot in the field of within-plate basalts. Finally, Zr-based discriminant diagrams (Fig. 11) show the mixed affinities of ANS ophiolitic lavas. On a plot of Ti vs. Zr (Fig. 11A), the samples are remarkably coherent, clustering about a mean Ti/Zr = 97 ± 41 (excluding Ethiopia and Eritrea), decreasing slightly to Ti/Zr = 95 ± 41 when data from Ethiopia and Eritrea are included. In contrast to Fig. 8, a calc-alkaline trend is not seen on the Ti-Zr diagram. Fig. 11A suggests that the ophiolitic lavas are simply related by differing degrees of melting and fractionation, with Sekerr ophiolite samples representing low degrees of melting (and/or extensive fractionation), whereas samples plotting near the origin (most samples from Ethiopia-Eritrea and some samples from Arabia, Sudan, and Egypt) are relatively high degree melts. Most of the

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Fig. 10. Cr-Y discrimination diagram for basalts (Pearce, 1982). Note that there is no field for back-arc basin basalts.

Sekerr upstream pillow lavas and a sample from Al Ays (sample 5 of Bakor et al., 1976), a sample from the Uogame basalts of Eritrea (ER54 of Woldehaimanot, 2000), and OPL/2 from Onib (Abdel-Rahman, 1993) lie along the extension of the MORB field in a region where within-plate basalts plot. Zr/Y vs. Zr systematics reveal similar affinities and also shows that a large proportion of samples plot to the low-Zr side of fields defined by arc basalts, MORB, and within-plate basalts. This is the area where boninites plot, using the summary data of Crawford et al. (1989), and samples from Arabia, Sudan, and especially Ethiopia and Eritrea plot in this field. 5.5. Sediments Pelagic sediments that rest immediately on the uppermost part of ANS ophiolites include dolomite and ribbon chert (e.g., Al Ays: Bakor et al., 1976; Hussein, 2000). These limestones are inferred to be similar to modern pelagic carbonates, whereas the cherts are thought to be due to silica-rich emanations without the involvement of radiolaria. Metased-

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Fig. 11. Zr-based discriminant diagrams for ANS ophiolitic basalts. (A) Ti vs. Zr (Pearce and Cann, 1973). Field labels: I = arc tholeiites, II = MORB, III = calc-alkaline basalt. A fourth field, situated within I but obscured by data points is defined by all three fields. No field for within-plate lavas is shown, but these plot along the upper right extension of the MORB field. (B) Zr/Y vs. Zr (Pearce and Norry, 1979). Gray field shows composition of boninites, from Table 1-1 of Crawford et al. (1989).

iments overlying the Ess ophiolite include shale and minor conglomerate (Shanti and Roobol, 1979). Ophiolites in the Central Eastern Desert of Egypt are overlain by sediments that include diamictite, tuffaceous siltstones, and banded iron formation (Stern, 1981; Sims and James, 1982). Future ophiolite studies should take care to carefully characterize the nature of the sedimentary sequence immediately above the pillowed lavas.

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All too little attention has been paid to the nature and composition of sediments that were deposited on ANS ophiolites, probably because investigations of these ophiolites have mostly been led by petrologists and igneous geochemists. Because the Neoproterozoic was an exceptional period in Earth history, characterized by extensive—perhaps global— glaciation, unparalleled excursions in seawater stable isotopic compositions, and explosive radiation of metazoa (Evans, 2000), sediments of this age are of increasing interest to the global geoscientific community. Efforts to understand the Neoproterozoic ‘Snowball Earth’ have concentrated to date on shallow water sequences, with very little of the record preserved in deep-sea sediments examined. Because sediments deposited on ANS ophiolites should preserve an excellent record of the composition of deep water of the Neoproterozoic ocean, these should become an increasing focus of research in the near future. 5.6. Alteration and Obduction-Related Metamorphism The ultramafic rocks associated with ANS ophiolites are generally highly altered, but it is often not known whether this alteration occurred before, during, or after emplacement. These rocks are largely converted to serpentinite or to mixtures of serpentine, talc, tremolite, magnesite, chlorite, magnetite, and carbonate. These rocks are variously called ‘talc-carbonate schists’, ‘Barramiya rocks’, or ‘listwaenite’. Strictly speaking, listwaenite should be reserved for rocks that are fuchsite-quartz-carbonate lithologies derived from ultramafic rocks by potassic and carbonate metasomatism (Halls and Zhao, 1995), but common usage in the ANS is more general. Silicified serpentinites are sometimes called ‘birbirites’ (Augustithus, 1965). Basta and Kader (1969) reported that lizardite is the main constituent of Egyptian serpentinites, whereas Akaad and Noweir (1972) identify antigorite as volumetrically dominant. The talc-carbonate rocks mainly consist of magnesite (± dolomite) and talc. The origin of the carbonate alteration fluids remains to be elucidated, but Stern and Gwinn (1990) argued on the basis of C and Sr isotopic studies that carbonate intrusions in the Eastern Desert of Egypt—which could be related to the carbonatizing fluids affecting ANS ultramafic rocks—are mixtures of mantlederived and remobilized sedimentary carbonate. Certainly the prevalence of carbonate alteration of ANS ophiolitic ultramafics suggests a tremendous flux of CO2 -rich fluids from the mantle during middle and late Neoproterozoic time (Newton and Stern, 1990; Stern and Gwinn, 1990). In contrast, Surour and Arafa (1997) argued that the ‘ophicarbonates’ of the Ghadir ophiolite are reworked oceanic calcites that formed after it was obducted. Regardless of how carbonatization of the ophiolitic ultramafics occurred, it has economic implications. A spatial and genetic relationship has been observed between carbonatized ultramafics, subsequent granite intrusions, and gold mineralization. Apparently the carbonatization preconcentrates gold up to 1,000 times that in the original ultramafic rocks, and interaction with hydrothermal systems associated with granite intrusions may further concentrate gold (Cox and Singer, 1986). Thrust contacts are documented at the base of some, but not all, ANS ophiolites (Fig. 5C). Metamorphic soles of ANS ophiolites have been studied along the west-

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ern part of the Allaqi-Heiani ophiolite belt in SE Egypt, where El-Naby and Frisch (1999) inferred temperatures of up to 700 ◦ C and pressures up to 8 kbar. Other thrusts at the base of the ophiolites are associated with no obvious thermal effects. Chloritites are developed around the peripheries of some ultramafic masses (Takla, 1991; Takla et al., 1992). The metamorphic sole of the Halaban ophiolite consists of tightly folded amphibolites, hornblendites, gneisses, magmatirestites, rodingites and serpentinized ultramafics, all of which are cut by granitic dikes associated with partial melting (Al-Saleh et al., 1998). The metamorphic sole of the Halaban ophiolite in Arabia is divided into a western segment that is dominated by greenschist to amphibolite facies metasedimentary rocks, structurally overlain by higher grade amphibolite-facies meta-igneous rocks. Such inverted metamorphic zonation is common in sub-ophiolitic complexes (Al-Saleh et al., 1998). Most amphibolites and meta-sediments in the latter segment have experienced varying degrees of calcium metasomatism due to the release of excess calcium during serpentinization, and the original assemblages have often been rodingitized. The Halaban ophiolitic sole includes exotic blocks of serpentinite and metamorphosed ultramafic rock, interpreted to be derived from the basal peridotites of the Halaban Ophiolite that became fragmented as thrusting progressed and were later incorporated within the underlying amphibolites.

6. ISOTOPIC DATA There has been a modest amount of isotopic work conducted on ANS ophiolites, including two ophiolites in Arabia, a concentration of work around the Gerf ophiolite in SE Egypt, and data for the Adola mafic-ultramafic complex. The most reliable isotopic data for ANS ophiolites comes from Sm-Nd isotopic systematics. This is because alteration makes it difficult to rely on Sr and Pb isotopic compositions, whereas Sm and Nd are relatively immobile and corrections for in situ radiogenic growth is simple if the age is known. These results are summarized in Table 4, which shows that all ANS ophiolites studied to date have strongly positive εNd (+5.0 to +7.7). This indicates that these melts were generated from a long-depleted (high Sm/Nd) mantle source. The Nd isotopic data indicate an asthenospheric source and a juvenile, ensimatic setting. There is a hint that the mantle source Table 4. Neodymium isotopic composition of ANS ophiolites Ophiolite Ess (2) Al ‘Ays (Wask) Harga Zarqa (10) Heiani (7) Gerf (20) Adola

Age (Ma) 780 743 750 750 750 789

εNd(t) 6.9 7.6 7.6 ± 0.5 7.7 ± 0.8 6.8 ± 0.7 5.0

Numbers in parentheses refer to number of samples used to calculate mean value.

Reference Claesson et al. (1984) Claesson et al. (1984) Zimmer et al. (1995) Zimmer et al. (1995) Zimmer et al. (1995) Worku (1996)

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for the northern ophiolites may have been more—or longer—depleted than those farther south, but the data at present are too sparse for this to be anything more than a suggestion. Os isotope data on chromites from the Al Ays ophiolite have recently been reported (Walker et al., 2002); these straddle the line defined for the evolution of the primitive mantle.

7. DISCUSSION Ophiolites of mid-Neoproterozoic age (690 to 890 Ma; mean = 781 ± 47 Ma) are abundant in NE Africa and Arabia. Ophiolites encompass an area of 3000 km N-S and > 1000 km E-W. These ophiolites are in various stages of dismemberment and alteration, but all of the diagnostic components can be found, including harzburgite, cumulate ultramafics, layered as well as higher level gabbro and plagiogranite, sheeted dikes, and pillowed basalt. Reconstructions of a few ophiolitic successions indicate a crustal thickness of 2.5 to 6 km. Many ANS ophiolites mark suture zones where smaller terranes coalesced; these suture zones indicate the location of fossil subduction zones. Some ANS ophiolites were emplaced while still hot enough to metamorphose underlying rocks, while others were emplaced cold. This gross tectonic setting is simplest to explain if ANS ophiolites generally were located on the hanging wall of a convergent plate margin and were emplaced when buoyant crust entered the subduction zone. Such an event could cause the subduction zone to fail, suturing the two terranes at the same time that the ophiolite was emplaced (Cloos, 1993). This describes a forearc setting for such ophiolites, an interpretation that finds increasing favor in the scientific community (Shervais, 2001). Mineral and lava compositions as well as limited isotopic data are consistent with the hypothesis that most ANS ophiolites formed in ‘suprasubduction zone’ (SSZ) settings. Most ANS ophiolites have the hallmarks of forearc ophiolites. Harzburgites are the most common type of mantle peridotite, and these contain magnesian olivines and spinels with compositions that indicate large extents of melting. Limited data for relict olivines in harzburgites show these to be significantly more Mg-rich than peridotites recovered from modern mid-ocean ridges and similar to olivines in harzburgites recovered from forearcs. Limited data for olivines from backarc basin peridotites are indistinguishable from MORB peridotites, arguing against a back-arc basin setting. This interpretation is consistent with spinel compositions. Cr# for spinels in ANS harzburgites are mostly > 60, again most like those recovered from modern forearcs and distinctly higher than those from mid-ocean ridges and the admittedly sparse database for backarc basin peridotites. ANS ophiolites are often associated with a thick (1–3 km) sequence of cumulate ultramafic rocks, which define a transition zone between the seismic and petrologic Mohos. These cumulates consist of a very high proportion of dunite but there are also a lot of pyroxene-rich lithologies. Chromites associated with dunites are often more Cr-rich than those in the underlying spinels. These cumulate ultramafics transition upwards into layered gabbro. Together these cumulate sequences indicate that the extensional magmatic systems represented by ANS ophiolites experienced significant fractionation. Several crystal-

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lization sequences are inferred from ANS transition zones and cumulate gabbro sections. Olivine and spinel are always first to crystallize: especially cpx-plag and cpx-opx-plag; opx-cpx-plag is rare. These are crystallization sequences (respectively) C, B, and A of Pearce et al. (1984), identified as hallmarks of SSZ ophiolites. This sequence contrasts with the crystallization sequence olivine-plagioclase-clinopyroxene which Pearce et al. (1984) interpreted to be diagnostic of ophiolites that formed at true mid-ocean ridges. The latter crystallization sequence has not been reported for ANS ophiolites. The overall sense of ANS ophiolites inferred from the transition zone and the gabbroic section is that of large, fractionating magma chambers that were repeatedly tapped and recharged but which generally evolved from a system dominated by primitive mafic magmas to ones dominated by highly evolved magmatic liquids. ANS ophiolitic lavas mostly define a subalkaline suite characterized by low K and moderate Ti contents that is moderately fractionated. They reveal both tholeiitic and calcalkaline affinities and include a significant if subordinate proportion of boninites. We find little in the structure or composition of ANS ophiolites to support the hypothesis that ANS crustal growth entailed widespread involvement of oceanic plateaus or large igneous provinces, although rare examples of within-plate lavas are identified. ANS ophiolitic lavas are quite fractionated (mean Mg# = 55) but have higher abundances of Cr (mean = 380 ppm) and Ni (mean = 135 ppm) than would be expected for this relatively low Mg#. ANS ophiolitic lavas include both LREE-depleted and LREE-enriched varieties, but as a group are slightly LREE-enriched: mean (Ce/Yb)n ∼ 2.2 and Sm/Nd ∼ 0.30. On a variety of discrimination diagrams, ANS ophiolitic lavas plot in fields for MORB, BABB, and arc tholeiites, along with a significant proportion of lavas with strong boninitic affinities. ANS lavas cluster reasonably tightly around Ti/Zr = 97, indicating that Ti-bearing phases did not precipitate early. These mixed subalkaline characteristics are characteristic of SSZ ophiolitic lavas, and the presence of boninitic lavas in particular supports a forearc origin. There is a tremendous amount of work that needs to be done on ANS ophiolites. Most of the information that we have on ANS ophiolites was collected in the 1980s and early 1990s and the field is ripe for new perspectives and for detailed studies using modern techniques. There have been few modern geochemical studies. Most of the trace element data summarized in this overview was generated with XRF techniques. We need to analyze ANS ophiolitic basalts using modern plasma analytical techniques. Modern ICP-MS techniques result in much better precision and accuracy for many more elements, including Nb and Ta, which are essential to understand tectonic setting of igneous rocks. Interestingly, one of the few ANS ophiolites that have been studied with modern geochemical techniques—the Gerf ophiolite—has strong trace element affinities to true MORB. The geochronologic and isotopic database is limited and additional data are needed if we are to understand when the crust and mantle lithosphere represented by these ophiolites were generated as well as when a given ophiolite was emplaced. Limited Sm-Nd isotopic data indicate derivation from depleted asthenospheric mantle. Isotopic data however do not resolve tectonic setting. There are no Hf isotopic data for these ophiolites, and such analyses should provide a valuable perspective on the evolution of the Lu-Hf isotopic sys-

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tem in the Neoproterozoic mantle. We also need to better understand the composition of primary magmatic liquids; compositions of olivines and spinels in ANS harzburgites indicate unusually high degrees of melting so primary liquids should be very primitive but lava compositions are very fractionated. One way to resolve this apparent contradiction would be to analyze clinopyroxenes in ANS ophiolitic harzburgites using an ion microprobe to determine compositions of REE and other trace elements in equilibrium liquids. This would allow a more explicit linkage between the composition of ophiolitic lavas and the extent of melting in the associated mantle section to be examined. Another very useful strategy would be to identify and study inclusions in spinels, which have recently been discovered to preserve primary magmatic phases. One of the most interesting unresolved problems concerns the nature of the carbonate alteration of especially the ultramafics. Where did all this CO2 come from? Is there any relationship between this alteration and gold mineralization in the ANS? Several workers suggest a link between more depleted, boninitic ophiolites and gold mineralization; is there relationship between gold mineralization in the ANS and ophiolite type? Modern regions of seafloor spreading often are associated with hydrothermal vents and associated mineralization and biota, are such associations preserved in ANS ophiolites? Such studies may be especially important because ANS ophiolites may have formed at a time when a variety of independent geologic observations suggest that the surface of the earth was covered with ice (Evans, 2000), and life may have been restricted to such vents. Similarly, the association of banded iron formations overlying ANS ophiolites (Sims and James, 1982) may reflect interactions between deep water and hydrothermal activity during a snowball earth episode. Another aspect of ANS ophiolites worthy of study is a possible relationship between economic chromite deposits and lava compositions. In some ophiolites (e.g., Zambales, Philippines) the affinity of the magmatic section is a reliable guide to the grade of associated chromite deposits (Evans and Hawkins, 1989).

8. CONCLUSIONS The last three decades of research on ANS ophiolites allows us to sketch the broad outlines of how these complexes formed, but most of this work was completed a decade or more ago. New research initiatives that focus on ANS ophiolites are needed and these promise to be rewarding. A few examples of what needs to be done are presented below. ANS ophiolites change style on either side of the Bi’r Umq-Nakasib suture zone. Are these correlative? We are not certain that the allochthonous mafic-ultramafic complexes of Eritrea and Ethiopia should be considered as ophiolites, although we have argued for this. Compositions of spinels in harzburgite are a powerful tool for understanding the magmatic evolution and tectonic setting of ophiolites, but we have too few analyses for ophiolitic harzburgites from Egypt, Eritrea, and Ethiopia. Similarly, we need more mineral chemical data for primary ophiolite phases (olivine, clinopyroxene, plagioclase). We need new campaigns of petrologic and geochemical studies using the most modern analytical techniques

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(ICP-MS, etc.) to determine the diagnostic trace element contents of ANS ophiolitic lavas. Some modern analytical techniques, such as using ion probe analyses of clinopyroxenes to independently infer trace element compositions of equilibrium liquids, have not yet been applied. More isotopic studies, especially Nd, Hf, and Os, are needed to understand the evolution of the mantle source region for these lavas. Similarly, we need more geochronological constraints on the age of ANS ophiolites. Ion probe ages of zircons would be especially useful because this best identifies inherited components but this technique has not been applied to ANS ophiolites. Precise determination of ophiolite age will also be critical for studies of deep water marine sequences immediately overlying the pillow lavas, a research direction that is needed to understand global change during Neoproterozoic time. We should identify the most complete ANS ophiolites and study them in detail. Interdisciplinary study of the best-preserved ophiolites by teams of structural geologists, petrologists, geochemists, and geochronologists focusing on the ophiolite itself in tandem with sedimentologists, geochemists, and paleontologists studying the overlying sedimentary succession should be encouraged. Efforts should be made to identify ancient hydrothermal vent deposits and investigate the associated fossilized biota. Economic considerations also favor renewed study of ANS ophiolites. The high grade of ANS chromites may be rich enough to mine. Given the observation that ANS gold mineralization often appears to be related to carbonate alteration of ANS ophiolitic peridotites, we can expect to better understand the former by focused studies of the latter. At present we have a very poor understanding of pervasive carbonate alteration of ANS ophiolitic peridotites. Research programs should be designed to better understand the age of this alteration and the origins of these fluids.

ACKNOWLEDGEMENTS Thanks to T. Tadesse (Ethiopian Geological Survey) and B. Woldhaimanot (U. Asmara) for their thoughts on the ‘ophiolites’ of Eritrea and Ethiopia. Thanks to J. Encarnacion and Y. Dilek for thoughtful reviews. This manuscript is dedicated to the memory of Ian Gass, who stimulated so many excellent studies of ANS ophiolites and whose enthusiasm for the ophiolites of the Arabian-Nubian Shield has yet to be matched.

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Sultan, M., Arvidson, R.E., Duncan, I.J., Stern, R.J., Kaliouby, B.E., 1988. Extension of the Najd shear system from Saudi Arabia to the central Eastern Desert of Egypt based on integrated field and Landsat observations. Tectonics 7, 1291–1306. Sultan, M., Arvidson, R.E., Sturchio, N.C., 1986. Mapping of serpentinites in the Eastern Desert of Egypt by using Landsat thematic mapper data. Geology 14, 995–999. Surour, A.A., Arafa, E.H., 1997. Ophicarbonates: calcified serpentinites from Gebel Moghara, Wadi Ghadir area, Eastern Desert, Egypt. Journal of African Earth Sciences 24, 315–324. Takla, M.A., 1991. Chloritites at the contacts of some ophiolitic ultramafics, Eastern Desert, Egypt. Egyptian Mineralogist 3, 151–165. Takla, M.A., Basta, F.F., Surour, A.A., 1992. Petrology and Mineral Chemistry of Rodingites associating the Pan-African Ultramafics of Sikait-Abu Rusheid area, South Eastern Desert, Egypt. Geology of the Arab World, 491–507. Walker, R.J., Prichard, H.M., Ishiwatari, A., Pimentel, M., 2002. The Osmium isotopic composition of convecting upper mantle deduced from ophiolite chromites. Geochimica et Cosmochimica Acta 66, 329–345. Wolde, B., Asres, Z., Desta, Z., Gonzalez, J.J., 1996. Neoproterozoic zirconium-depleted boninite and tholeiitic series rocks from Adola, southern Ethiopia. Precambrian Research 80, 261–279. Woldehaimanot, B., 2000. Tectonic setting and geochemical characterization of Neoproterozoic volcanics and granitoids from the Adobha Belt, northern Eritrea. Journal of African Earth Sciences 30, 817–831. Woldehaimanot, B., Behrmann, J.H., 1995. A study of metabasite and metagranite chemistry in the Adola region (south Ethiopia): Implications for the evolution of the East African orogen. Journal of African Earth Sciences 21, 459–476. Worku, H., 1996. Geodynamic Development of the Adola Belt (Southern Ethiopia) in the Neoproterozoic and Its Control on Gold Mineralization. Verlag Dr. Köster, Berlin, p. 156. Yibas, B., Reimold, W.U., Anhaeusser, C.R., Koeberl, C., 2003. Geochemistry of the mafic rocks of the ophiolitic fold and thrust belts of southern Ethiopia: constraints on the tectonic regime during the Neoproterozoic (900–700 Ma). Precambrian Research 121, 157–183. Yihunie, T., 2002. Pan-African deformations in the basement of the Negele area, southern Ethiopia. International Journal of Earth Science 91, 922–933. Zimmer, M., Kröner, A., Jochum, K.P., Reischmann, T., Todt, W., 1995. The Gabal Gerf complex: A Precambrian N-MORB ophiolite in the Nubian Shield, NE Africa. Chemical Geology 123, 29–51.

Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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NEOPROTEROZOIC OPHIOLITES IN THE ARABIAN SHIELD: FIELD RELATIONS AND STRUCTURE PETER R. JOHNSONa , FAYEK H. KATTANa AND AHMED M. AL-SALEHb a Saudi Geological Survey,

PO Box 54141, Jiddah 21514, Saudi Arabia Department, King Saud University, PO Box 2009, Riyadh, Saudi Arabia

b Geology

Ophiolites make up a small but tectonically important part of the Arabian shield. Where most complete, they consist of serpentinized peridotite, gabbro, dike complex, basalt, and pelagic rocks. However, because of folding and shearing, the majority of the ophiolites lack one or more of these diagnostic lithologies. Nonetheless, the incomplete assemblages are identified as ophiolites because they minimally include peridotite and gabbro, in many cases are associated with basalt, and in all cases show evidence of emplacement by thrusting and shearing rather than intrusion. The ophiolites range in age from ∼ 870 Ma to ∼ 695 Ma, documenting a 200-million year period of oceanic magmatism in the Arabian shield, and are caught up in ∼ 780 Ma to ∼ 680 Ma suture zones that reflect a 100-million year period of terrane convergence. All the ophiolites are strongly deformed, metamorphosed, and altered by silicification and carbonatization. Low-grade greenschist facies metamorphism predominates, but in places the rocks reach amphibolite grade. Alteration resulted in the development of listwaenite, particularly in shear zones, and locally the only evidence that mafic-ultramafic rocks underlie a given area is the presence of upstanding ridges of listwaenite that are resistant to erosion. S/C fabrics are widespread and indicate that the ophiolites were affected by both strike-slip and vertical displacements. Variations in senses of shear observed along and across the strike evidence considerable strain partitioning during deformation. However, prevailing senses of shear can be discerned for several of the ophiolites that, in conjunction with other structural observations, indicate the main shear trajectories of the shear zones containing the ophiolites. Jabal Ess, Jabal Tharwah, and Bi’r Umq ophiolites were emplaced during periods of dextral transpression on the Yanbu and Bi’r Umq sutures, respectively. The Bi’r Tuluhah ophiolite was emplaced during sinistral transpression of the Hulayfah-Ad Dafinah-Ruwah suture, and the Halaban ophiolite was emplaced during west-directed convergence on the Halaban suture. DOI: 10.1016/S0166-2635(04)13004-1

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1. INTRODUCTION Neoproterozoic mafic-ultramafic complexes make up less than 1% of the surface area of the Arabian shield but, starting with the pioneering work of Al-Shanti and Mitchell (1976) and Bakor et al. (1976), they figure prominently in discussions of the origins of the shield because of their possible tectonic significance as remnants of oceanic crust and indicators of arc-arc suturing (Pallister et al., 1987; Stoeser and Camp, 1985). Certainly not all mafic-ultramafic complexes in the region are ophiolites—some are nondiagnostic lenses of sheared serpentinite, some are intrusions in the base of volcanic arcs, and some are layered intrusions—and care must be taken to avoid misidentification, misinterpretation, and spurious correlations (Church, 1988, 1991). Nevertheless, a significant number of complexes have the hallmarks of ophiolites. They are widespread in the shield along shear zones (Fig. 1) and, together with stratigraphic, geochronologic, and structural data, provide evidence of active ocean-floor magmatism in association with development of the tectonostratigraphic terranes in the shield and the process of suturing during terrane amalgamation (Pallister et al., 1988; Johnson and Woldehaimanot, 2003; Genna et al., 2002). The mineralogy, chemistry, and tectonic settings of Arabian and Nubian shield ophiolites are reviewed by R.J. Stern and colleagues in a companion chapter in this volume. The purpose of this report is to describe the lithology, structure, and field relations of selected Arabian shield ophiolites, thereby providing examples of Neoproterozoic ophiolites and illustrating the outcrop characteristics and degrees of dismemberment and structural complexity that may be expected of Neoproterozoic ophiolites elsewhere. Of the ophiolites selected, Jabal Ess lies on the Yanbu suture at the join between the Midyan and Hijaz terranes in northwestern Saudi Arabia (Fig. 1). The Jabal Tharwah and Bi’r Umq ophiolites lie on the Bi’r Umq suture between the Hijaz and Jiddah terranes. The Bi’r Tuluhah ophiolite is at the northern end of the Hulayfah-Ad Dafinah-Ruwah suture joining the Hijaz-JiddahAsir terranes and the Afif terrane. The Halaban and Jabal al Uwayjah ophiolites are parts of the Halaban suture between the Afif and Ad Dawadimi terranes, and the Jabal Tays ophiolite is within the Ad Dawadimi terrane east of the Halaban suture. The Jabal Ess, Jabal Tharwah, Bi’r Umq, Bi’r Tuluhah, Halaban, and Jabal al Uwayjah ophiolites are believed to be rooted in the shear zones with which they are associated and, as such, mark the sites of consumption of oceanic crust. The Jabal Tays ophiolite, in contrast, appears to be part of a structurally detached ophiolite allochthon far traveled from its root zone.

2. JABAL ESS OPHIOLITE The Jabal Ess ophiolite (Figs. 2, 3) comprises mantle peridotite, isotropic and layered gabbro, a dike complex, pillow basalt, and pelagic sediments metamorphosed in the greenschist facies and is the most complete ophiolite in the Arabian shield (Al-Shanti, 1982). It covers an area of approximately 30 km east-west, and 5 km north-south, and has a minimum estimated thickness of about 3 km. The ophiolite crops out in hills that rise 300 m

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Fig. 1. Mafic-ultramafic complexes and major faults and shear zones in the Arabian shield, Saudi Arabia with an inset showing terranes and interterrane ophiolite-decorated sutures and shear zones. Complexes that satisfy criteria as ophiolites, are named; other mafic-ultramafic complexes (mostly serpentinite lenses along fault zones) are shown by initials. Complexes: A = Al Amar; D = Ad Dafinah; H = Hamdah; HA = Hakran; I = Ibran; M = Muklar; N = Nabitah; R = Rahah; RD = Ar Ridaniyah; TA = Tabalah-Tarj; TW = Tawilah; UF = Umm Fawrah. Terranes: Af = Afif; As = Asir; D = Ad Dawadimi; H = Hijaz; Hi = Hail; J = Jiddah; M = Midyan; R = Ar Rayn. Boxes outline areas of Figs. 2, 9, and 11.

above the surrounding valley bottoms and is well dissected by east- and northeast-flowing drainages, which provide good exposures of the underlying geology. To the east, the northern and southern boundary faults of the ophiolite converge, and the ophiolite tapers and ceases to be recognizable. To the west, the ophiolite is cut by the northwest-trending sinistral fault system of the Da’bah and Durr shear zones. Mafic-ultramafic rocks continue to the south as the Sahluj mélange (named here after Jabal Sahluj) and the Jabal Wask ophiolite (Fig. 2). Together with the Jabal Ess ophiolite, these mafic-ultramafic rocks are

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Fig. 2. Simplified geologic map and geochronologic data for the Jabal Ess-Jabal Wask ophiolite zone, which marks the Yanbu suture in the northwestern Arabian shield. Mapping after Kemp (1981) and Hadley (1987). Geochronologic data after Kemp et al. (1980), Ledru and Augé (1984), Claesson et al. (1984), and Pallister et al. (1988). Box shows area of Fig. 3.

parts of the zone of deformed rocks that constitute the Yanbu suture in Saudi Arabia, and its extension in Northeast Africa, the Allaqi-Sol Hamid suture (Kröner et al., 1987; Abdelsalam and Stern, 1995; Johnson and Woldehaimanot, 2003). (For details on the Jabal

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Fig. 3. Simplified map and cross section of the Jabal Ess ophiolite (after Al-Shanti, 1982; Chevrèmont and Johan, 1982b; and this report). Geochronologic data from Claesson et al. (1984); Pallister et al. (1988).

Wask ophiolite, see Bakor et al., 1976; Chevrèmont and Johan, 1982a; Ledru and Augé, 1984.) Peridotite is mainly exposed on the southern slope of Jabal Ess in the central part of the ophiolite shown in Fig. 3. It is strongly altered and is chiefly black, massive serpentinite in which original textures are rarely preserved although sufficient primary features remain to locally indicate the presence of harzburgite, subordinate tec-

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tonized dunite, and cumulate wehrlite, orthopyroxenite, and serpentinite (Al-Shanti, 1982; Chevrèmont and Johan, 1982b). The harzburgite and dunite are concentrated in a zone of mantle peridotite as much as 500 m wide in outcrop. The harzburgite contains 5–20% bastite pseudomorphs after euhedral orthopyroxene in a serpentinized olivine ground mass (Shanti and Roobol, 1979). Dunite contains bastatized pyroxene ghosts and local disseminated chromite and podiform chromite lenses 20 cm across. Enstatite banding and trains of ovoid, stretched chromian spinel define a metamorphic foliation and lineation, which are suggestive of high-temperature subsolidus deformation possibly as a result of plastic mantle flow beneath a spreading ridge (Pallister et al., 1988). Olivine in the peridotite is magnesian rich with fosterite in the range Fo91 –Fo92.7, orthopyroxene is close to enstatite, and clinopyroxene is mainly diopside (Chevrèmont and Johan, 1982b; Shanti, 1983). Chromiferous spinel is similar to that from modern forearc peridotites (Stern et al., 2004) and is present as anhedral grains enclosing olivine and less commonly as euhedral grains interstitial to the cumulate olivine in the dunite or enveloped in orthopyroxene in harzburgite (Chevrèmont and Johan, 1982b). The cumulate rocks, inferred by Al-Shanti (1982) to be a unit about 400 m wide overlying the harzburgite, consists of serpentinized wehrlite and orthopyroxenite in layers a meter or so thick intercalated with serpentinite several meters thick. Gabbro is predominantly a dark, relatively featureless, massive rock, but locally has well-developed igneous lamination and rhythmic alternations of melanocratic gabbro and leucocratic anorthosite in layers as much as 20 cm thick (Fig. 4C). The gabbro is metamorphosed and, where strongly deformed, is mylonitized and brecciated, particularly in a narrow zone southeast of Jabal Ess where gabbro is exposed between peridotite and the dike complex and is tectonically intercalated with peridotite (Shanti and Roobol, 1979). The sheeted dike complex is a unit as much as 600 m wide composed of metadolerite dikes 30 cm to 2 m wide. Its contacts with gabbro, below, and pillow basalt, above, are transitional. Outcrop features of the complex are commonly obscured by extensive desert varnish but, where exposure is favorable, the complex is seen to consist entirely of dikes that have fine-grained chilled margins and fine- to medium-grained cores (Shanti and Roobol, 1979). The basalt unit includes pillow basalt (Fig. 4D), subordinate massive basalt flows as much as 10 m thick, and very sparse basalt breccia. It is estimated to be up to 300 m thick but because of folding and fault repetition is exposed over a width of nearly 2.5 km (AlShanti, 1982). Thin-skinned pillow lava characterized by amygdaloidal cores and rims of chloritized and (or) spherulitic basalt predominates. Khaki, locally siliceous shale and laminated chert crop out as interbeds 50 m thick in the pillow basalt and as isolated, strongly

Fig. 4. Features of the Jabal Ess ophiolite. (A) View of Jabal Ess from the south showing north-dipping shear surfaces and (along ridge line) a north-dipping sheet of carbonated and silicified peridotite. Relief about 150 m. (B) South-dipping thrusts in serpentinite mélange at the southern margin of the ophiolite. (C) Rhythmic layering in metagabbro showing anorthosite intercalated with melanocratic gabbro. (D) Pillow basalt. (Photos C and D after Al-Shanti, 1982.)

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sheared exposures at the northern boundary fault of the ophiolite (Shanti and Roobol, 1979), and are interpreted as pelagic sediments at the top of the ophiolite succession. The complex is steeply dipping, and the exposures are effectively a cross-section through the ophiolite. The gross distribution of rock types suggests an ophiolite succession younging from south to north but the succession is disrupted by deformation. Igneous layering in the peridotite dips 40◦ –90◦ , mostly to the south; the basalt is folded into a series of anticlines and synclines; and the gabbro is repeated by folding and/or thrusting north of the basalt. Shear zones, characterized by serpentinite schist, secondary listwaenite, magnesite, and mélange, are abundant and dip between 30◦ and 90◦ south and north. Mélange, composed of angular to subrounded blocks of massive serpentinite, gabbro, dolerite, and basalt up to 50 m in diameter in a yellow to black serpentinite schist matrix, is particularly conspicuous as a subvertical shear zone as much as 500 m wide in the southern part of the ophiolite. A north-dipping unit of chert and listwaenite (Fig. 4A) marks the shear zone at the northern boundary of the ophiolite on Jabal Ess. South-dipping shear zones are common in the northern unit of gabbro and in the southern part of the ophiolite (Fig. 4B), and a subvertical shear zone forms the southernmost boundary of the ophiolite. S/C shear fabrics are widespread. A dextral sense of horizontal shear predominates, but variation in the sense of shear along and across the ophiolite indicates that there was a large degree of strain partitioning in the ophiolite during deformation. The northern boundary fault, minor shear zones in the northern gabbro, and the shear zone at the southern margin of the ophiolite are dextral (Fig. 5A). A shear zone about 100–200 m south of the northern boundary fault about 1 km east of Jabal Ess summit is sinistral (Fig. 5B). The south dipping shear zones in mélange in the southern part of the ophiolite are both dextral and sinistral (Figs. 5C, D). Indicators of the sense of vertical movement on the shear zones have not been observed, but it is conceivable that south-dipping shears throughout the ophiolite are north-vergent thrusts. Worldwide, suture zones commonly display combinations of horizontal shearing and thrusting that reflect deformation during transpression. The structures at Jabal Ess, suggestive of north-vergent thrusting and regional dextral horizontal shear, are consistent with development under conditions of dextral transpression. The Jabal Ess ophiolite is directly dated by means of a 780 ± 11 Ma U-Pb zircon age obtained from gabbro in the eastern part of the ophiolite (Pallister et al., 1988) and a 782 ± 36 Ma Sm-Nd mineral model age obtained from the same gabbro sample (Claesson et al., 1984). A younger U-Pb zircon age of 706 ± 11 Ma obtained from trondhjemite that intrudes already serpentinized and sheared gabbro provides a minimum age for ophiolite formation (Pallister et al., 1988). Both U-Pb ages are model ages, obtained by forcing the lower intercept through a fixed point of 15 ± 15 Ma, a procedure commonly applied to U-Pb geochronologic data in the Arabian shield (Cooper et al., 1979).

3. THARWAH OPHIOLITE COMPLEX The Tharwah ophiolite complex (Nassief, 1981; Nassief et al., 1984; Pallister et al., 1988) consists of mafic-ultramafic rocks preserved as a stack of steeply dipping, northwest-

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Fig. 5. Typical shear fabrics in the Jabal Ess ophiolite. (A) Dextral shear fabric in shear zone at the northern margin of the ophiolite. Knife for scale, about 10 cm long. (B) Sinistral shear fabric in serpentinite mélange south of the northern margin of the ophiolite. Hammer for scale. (C and D) Dextral and sinistral shear fabrics, respectively, about 5 m apart along strike in serpentinite mélange close to the southern margin of the ophiolite. Pen for scale in (C), hammer for scale in (D).

and southeast-vergent thrust sheets exposed over an area 13 km east-west and 6 km northsouth (Fig. 6). Together with adjacent pelagic rocks, the ophiolite is part of the Labunah thrust zone (Ramsay, 1986) and lies in the zone of deformed rocks that constitutes the southwestern part of the Bi’r Umq suture (Pallister et al., 1988; Johnson et al., 2002). The ophiolite is exposed in hills rising 150–200 m above the adjacent Red Sea Coastal Plain. Weathering is locally intense and most rock surfaces are coated in desert varnish that, compounded by pervasive metamorphism and shearing, makes rock identification difficult. The succession is disrupted and locally inverted. Serpentinized depleted-mantle harzburgite and subordinate dunite together with minor lenses and dike-like bodies of lherzolite and gabbro make up the central part of the complex. Harzburgite contains relict olivine (Fo89.5–93.4) (70–90 mode%), bastite pseudomorphs of orthopyroxene (En90–92) (15–30%), chromite (< 1%), and clinopyroxenes (< 1%). Dunite is largely serpentinized olivine. Chemically, the chromites resemble chromitiferous spinels in modern-day forearcs (Stern et al., 2004). The rocks are tectonized and have a strong, high-temperature foliation

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Fig. 6. Simplified map and cross section of the Jabal Tharwah ophiolite (after Nassief, 1981; Nassief et al., 1984; Ramsay, 1986; Pallister et al., 1988; Johnson, 1998). Geochronologic data after Pallister et al. (1988).

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Fig. 7. Features of the Jabal Tharwah ophiolite. (A) Cumulate peridotite showing interlayering of dunite and pyroxenite (circle encloses bent-over person for scale). (B) Rhythmic layering in gabbro (pen for scale 12 cm long). (C) Sinistral shear fabric in the interior of the ophiolite (hammer for scale).

composed of orthopyroxene and chromite grains (Nassief et al., 1984). Cumulate ultramafic rocks crop out in the northern part of the complex as a unit of dunite, lherzolite, and pyroxenite intercalated in layers 1–10 m thick (Fig. 7A). The cumulate rocks are as much as 2.9 km thick, but are probably thickened by deformation from an original thickness of about 1 km (Nassief et al., 1984). Oriented pyroxene produces a weak igneous lamination, but a high-temperature deformational foliation of the type displayed by the peridotite is absent. Olivine is less magnesium rich than in the mantle peridotite (Fo86.5–89.5) and clinopyroxene is chiefly diopside. Orthopyroxene, mostly present as exsolution laminae in clinopyroxene, is largely replaced by bastite but is preserved locally as grains of En82–89 (Nassief et al., 1984). Layered to locally massive gabbro is present in the north and south of the complex. It is metamorphosed in the greenschist facies, but has well-developed igneous lamination, cm- to m-scale plagioclase- and pyroxene-rich phase laying (Fig. 7B), and, to a lesser de-

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gree, grain-size layering (Nassief et al., 1984). Sheeted dikes crop out in fault-bounded units 10 km long and over 300 m thick at the southern margin of the ophiolite. The dikes are not as well exposed as the dike complex in the Jabal Ess ophiolite and, because of shearing and alteration, their protoliths are not always evident. However, Nassief et al. (1984) report that, in places, dikes comprise 50–90% of the outcrop and are observed to be 1– 1.5 m wide, separated by screens of altered gabbro and basalt. Pillow basalt occurs in part of the dike complex. The rocks are strongly sheared and altered but bulbous pillow forms are still discernible (Nassief et al., 1984). Fine-grained argillaceous and cherty sedimentary rocks, carbonates, and basalt interpreted to be pelagic, ocean-floor deposits at the top of the ophiolite are faulted against the main mass of the Tharwah ophiolite south of Jabal Tharwah along the Thamrih fault and appear to be in depositional contact with the ophiolite northeast of Wadi Qirba’. Small lenses and veins of gabbroic pegmatite and leucodiorite or trondhjemite, probably representing late-fractionated derivatives of the gabbroic magma, intrude massive gabbro at the extreme northern edge of the complex along the Qirba’ fault. Layering in the cumulate unit is locally moderately inclined (Fig. 7A), but most structures in the Tahrwah ophiolite are steep. The Qirba’ fault dips 50◦ –80◦ to the southeast; the Thamrih fault and shear zones within the complex are subvertical. The pelagic rocks south of the complex make up the southwest plunging (40◦ –50◦) Farasan synform (Fig. 6). This has a steeply dipping northwest limb, a more gently inclined southeast limb, and is truncated by a southeast-vergent thrust along the southern flank of Jabal Farasan. The Ukaz fault at the southern boundary of the Labunah thrust zone is an oblique dextral, hangingwall-up-to-the south steep reverse fault that juxtaposes the pelagic rocks with the Samran group (Johnson, 1998). An exceptional gently inclined fault exposed at a location about 6 km northeast of Jabal Farasan may be a remnant of an original thrust dipping 35◦ –43◦ to the northwest. The structure of the Tharwah ophiolite and the Labunah thrust zone is believed to reflect two phases of progressive deformation (Johnson, 1998). The early phase caused northeast- and southwest-trending tight to isoclinal folding, the development of beddingparallel shear surfaces, and thrusting. Folding during the second phase created the Farasan synform, folded and steepened early thrusts and shear surfaces, and refolded early isoclinal folds and lineations. The rocks were pervasively affected by non-coaxial strain during both phases of deformation and S/C fabrics, asymmetrical extensional-shear bands, winged porphyroclasts, and quartz-mosaic ribbons are widespread (Johnson, 1998). The Qirba’ fault shows evidence of both sinistral and dextral horizontal as well as top-to-the-northwest reverse-slip movements; the Ukaz fault shows top-to-the-southeast and dextral horizontal movements; whereas shears interior to the ophiolite are commonly sinistral (Fig. 7C). Johnson (1998) proposes that the Tharwah ophiolite is a flower structure that developed in a zone of dextral transpression (see the cross section in Fig. 6) and, as in the case of the Jabal Ess ophiolite, the variations in sense of shear indicate considerable strain partitioning during its formation. Zircon grains from gabbro in the northern and southern parts of the ophiolite yield a near-concordant U-Pb age of 870 ± 11 Ma (Pallister et al., 1988). Other gabbro zircon

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fractions are highly discordant, yielding 207 Pb/206Pb model ages of about 1250 Ma. The 870 Ma result is not robust, but if provisionally accepted as a crystallization age, suggests ocean-floor magmatism 100 million years earlier than the Jabal Ess magmatism. The 1250 Ma age is likely to be an artifact caused by assimilation of xenocrystic zircons, similar to the explanation of anomalously old ages for some Bi’r Umq ophiolite zircon samples (Calvez et al., 1985).

4. BI’R UMQ OPHIOLITE COMPLEX The Bi’r Umq ophiolite complex consists of serpentinized and carbonate-altered peridotite, gabbro, and mélange (undivided in Fig. 8), and a 1500-m thick succession of spilitic metabasalt, chert, and metatuff assigned to the Sumayir formation (Al-Rehaili, 1980; Al-Rehaili and Warden, 1980; Le Metour et al., 1982; Kemp et al., 1982; Pallister et al., 1988). The complex crops out in an area of about 60 km by 20 km at the northeastern end of the Bi’r Umq suture (Johnson et al., 2002). The ultramafic rocks and gabbro are concentrated in the south close to the Bi’r Umq and Wobbe faults and are a disproportionately small component of the ophiolite in comparison with other ophiolites described here. The ultramafic rocks and gabbro form a chain of discontinuous hills that have moderate relief of 50–75 m and are partly held up by more resistant listwaenite and chert. The Sumayir formation crops out in low-lying exposures north of the Bi’r Umq fault and is extensively covered by colluvium and alluvium. Mélange is mostly in discontinuous exposures along the Bi’r Umq fault. The ophiolite is truncated by the Arj fault on the west (west of the area shown in Fig. 8), a sinistral strike-slip structure belonging to the Najd fault system. The Raku-Mandisa faults truncate the ophiolite on the east. The Raku fault is a dextral shear of uncertain origin; because of poor exposure little is known about the Mandisa fault other than its trace, identified by a narrow, linear zone of listwaenite. Peridotite at Bi’r Umq is interpreted by Le Metour et al. (1982) to be an ultramafic cumulate consisting of dunite and subordinate, locally cumulus harzburgite. The rocks were pervasively sheared during ophiolite emplacement (Le Metour et al., 1982) and are extensively serpentinized, carbonated, and silicified, which results in the common development of Cu- and Ni-rich listwaenite along shear zones. Olivine in the dunite is thoroughly replaced by serpentine and is only recognized as ghost pseudomorphs. Harzburgite contains cumulus serpentinized olivine and intercumulus bastite-altered orthopyroxene (Le Metour et al., 1982). Small intrusions of hypabyssal trondhjemite, plagiogranite (termed keratophyre by Pallister et al., 1988), diorite, hornblende gabbro, metadiabase, and basalt occur at, or in a separate thrust slice south of, the Bi’r Umq fault (Le Metour et al., 1982; Pallister et al., 1988). The Sumayir formation is predominantly a homogeneous, monotonous unit of fine-grained greenstone derived from basalt flows and tuffs and subordinate pillow basalt and basaltic breccia (Al-Rehaili and Warden, 1980). Diagnostic of their metamorphic grade, the rocks contain sodic plagioclase (An5–20), secondary green tremolite, chlorite, epidote, iron oxide, and carbonate, with local relict clinopyroxenes. Minor metasedimentary units in the greenstone consist of thin-bedded felsic tuff, limestone,

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Fig. 8. Simplified map and cross section of the Bi’r Umq ophiolite and adjacent areas. Mapping after Al-Rehaili (1980); Kemp et al. (1982); Le Metour et al. (1982); Pallister et al. (1988); Johnson et al. (2002). Geochronology after Dunlop et al. (1986); Pallister et al. (1988).

chert, and siltstone, locally altered to mafic and felsic schist and amphibolite. Mélange consists of blocks of serpentinite, spilitic basalt, dolerite, and gabbro a few to several hundred meters across in a sheared serpentinite matrix.

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Regionally, the Bi’r Umq ophiolite trends east-west, but is broadly folded about a northsouth axis and in detail trends west-southwest in the west and east-southeast in the east (Fig. 8). The Bi’r Umq fault at its southern margin is a steeply (50◦ –70◦ ) north-dipping reverse fault, and the Wobbe fault is a moderately (40◦ –50◦ ) southwest-dipping reverse fault. The northern Shuwaykah fault is believed to be a high-angle southeast-dipping reverse fault (Johnson et al., 2002). The dip across the Bi’r Umq complex changes from steeply northwest in the south to steeply southeast in the north and the complex is interpreted by Al-Rehaili (1980) to be a large asymmetric synform, although a flower structure of the type shown in the cross section in Fig. 8 is our preferred interpretation. Pillow structures indicate that basalt units are right way up (Al-Rehaili and Warden, 1980). Shear fabrics and the down-dip plunge of stretching lineations indicate an early phase of top-to the south reverse dip-slip movement on a south-vergent thrust along the Bi’r Umq fault; later movement included dextral and sinistral horizontal shear (B. Blasband, written communication, 2001). The kinematics of the Wobbe fault are unknown. The ophiolite is directly dated by a three-point U-Pb zircon model age of 838 ± 10 Ma obtained from diorite in the ophiolite close to the Bi’r Umq fault (Pallister et al., 1988). Trondhjemite and a pyroxene separate obtained from nearby gabbro yield a composite Sm-Nd isochron age of 828 ± 47 Ma (Dunlop et al., 1986), and Sumayir-formation basalt yields a three-point Rb-Sr whole-rock isochron of 831 ± 47 Ma (Dunlop et al., 1986). Single-point zircon model ages of 764 ± 3 Ma and 782 ± 5 Ma obtained from plagiogranite (or keratophyre) that cuts already serpentinized and carbonated peridotite and is interpreted to be a post-serpentinization and post-obduction intrusion constrain the minimum age of ophiolite emplacement (Pallister et al., 1988).

5. BI’R TULUHAH OPHIOLITE The Bi’r Tuluhah ophiolite crops out in the northern (Hulayfah) part of the HulayfahAd Dafinah fault zone in the north-central part of the Arabian shield (Fig. 9). The rocks are strongly folded and sheared and together with rocks of the Nuqrah formation constitute a subvertical brittle-ductile shear zone that resulted from sinistral transpression during suturing between the Afif and Hijaz terranes (Quick and Bosch, 1989; Johnson and Kattan, 2001). The suture continues as an ophiolite-decorated shear zone over 500 km to the south and southeast, and is one of the longer sutures recognized in the Arabian shield. Lithologic contacts in the fault zone are mostly faults so that original stratigraphic and structural relations are obscure, but an ophiolite is identified at Bi’r Tuluhah on the basis of the presence of amphibolite, serpentinized peridotite, layered gabbro, and noncumulus gabbro (Delfour, 1977). The ophiolite, the central part of which is shown in Fig. 10, is about 30 km long in a north-south direction and 6 km wide, and crops out in low-relief hills rising 10–40 m above the wadi plain (Kattan, 1983). The rocks are strongly weathered and intense coatings of desert varnish commonly cover outcrop surfaces. Volcanic and volcaniclastic rocks of the Hulayfah formation flank the fault zone on the west and epiclastic rocks and bimodal basalt and rhyolite of the Shammar group and Shammar

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Fig. 9. Simplified geologic map showing the location of the Bi’r Tuluhah, Darb Zubaydah, and Wadi Khadra ophiolitic complexes along and east of the Hulayfah-Ad Dafinah fault zone (HDFZ), part of the Hulayfah-Ad Dafinah-Ruwah suture. The fault zone is overlain and intruded by post-amalgamation basins and granites and displaced by Neoproterozoic III northwest-trending Najd faults. Map after Johnson and Kattan (2001). Box outlines area of Fig. 10.

5. Bi’r Tuluhah Ophiolite

145

Fig. 10. Simplified map of the Bi’r Tuluhah ophiolite and adjacent areas. Map after Kattan (1983); Le Metour et al. (1983); Quick and Bosch (1989); and Johnson et al. (1989). Geochronologic data after Calvez et al. (1984); Stuckless et al. (1984); Calvez and Kemp (1987); Pallister et al. (1988). HFZ = Hulayfah fault zone.

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granites overlie and intrude the fault zone on the east. Post-Shammar northwest-trending sinistral Najd faults dislocate the ophiolite and fault zone on the north and south. The most extensive unit in the ophiolite is peridotite, which crops out in a zone 2.5 km wide north and south of Bi’r Tuluhah. The rocks are strongly serpenitinized but harzburgite and dunite protoliths are recognized. The rocks are mylonitized and have a strong cataclastic texture although ghost equant grains and a relict granoblastic texture are recognized in thin sections (Kattan, 1983). Serpentinized harzburgite comprises about 15% orthopyroxene, 80% serpentinized olivine presudomorphs, and minor chromite and magnetite. In hand specimen, serpentinized dunite is fine grained and dark gray to green, and in thin section is an aggregate of serpentine minerals that locally have a well-developed boxwork texture derived from the original olivine (Kattan, 1983). It contains anhedral grains of chromite and magnetite, and small lenses of massive chromite. As shown by Stern et al. (2004), the chromites plot on the low-Mg side of the field of present-day forearc chromian spinels, suggesting a suprasubduction origin for the ophiolite. Layered gabbro and dunite with minor wehrlite, lherzolite, websterite, and olivine clinopyroxenite layered on scales of centimeters to meters occur in a narrow band west of the serpentinized peridotite (Le Metour et al., 1983). Serpentinization obscures primary textures, and it is not clear whether the layered rocks are mantle tectonites or ultramafic cumulates (Quick and Bosch, 1989). A narrow zone of mafic plutonic rocks farther west consists of gabbro and diorite in the north, in the northern half of Fig. 10, and fault slivers of layered gabbro in the south. The northern gabbro and diorite either intrude or are faulted against the ophiolitic rocks, and may postdate ophiolite magmatism (Le Metour et al., 1983). The southern layered gabbro has a cumulus texture in pyroxene- and hornblenderich phases and may be a cumulate part of the ophiolite succession. Amphibolite, treated by Delfour (1977) as part of the ophiolite, crops out east of the peridotite. Fine-grained amphibolite is strongly schistose and lacks clear textural or relict mineralogic indications of its protoliths. Coarse-grained amphibolite appears to be the result of epidote-amphibolite facies metamorphism of gabbro and diabase (Quick and Bosch, 1989). Massive, locally pillowed metabasalt and chert together with fine-grained sandstone, keratophyre, and interbedded felsic tuff and minor basalt make up the volcanicvolcaniclastic rocks of the Nuqrah formation located on either side of the mafic-ultramafic units along the fault zone (Le Metour et al., 1983; Quick and Bosch, 1989). Considered in isolation, it is conceivable that the basalt and chert represent pelagic rocks at the top of the ophiolite succession but their association with felsic tuffs and sandstone suggest that they are part of a suprasubduction volcanic arc. The metabasalt is a fine-grained, light gray to gray-green rock, the original structure and texture of which are virtually obliterated by metamorphism. The rock is identified as basalt in the field by its mafic composition, generally massive appearance, and the local presence of pillow structure. In thin section, the basalt has a strongly developed, fine-grained metamorphic foliation composed of saussuritized plagioclase, epidote, carbonate, chlorite, clinozoisite, and iron oxides (Kattan, 1983). Dikes of diabase, gabbro, plagiogranite, and diorite cut all the serpentinized ultramafic rocks, and plutons of diorite and quartz diorite intrude the southern part of the ophiolite.

6. Halaban Ophiolite

147

Structurally, the Bi’r Tuluhah ophiolite is a set of fault-bounded lenses. Together with the flanking Nuqrah formation, the rocks are pervasively cleaved and sheared, and deformation is spread across the entire width of the Hulayfah fault zone, although narrow zones of ultramylonite, schist, and carbonate-altered serpentinite identify discrete shears within the fault zone. All shear surfaces are subvertical, and any low-angle thrusts that may have been originally present have been obliterated or steepened by subsequent deformation. Model U-Pb zircon ages of 847 ± 14 Ma and 823 ± 11 Ma obtained from plagiogranite (trondhjemite) dikes that intrude serpentinized harzburgite in the center of the ophiolite (Pallister et al., 1988) provide a minimum age for the ophiolite. The dikes have rodingite margins indicating intrusion prior to complete serpentinization and they are interpreted by Pallister and colleagues as forming late during ophiolite magmatism. Together with a UPb zircon age of 839 ± 23 Ma obtained from Nuqrah formation rhyolite 40 km east of the Bi’r Tuluhah ophiolite (Calvez et al., 1984), the Bi’r Tuluhah model ages suggest that oceanic crust and volcanic arc rocks were actively forming in the region between 840 and 820 Ma. The age of the Hulayfah formation is weakly constrained by a U-Pb zircon age of 720 ± 10 Ma obtained from tonalite that intrudes the formation west of the Hulayfah fault zone (Calvez et al., 1984). An approximate U-Pb zircon age of 710 Ma obtained from a quartz diorite pluton intruded into the southern part of the ophiolite and fault zone (Fig. 10) constrains the minimum age for ophiolite deformation and alteration (J.S. Stacey, personal communication, cited by Quick, 1991) and approximate U-Pb and Rb-Sr ages between 630 Ma and 615 Ma obtained from the post-amalgamation Shammar group and Shammar “stitching” granites (Stuckless et al., 1984; Calvez and Kemp, 1987) give a minimum age for completion of suturing along the fault zone.

6. HALABAN OPHIOLITE The Halaban ophiolite is a zone of mafic-ultramafic rocks exposed north and south of Halaban in the eastern part of the Arabian shield (Figs. 11, 12) and located along the Halaban suture at the join between the Afif and Ad Dawadimi terranes. The Ad Dawadimi terrane is strongly deformed and treated by some authors as, itself, part of a larger suture zone referred to as the Al Amar suture (Stoeser and Camp, 1985). The ophiolite consists of metagabbro and subordinate serpentinite. It is bounded on the west by the HalabanZarghat fault zone, a complex structure including a west-vergent thrust in the south and a down-to-the-west normal fault in the north, and on the east by the Eastern shear zone (Fig. 12). The rocks west of the ophiolite include mafic plutons, orthogneiss, and amphibolite referred to as the Suwaj domain and late Neoproterozoic sedimentary rocks of the Jibalah group deposited in the Antaq basin. Rocks to the east are low-grade metasedimentary units of the Abt formation and post-amalgamation granitoids. The ophiolitic rocks crop out in hills with relief of about 50 m and are generally well exposed. Unfortunately, the margins of the ophiolite tend to coincide with valleys so that structural details at the boundaries of the ophiolite are largely obscured. From the point of view of the Penrose definition, the rocks in the vicinity of Halaban do not include mantle peridotite, a dike

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Fig. 11. Simplified geologic map and geochronologic data for the Halaban, Jabal Tays, and Jabal al Uwayjah ophiolite complexes in the eastern Arabian shield. Map after Delfour (1979); Manivit et al. (1985); and this report. Geochronologic data after Calvez et al. (1984); Stacey et al. (1984). Abbreviations: AFZ = Al Amar fault zone; EF = East fault (magnetically inferred); HYFZ = Hufayrah fault zone; HZF = Halaban-Zarghat fault zone; ARFZ = Ar Rika fault zone; WF = West fault (magnetically inferred).

complex, or pillow basalt and, at best, are an incomplete ophiolite. However, north of the area described here, the on-strike continuation of the Halaban rocks includes peridotite, gabbro, serpentinite, listwaenite, and basalt (Al-Shanti and El-Mahdy, 1988). It is possi-

6. Halaban Ophiolite

149

Fig. 12. Simplified map and cross section of the Halaban ophiolite complex and adjacent units, showing the locations and results of geochronologic dating. Map after Delfour (1979), Al-Saleh (1993), and this report; geochronologic data after Stacey et al. (1984), Al-Saleh (1993), and Al-Saleh et al. (1998).

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ble that the rocks in the vicinity of Halaban village are the mafic plutonic section of an ophiolite, whereas the mantle part is preserved in the north. The Halaban rocks are predominantly pale green, well-foliated metagabbro (Al-Saleh et al., 1998). Al-Shanti and El-Mahdy (1988) interpret the foliation as igneous layering and describe microscale igneous lamination and large-scale rhythmic layering caused by differences in mineralogy, grain size, and texture. Some of the foliation, however, is clearly secondary in origin, with oriented metamorphic minerals and quartz ribbons, ductile folding, as well as S/C shear fabrics and Al-Saleh et al. (1998) interpret much of the foliation to be the result of sea-floor metamorphism under greenschist- to amphibolite-facies conditions. The widespread development of actinolite, chlorite, clinozoisite, and albite is inferred to reflect ubiquitous, low-grade, low-temperature, off-axis metamorphism, whereas local amphibolitization of gabbro is inferred to reflect metamorphism close to the spreading axis in conjunction with shearing. Primary feldspar in the gabbro is commonly saussuritized and clinopyroxene tends to be replaced by chlorite and quartz. Lenses of massive black serpentinite and serpentinized lherzolite and olivine websterite occur sporadically along the western margin and axis of the gabbro (Fig. 12). The serpentinite lenses have sharp contacts with surrounding gabbro and are interpreted as ultramafic diapirs emplaced in the gabbro from an originally lower stratigraphic position in the ophiolite (Al-Saleh et al., 1998). Chromite from a small pod south of Halaban village plots close to the fields of chromian spinels from boninite and forearc ophiolites (Stern et al., 2004). South of Halaban village the gabbro outcrops taper and are structurally underlain by metamorphosed mafic and ultramafic rocks belonging to a sub-ophiolitic metamorphic complex (Al-Saleh et al., 1998). The eastern part of the metamorphic complex contains abundant blocks of serpentinite, 1–20 m across, sheathed by soapstone and small lenses of chromite in a matrix of orthoamphibolite and rodingite. It forms an inhomogeneous unit that may represent an obduction-related mélange. To the west, the metamorphic complex becomes more felsic. It contains no ophiolitic material and was probably largely derived from diorite and tonalite belonging to the Suwaj magmatic arc. The metamorphosed rocks were affected, particularly at their contact with the Halaban gabbro, by partial melting, which resulted in the development of migmatitic gneiss composed of coarse-grained hornblendite, amphibolite and gneissic gabbro and diorite intruded by numerous veins and irregular lenses of trondhjemite. Petrologic studies indicate that the mafic paleosome of the gneiss was partially melted under hydrous conditions; the neosome segregations are chiefly quartz and andesine plagioclase (Al-Saleh et al., 1998). The Eastern shear zone consists of strongly deformed gabbro, talc schist mélange, and pelitic schist exposed in a zone as much as 1 km wide. The rocks are tectonically intercalated with each other or are present as a mélange comprising irregular, scattered blocks of the various rock types in an anthophyllite-talc schist matrix (Fig. 13A). Gabbro in the Eastern shear zone commonly has a mylonitic texture and is cut by shear zones several centimeters thick that contain S/C fabrics. Altered basalt consists of pumpellyite pseudomorphs of original plagioclase phenocrysts, chlorite, quartz, and hematite, and the pelitic schists are rich in Ca and Mg silicates and believed to be derived from deep-marine argillaceous carbonates (Al-Saleh, 1993).

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151

Fig. 13. Features of the Halaban ophiolite. (A) Mélange from the Eastern shear zone. (B) Brittleductile shear in the partial melt zone showing top-up-to-the west displacement.

The steep dips of schistosity and shear surfaces indicate that the Eastern shear zone is subvertical. The western boundary of the ophiolite west of Halaban, in contrast, is inferred to be east dipping in conformity with east-dipping foliations and shear surfaces in the gabbro in proximity to the boundary. Sense-of-shear indicators in the partial melt zone beneath the gabbro (Fig. 13B) suggest that this western boundary was affected by west-

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directed shearing and the boundary is interpreted by us as a west-vergent thrust that placed the Halaban ophiolite above the Suwaj domain. Radiometric age determinations are reported from many units in the area. A U-Pb model zircon age of 694 ± 8 Ma obtained from a hypersthene gabbro in the southern part of the ophiolite about 10 km NNW of Halaban (Stacey et al., 1984) constrains the magmatic age of the ophiolite. Several 40 Ar/39Ar ages of about 680 Ma obtained from amphibolites of the sub-ophiolitic metamorphic complex and from metamorphic hornblendes in the ophiolite gabbro are interpreted to reflect rapid cooling and obduction of the ophiolite (Al-Saleh et al., 1998). U-Pb and 40 Ar/39Ar dates obtained from Suwaj diorite suggest that the Suwaj domain developed between 681 Ma and 675 Ma (Stacey et al., 1984; Al-Saleh et al., 1998).

7. JABAL TAYS OPHIOLITE The Jabal Tays ophiolite crops out in the central part of the Ad Dawadimi terrane, 75 km east-southeast of the Halaban ophiolite. The exposures form a group of prominent hills that have a local relief of 220 m rising to a summit of 1057 m above sea level at Jabal Tays, and are surrounded by low-relief exposures of low-grade sandstone, siltstone, conglomerate, and limestone of the Abt formation (Fig. 14). Isolated bodies of gabbro and mafic dikes exposed south of the area shown in Fig. 14 may be detached parts of the ophiolite (AlShanti and Gass, 1983), but their exact relation to the Jabal Tays exposures are not clear at this stage because of surficial cover. Mafic-ultramafic rocks at Jabal Tays include a large amount of undifferentiated serpentinite, subordinate amounts of gabbro intruded by mafic dikes, mélange, serpentinite schist, and listwaenite. Gabbro is variably serpentinized but is fresh enough that igneous lamination and cyclic layering of melanocratic, olivine- and pyroxene-rich gabbro and anorthosite are recognized (Al-Shanti and Gass, 1983). Plagioclase and clinopyroxene in the gabbro have a cumulate texture; orthopyroxene is mostly replaced by chlorite. The serpentinite, which makes up the bulk of Jabal Tays, is variably sheared and typically consists of relatively massive serpentinite cut by shear zones marked by serpentinite schist. The serpentinite protoliths have not been identified but are presumably varieties of mafic and ultramafic rocks. Mélange is uniformly developed at the outer margins of the ophiolite as a zone up to 500 m wide. It comprises irregular blocks of gabbro and massive to schistose serpentinite from a few centimeters to tens of meters across in a serpentinite and talc-schist matrix. Along the western side of Jabal Tays, the mélange creates a distinctive rugged terrain in which the mélange clasts weather out as protuberances (Fig. 15A). Carbonate alteration is widespread in the area, but is particularly prominent in west-dipping shear zones on the southern flank of Jabal Tays, on which basis the mountain is interpreted to be a stack of west-dipping thrusts. The external contacts of the ophiolite are poorly and discontinuously exposed, but the manner in which the outer contact and mélange zone wraps around Jabal Tays suggests that the ophiolite is a synform (Fig. 14). Where exposed, the outer, structurally lower contact is a shear zone 1–5 m thick discordant with respect to the underlying Abt formation.

7. Jabal Tays Ophiolite

153

Fig. 14. Geologic map and cross section of the Jabal Tays ophiolite complex. Map after Al-Shanti and Gass (1983) and this report. Geochronologic data after Al-Shanti et al. (1984).

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Fig. 15. Features of the Jabal Tays and Jabal Uwayjah ophiolites. (A) View to the southwest of the mélange zone on the western side of the Jabal Tays ophiolite. (B) Dextral shear in the footwall of the Jabal Tays basal thrust. (C) Strongly developed deformational (?) foliation in Jabal Uwayjah gabbro. (D) West-dipping shear surfaces in Jabal Uwayjah serpentinized peridotite. (E) Shear fabric in serpentinized peridotite from locality (D) looking south, showing top-to-left, that is hanging-wall-up-to-east, displacement.

8. Jabal al Uwayjah Ophiolite

155

Mesoscale folding in the basal shear zone is indicated by changes in dip of shear surfaces from flat lying (25◦ –35◦ ) to subvertical over distances of a few meters. S/C shear fabrics indicate dextral-horizontal (Fig. 15B) and hanging-wall up-to-the-east vertical movements along the eastern and southeastern parts of the shear zone, suggestive of the ophiolite being part of an easterly vergent nappe. This inferred transport direction is compatible with an E-W elongated chromite lineation described from deformed gabbro in the central part of the ophiolite (Al-Shanti and Gass, 1983) and with the southwesterly plunge of stretched pebbles observed by us in Abt formation conglomerate in the footwall of the ophiolite on the southeastern flank of Jabal Tays. Whether the mélange along the western margin of the ophiolite is part of the basal thrust upturned by synclinal folding or a secondary mélange created along the north-trending steep fault that appears to truncate the ophiolite on the west is not yet established. The mafic-ultramafic rocks of the Jabal Tays ophiolite are not directly dated. Their minimum age is weakly constrained by an Rb-Sr whole-rock isochron of 620 ± 40 Ma obtained from trondhjemite that intrudes the mélange zone (Al-Shanti et al., 1984). However, granitoids elsewhere in the Ad Dawadimi terrane are dated 670–640 Ma (Stacey et al., 1984), and the Rb-Sr age is too young to be a meaningful constraint on the ophiolite. By comparison with the Halaban ophiolite, the Jabal Tays body is probably more likely to be about 680 Ma.

8. JABAL AL UWAYJAH OPHIOLITE The Jabal al Uwayjah ophiolite is exposed at the eastern edge of the Arabian shield in a group of isolated hills of low relief (< 80 m). Because of extensive cover by Quaternary eolian sand, pediment gravel, and wadi alluvium (Fig. 16), exposure is poor, and this characteristic in combination with little recent mapping means that the ophiolite is the least well known in the shield. Overall, it appears to occupy an area of about 45 km N-S and 12 km E-W on the shield but, judging by its aeromagnetic signature, continues 40 km to the southeast beneath the Permian, making it one of the larger ophiolites in the region. The contacts of the ophiolite are concealed and its structure is obscure, but prominent magnetic lineaments suggest major faults occur along the axis and western margin of the ophiolite. Permian sandstone and limestone are unconformable on the ophiolite on the east and amphibolite-grade metadiorite, metagabbro, and amphibolite, and garnet-amphibole gneiss of the Ghadaniyah complex (701 ± 5 Ma; Agar et al., 1992) flank the ophiolite on the west. The Ghadaniyah complex is lithologically and geochronologically similar to the Suwaj domain rocks in the Halaban area, and the two are correlated by Johnson (1996), as is implied by use of the same graphic symbol for the two rock units in Fig. 11. The most extensive exposures of the ophiolite are in low hills north and south of Jabal al Uwayjah and include serpentinized peridotite, pyroxenite, metagabbro, undifferentiated serpentinite, and minor metabasalt and metaandesite (Manivit et al., 1985). Pyroxenite has relict orthopyroxene (enstatite) and clinopyroxene (augite and diallage), and metagabbro, which is fine grained, strongly foliated (Fig. 15C), and closely resembles the Halaban

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Fig. 16. Map of the Jabal al Uwayjah ophiolite and adjacent units. Abbreviations: EF = Eastern fault; HYFZ = Hufayrah fault zone; WF = Western fault. Map after Bois and Shanti (1970), Brosset (1974), Manivit et al. (1985), and this report.

gabbro, contains relict olivine and pyroxenes pervasively altered to epidote and serpentine. Serpentinite is a black to green rock composed of antigorite, talc, and relict pyroxene. A fine-grained brick-red colloidal and ferruginous siliceous unit located at the contact be-

9. Summary and Discussion

157

tween the ophiolite and Permian rocks probably represents silicification of serpentinite as a result of weathering. A low-lying area west of the Jabal al Uwayjah hills is virtually devoid of exposures other than discontinuous north-trending ridges of carbonate. Some of the ridges are fine-grained gray or variegated white and gray, massive to thinly layered marble that resembles sedimentary marbles in other parts of the shield. Others, however, are listwaenite, which suggests that, despite the lack of outcrop, bedrock includes a significant amount of ultramafic rock. Structures in the Jabal al Uwayjah ophiolite are predominantly north trending. They include cleavage in serpentinite and gabbro, foliation in metagabbro, linear outcrops of listwaenite and carbonate, and magnetically inferred faults, two of which dominate the region. The western fault separates the ophiolite from the Ghadaniyah complex. The eastern fault, which locally coincides with ridges of listwaenite, separates the low hills north and south of Jabal al Uwayjah from the area of poor exposure to the west. West-dipping shear zones in serpentinite at the western edge of these hills contain S/C fabrics indicating top-to-the-east reverse slip (Figs. 15D, E), which suggests that the eastern fault may be an east-vergent thrust. The Jabal al Uwayjah ophiolite is not directly dated but on the basis of lithology is correlated by Brosset (1974) with the Halaban ophiolite. We concur with this correlation, which is consistent with the correlation mentioned above between the Ghadaniyah complex and Suwaj domain on the western flanks of the Jabal al Uwayjah and Halaban ophiolites, respectively, and provisionally infer that the Jabal al Uwayjah ophiolite is about 680 Ma.

9. SUMMARY AND DISCUSSION Ophiolites are widespread in the Arabian shield. However, as is evident from this review, they are ubiquitously deformed, with the consequence that typical ophiolite successions are not preserved at every occurrence. Nevertheless, sufficient diagnostic lithologic criteria are present to confidently conclude that numbers of the mafic-ultramafic complexes of the shield are indeed ophiolites. Of these, Jabal Ess is one of the most complete (Table 1); Jabal Tays the least complete. Jabal al Uwayjah is the least well known. Available geochronologic data indicate that the ophiolites developed over a 200-million year period delimited by Jabal Tharwah (∼ 870 Ma) and Halaban (∼ 695 Ma). Jabal Tharwah and Halaban are, in fact, the oldest and youngest ophiolites known in the entire Arabian-Nubian shield (Stern et al., 2004). The western and eastern geographic locations of the Tharwah and Halaban ophiolites conceivably suggest an eastward migration of oceanic floor magmatism in this period in the Arabian shield. However, in the larger setting of the entire region of juvenile Neoproterozoic rocks represented by the Arabian and Nubian shields, there is no unidirectional time-space distribution. Jabal Tharwah is in the middle of the combined Arabian-Nubian shield and younger ophiolites occur to the east, west, north, and south. Although not a specific topic of this review, it is also evident, in conjunction with a range of structural and stratigraphic information about the adjacent rocks, that ophiolites

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Table 1. Summary table showing lithologic components and estimated magmatic ages of selected Arabian shield ophiolites Ophiolite Jabal Ess Jabal Tharwah Bi’r Umq Bi’r Tuluhah Halaban Jabal Tays Jabal al Uwayjah √

Age (to nearest 5 Ma) 780 870 840–830 845–825 695

Peridotite

Gabbro

Dikes

Basalt

√ √ √ √ √ minor √ serpentinite √

√ √

√ √

? √ √ √ √

? ? ? ? ?

√ √ √ √ √ √

in the north

Pelagic sediments √ √ √ ? ? ? ?

= lithology observed; ? = lithology not observed to date.

in the Arabian shield occur along suture zones. The structure and geochronology of the ophiolites are therefore important constraints on the history of suturing. Dating of the Jabal Tharwah-Bi’r Umq ophiolites and associated intrusions imply ocean-floor magmatism in the northwestern part of the shield between ∼ 870–830 Ma and is consistent with development of the Bi’r Umq suture at about 780–760 Ma (Johnson et al., 2002). The 706 Ma age of post-ophiolite trondhjemite at Jabal Ess is consistent with convergence along the Yanbu suture and its Northeast African extension between ∼ 700 Ma and ∼ 600 Ma. The Bi’r Tuluhah ophiolite is the oldest known example of oceanic-floor magmatism along the Hulayfah-Ad Dafinah-Ruwah suture (Johnson and Kattan, 2001), and the Halaban ophiolite constrains oceanic magmatism and suturing in the eastern shield at ∼ 695 Ma and ∼ 680 Ma, respectively. Structurally, all the ophiolites are complex, and exhibit multiple phases of folding and shearing. Most structures are steep, and unlike some of the Neoproterozoic ophiolites in the Nubian shield (Abdelsalam and Stern, 1993; Schandelmeier et al., 1994), the Arabian examples have few preserved low-angle thrusts. The only candidates for original thrusts identified to date are low- to moderately inclined shears in the southern part of Jabal Ess, in eastern Jabal Tharwah, at the western contact of Halaban, along parts of the Jabal Tays basalt contact, and at the Eastern fault at Jabal al Uwayjah. Other shear zones are subvertical, either because they are folded thrusts, similar to the folding evident in the basal thrust at Jabal Tays, or are steep shear zones that developed during other phases of deformation. Overall, the available structural evidence is permissive of modeling the Jabal Ess ophiolite as a stack of steepened north-vergent thrusts and horizontal shear zones that resulted from a period of dextral transpression. The Jabal Tharwah and Bi’r Umq ophiolites are possibly both flower structures related to southeast- and northwest-vergent thrusting along the Bi’r Umq-Nakasib suture (Johnson et al., 2002). The Bi’r Tuluhah ophiolite is preserved in a subvertical shear zone that forms the northern part of the Hulayfah-Ad Dafinah-Ruwah shear zone created during sinistral transpression. The Halaban ophiolite is part of a westvergent allochthon thrust over the eastern margin of the Afif terrane at the Halaban suture. The Jabal al Uwayjah ophiolite is an extension of the Halaban ophiolite, detached from the Halaban rocks by the sinistral and top-to-the-north Hufayrah fault zone (Fig. 11), and its

References

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location is evidence that the Halaban suture continues to the edge of the shield and beyond, beneath Permian rocks that flank the shield. Available information about the Arabian shield ophiolites varies in quality and quantity. Unfortunately, none are known in sufficient detail to fully determine their tectonic setting. Additional, and in some cases, original, petrologic, geochemical, geochronologic, and structural research are required. The ophiolites have a small surface area, but are critical for our understanding of the tectonic history of the shield, and warrant ongoing study and exploration. They testify to the juvenile tectonic environment of the Arabian-Nubian shield; they document the creation of oceanic floor following the breakup of Rodinia; and in their deformed and metamorphosed state they record stages in the subduction and closure of the Mozambique Ocean concurrent with the amalgamation and suturing of the tectonostratigraphic terranes that make up the shield.

REFERENCES Abdelsalam, M.G., Stern, R.J., 1993. Structure of the late Proterozoic Nakasib suture, Sudan. Journal of the Geological Society of London 150, 1065–1074. Abdelsalam, M.G., Stern, R.J., 1995. Sutures and shear zones in the Arabian-Nubian Shield. Journal of African Earth Sciences 23, 289–310. Agar, R.A., Stacey, J.S., Whitehouse, M.J., 1992. Evolution of the southern Afif terrane—a geochronologic study. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report DGMR-OF-10-15, p. 41. Al-Rehaili, M.H., 1980. Geology of the mafic-ultramafic complex of Bi’r Umq area. M.Sc. thesis. King Abdulaziz University, Jiddah, p. 160. Al-Rehaili, M.H., Warden, A.J., 1980. Comparison of the Bi’r Umq and Hamdah ultrabasic complexes, Saudi Arabia. Institute of Applied Geology Bulletin 3 (4), 143–156. Al-Saleh, A.M., 1993. Origin, age and metamorphism of the Halaban ophiolite and associated units: implications for the tectonic evolution of the eastern Arabian Shield. Ph.D. thesis. University of Liverpool, p. 274. Al-Saleh, A.M., Boyle, A.P., Mussett, A.E., 1998. Metamorphism and 40 Ar/39 Ar dating of the Halaban ophiolite and associated units: evidence for two-stage orogenesis in the eastern Arabian shield. Journal of the Geological Society of London 155, 165–175. Al-Shanti, A.M., El-Mahdy, O.R., 1988. Geological studies and assessment of chromite occurrences in Saudi Arabia. King Abdulaziz City for Science and Technology Project No. AT-6-094 Final Report, p. 165. Al-Shanti, A.M., Gass, I.G., 1983. The Upper Proterozoic ophiolite mélange zones on the easternmost Arabian shield. Journal of the Geological Society of London 140, 867–876. Al-Shanti, A.M.S., Abdel-Monem, A.A., Marzouki, F.H., 1984. Geochemistry, petrology and Rb-Sr dating of trondhjemite and granophyre associated with Jabal Tays ophiolite, Idsas area, Saudi Arabia. Precambrian Research 24, 321–334. Al-Shanti, A.M.S., Mitchell, A.H.G., 1976. Late Precambrian subduction and collision in the Al Amar-Idsas region, Arabian Shield, Kingdom of Saudi Arabia. Tectonophysics 30, T41–T47. Al-Shanti, M.M.S., 1982. Geology and mineralization of the Ash Shizm-Jabal Ess area. Ph.D. thesis. King Abdulaziz University, Jiddah, p. 291.

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Bakor, A.R., Gass, I.G., Neary, C.R., 1976. Jabal al Wask, northwest Saudi Arabia: an Eocambrian back-arc ophiolite. Earth and Planetary Earth Sciences Letter 30, 1–9. Bois, J., Shanti, M., 1970. Mineral resources and geology of the As Sakhin quadrangle, photomosaic sheet 130. Bureau de Recherches et Géologiques et Minières Technical Record 70-JED-6, scale 1:100,000. Brosset, R., 1974. Geology and mineral exploration of the Umm Sulaym quadrangle, 22/45C. Bureau de Recherches Géologiques et Minières Technical Record 74-JED-9, scale 1:100,000. Calvez, J.-Y., Kemp, J., 1987. Rb-Sr geochronology of the Shammar group in the Hulayfah area, northern Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-07-11, p. 22. Calvez, J.-Y., Alsac, C., Delfour, J., Kemp, J., Pellaton, C., 1984. Geochronological evolution of western, central, and eastern parts of the northern Precambrian shield, Kingdom of Saudi Arabia. Faculty of Earth Sciences, King Abdulaziz University, Jiddah, Bulletin 6, 24–48. Calvez, J.-Y., Delfour, J., Kemp, J., Elsass, P., 1985. Pre Pan-African inherited zircons from the northern Arabian shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-05-13, p. 22. Chevrèmont, P., Johan, Z., 1982a. The Al Ays ophiolite complex. Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-02-5, p. 65. Chevrèmont, P., Johan, Z., 1982b. Wadi al Hwanet-Jabal Iss ophiolite complex. Deputy Ministry for Mineral Resources Open-file Report BRGM-OF-02-14, p. 30. Church, W.R., 1988. Ophiolites, structures, and micro-plates of the Arabian-Nubian shield: a critical comment. In: El-Gaby, S., Greiling, R.O. (Eds.), The Pan-African Belt of Northeast Africa and Adjacent Areas. Veiweg, Braunschweig/Wiesbaden, pp. 289–316. Church, W.R., 1991. Discussion of ophiolites in northeast and east Africa: implications for Proterozoic crustal growth. Journal of the Geological Society of London 148, 600–601. Claesson, S., Pallister, J.S., Tatsumoto, M., 1984. Samarium-neodymium data on two late Proterozoic ophiolites of Saudi Arabia and implications for crustal and mantle evolution. Contribution to Mineralogy and Petrology 85, 244–252. Cooper, J.A., Stacey, J.S., Stoeser, D.B., Fleck, R.J., 1979. An evaluation of the zircon method of isotopic dating in the southern Arabian craton. Contributions to Mineralogy and Petrology 68, 429–439. Delfour, J., 1977. Geology of the Nuqrah quadrangle, 25E, Kingdom of Saudi Arabia. Saudi Arabian Directorate General of Mineral Resources Geologic Map GM 28, 1:250,000 scale. Delfour, J., 1979. Geologic map of the Halaban quadrangle, sheet 23G, Kingdom of Saudi Arabia. Saudi Arabian Directorate General of mineral Resources Geologic Map GM-46, scale 1:250,000. Dunlop, H.M., Kemp, P., Calvez, J.-Y., 1986. Geochronology and isotope geochemistry of the Bi’r Umq mafic-ultramafic complex and Arj group volcanic rocks, Mahd adh Dhahab quadrangle, central Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-07-7, p. 38. Genna, A., Nehilg, P., Le Goff, E., Guerrot, C., Shanti, M., 2002. Proterozoic tectonism of the Arabian Shield. Precambrian Research 117, 21–40. Hadley, D.G., 1987. Geologic map of the Sahl Al Matran quadrangle, sheet 26C, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM-86, scale 1:250,000. Johnson, P.R., 1996. Geochronologic and isotopic data for rocks in the east-central part of the Arabian shield: stratigraphic and tectonic implications. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-96-3, p. 47.

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Johnson, P.R., 1998. The structural geology of the Samran-Shayban area, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Technical Report USGS-TR-98-2, p. 45. Johnson, P.R., Kattan, F., 2001. Oblique sinistral transpression in the Arabian shield: the timing and kinematics of a Neoproterozoic suture zone. Precambrian Research 107, 117–138. Johnson, P.R., Woldehaimanot, B., 2003. Development of the Arabian-Nubian Shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. Geological Society of London Special Publication 206, 289–325. Johnson, P.R., Abdelsalam, M., Stern, R.J., 2002. The Bi’r Umq-Nakasib shear zone: Geology and structure of a Neoproterozoic suture in the northeastern East African Orogen, Saudi Arabia and Sudan. Saudi Geological Survey Technical Report SGS-TR-2002-1, p. 33. Johnson, P.R., Quick, J.E., Kamilli, R.J., 1989. Geology and mineral resources of the Bi’r Tuluhah quadrangle, Kingdom of Saudi Arabia. Saudi Arabian Directorate General of Mineral Resources Technical Record USGS-TR-09-1, p. 42. Kattan, F.H., 1983. Petrology and geochemistry of the Tuluhah belt, northeast Arabian shield. M.S. thesis. King Abdulaziz University, Jiddah, p. 111. Kemp, J., 1981. Geologic map of the Wadi al Ays quadrangle, sheet 25C, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geologic Map GM 53, scale 1:250,000. Kemp, J., Gros, Y., Prian, J.-P., 1982. Geologic map of the Mahd adh Dhahab quadrangle, sheet 23E, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geologic Map GM 64, scale 1:250,000. Kemp, J., Pellaton, C., Calvez, J.-Y., 1989. Geochronological investigations and geologic history of the Precambrian of northwestern Saudi Arabia. Saudi Arabian Directorate General of Mineral Resources Open-File Report BRGM-OF-01-1, p. 120. Kröner, A., Greiling, R., Resichmann, T., Hussein, I.M., Stern, R.J., Dürr, S., Krüger, J., Zimmer, M., 1987. Pan-African crustal evolution in the Nubian segment of Northeast Africa. In: Kröner, A. (Ed.), Proterozoic Lithospheric Evolution. In: Geodynamic Series, vol. 17. American Geophysical Union, pp. 235–257. Ledru, P., Augé, T., 1984. The Al Ays ophiolitic complex; petrology and structural evolution. Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-04-15, p. 57. Le Metour, J., Johan, V., Tegyey, M., 1982. Relationships between ultramafic-mafic complexes and volcanosedimentary rocks in the Precambrian Arabian Shield. Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-12-15, p. 90. Le Metour, J., Johan, V., Tegyey, M., 1983. Geology of the ultramafic-mafic complexes in the Bi’r Tuluhah and Jabal Malhijah areas. Deputy Ministry for Mineral Resources Open-Field Report BRGM-OF-03-40, p. 47. Manivit, J., Pellaton, C., Vaslet, D., Le Nindre, Y.-M., Brosse, J.-M., Fourniguet, J., 1985. Geologic map of the Wadi al Mulayh quadrangle, sheet 22H, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM-92, scale 1:250,000. Nassief, M.O., 1981. Geology and petrology of the Jabal Thurwah area, Western Province, Saudi Arabia. Ph.D. thesis. University of Lancaster, p. 180. Nassief, M.O., Macdonald, R., Gass, I.G., 1984. The Jabal Thurwah upper Proterozoic ophiolite complex, western Saudi Arabia. Journal of the Geological Society of London 141, 537–546. Pallister, J.S., Stacey, J.S., Fischer, L.B., Premo, W.R., 1987. Arabian Shield ophiolites and late Proterozoic microplate accretion. Geology 15, 320–323. Pallister, J.S., Stacey, J.S., Fischer, L.B., Premo, W.R., 1988. Precambrian ophiolites of Arabia: Geologic settings, U-Pb geochronology, Pb-isotope characteristics, and implications for continental accretion. Precambrian Research 38, 1–54.

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Quick, J.E., 1991. Late Proterozoic transpression on the Nabitah fault system–implications for the assembly of the Arabian Shield. Precambrian Research 53, 119–147. Quick, J.E., Bosch, P.S., 1989. Tectonic history of the northern Nabitah fault zone, Arabian Shield, Kingdom of Saudi Arabia. Directorate General of Mineral Resources Technical Record USGSTR-08-2, p. 87. Ramsay, C.R., 1986. Geologic map of the Rabigh quadrangle, sheet 22D, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM-84, scale 1:250,000. Schandelmeier, H., Wipfler, E., Küster, D., Sultan, M., Becker, R., Stern, R.J., Abdelsalam, M.G., 1994. Atmur-Delgo suture: a Neoproterozoic oceanic basin extending into the interior of northeast Africa. Geology 22, 563–566. Shanti, M., 1983. The Jabal Ess ophiolite complex. Faculty of Earth Sciences, King Abdulaziz University, Jiddah, Bulletin 6, 289–317. Shanti, M., Roobol, M.J., 1979. A late Proterozoic ophiolite complex at Jabal Ess in northern Saudi Arabia. Nature 279, 488–491. Stacey, J.S., Stoeser, D.B., Greenwood, W.R., Fischer, L.B., 1984. U-Pb zircon geochronology and geologic evolution of the Halaban-Al Amar region of the eastern Arabian shield, Kingdom of Saudi Arabia. Journal of the Geological Society of London 141, 1043–1055. Stern, R.J., Johnson, P.R., Kröner, A., Yibas, B., 2004. Neoproterozoic ophiolites of the ArabianNubian Shield. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 95–128. Stoeser, D.B., Camp, V.E., 1985. Pan-African microplate accretion in the Arabian shield. Geological Society of America Bulletin 96, 817–826. Stuckless, J.S., Hedge, C.E., Wenner, D.B., Nkomo, I.T., 1984. Isotopic studies of postorogenic granites from the northeastern Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-04-42, p. 40.

Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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THE WADI ONIB MAFIC-ULTRAMAFIC COMPLEX: A NEOPROTEROZOIC SUPRA-SUBDUCTION ZONE OPHIOLITE IN THE NORTHERN RED SEA HILLS OF THE SUDAN I.M. HUSSEIN, A. KRÖNER AND T. REISCHMANN Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, Germany

The Wadi Onib mafic-ultramafic complex represents the best preserved, though tectonically dismembered, Neoproterozoic (Pan-African) ophiolite in the northern Red Sea Hills of the Sudan. Forming part of a regionally distinct, southwest to northeast trending ophiolite-decorated shear belt (Onib-Sol Hamed suture) it consists, from bottom to top, of a basal peridotite unit, an exceptionally thick (2–3 km) transitional zone (TZ) of interlayered cumulates, an isotropic gabbroic mass with plagiogranite bodies, and a sheeted basic dyke complex. The highest stratigraphic section of the ophiolitic sequence is represented by pillowed basaltic lavas (containing fragmentary lenses of ribbon chert and/or graphitic to shaly carbonates) which are tectonically juxtaposed against the basal peridotite. Whereas the basal unit is strongly serpentinized and/or carbonatized, the transitional zone comprises abundant and well preserved pyroxenites, some of which are enriched in Cr relative to TiO2 . The TZ also shows a polycyclic cumulate arrangement that possibly originated from multiple magma pulses rather than from tectonic interslicing. Moreover, mineral grading, gravity stratification and a spectrum of folds with varying geometrical dispositions and amplitudes within discrete layers as well as a vertical metamorphic zonation (suggesting seafloor hydrothermal processes) are evident within the Onib ophiolitic sequence. In particular, the volcanic component is Ti-rich, has a transitional IAT/MORB character and is indistinguishable from anomalous MORB and/or marginal basin basalts. Thus, the Onib is envisaged to be of arc/back-arc (marginal) basin affiliation, and it is classified as a supra-subduction zone (SSZ) rather than normal MORB-type ophiolite. It was generated at 808 ± 14 Ma as documented by a plagiogranite single zircon Pb-Pb age. The ophiolitic sequence probably resulted from parental magma(s) generated through multistage partial fusion of mantle peridotite. 1. INTRODUCTION The Northern Red Sea Hills (NRSH) of the NE Sudan (Fig. 1) cover some 100,000 km2 between latitudes 19◦ N and 23◦ N and longitude 34◦ 30 E and the Red Sea. They form part DOI: 10.1016/S0166-2635(04)13005-3

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Fig. 1. Simplified geological map of part of the Nubian segment of the Arabian-Nubian shield (after Kröner et al., 1987) showing location of the Wadi Onib ophiolite (arrow), northern Red Sea Hills, Sudan.

of the Neoproterozoic (Pan-African) Nubian segment of the Arabian-Nubian shield (ANS) that straddles the Red Sea perimeter in NE Africa and Arabia. The ANS is the northern continuation of the Mozambique belt and, together, they have been referred to as the East African Orogen (EAO) (Stern, 1994). The ANS represents an excellent example of the Pan-

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African orogenic cycle that has long been recognized as a period of major crustal accretion (Kröner, 1984), where continental, island-arc, and oceanic terranes were brought together (Gass, 1981; Kröner et al., 1991; Reischmann and Kröner, 1994; Kusky et al., 2003) to form the crystalline basement of the African continent as part of the late Neoproterozoic supercontinent Gondwana. The Nubian segment of the ANS contains numerous ophiolite occurrences (Fitches et al., 1983; Hussein et al., 1984; Kröner et al., 1987; Abdel-Rahman, 1993; Zimmer et al., 1995; Reischmann, 2000; Hussein, 2000; Stern et al., 2004; Johnson et al., 2004). These classify as Tethyan-type (found in shallow structural positions with a preponderance towards the margin of the older continent—the Nile craton to the west) or as Cordilleran-type (interspersed within steep suture zones separating the arc terranes) (Reischmann, 2000). Modern-style plate tectonic processes are perceived to have operated within the NRSH, and this is because the bulk of the region is made up of voluminous arc-related volcanoplutonic terranes that are separated by a number of distinct high-strain ophiolite-decorated sutures (Kröner et al., 1987; Hussein, 2000; Kusky and Ramadan, 2002). Lateral accretion and suturing thus contributed significantly to the evolution of the NRSH. The Wadi Onib mafic-ultramafic complex constitutes one of the best preserved, though tectonically fragmented, ophiolitic sequence within the ANS and makes up the major part of the prominent, southwest to northeast oriented Onib-Sol Hamed suture. An improved understanding of the composition and history of this complex has important regional tectonic significance and helps to understand global crustal evolution in the Neoproterozoic.

2. PREVIOUS WORK IN THE ONIB REGION Early geological data on the Wadi Onib and adjacent Deraheib regions were provided by Bagnall (1955) and Gabert et al. (1960) who thought that gabbroic and serpentinitic rocks of the NRSH represent intrusive bodies and/or dykes. The latter authors, in particular, have shown most of the Onib area to be underlain by greenstones which they incorporated into the Nafirdeib Series, a lithostratigraphic term (Ruxton, 1956) embracing widespread volcano-sedimentary sequences, the type area of which is Wadi Nafirdeib in the northern Red Sea Hills. In the late 1970’s a joint French/Sudanese mineral exploration project carried out the first systematic regional geological mapping and mineral prospecting of the Deraheib region. Within the scope of this project Stolojan et al. (1978) and Hussein et al. (1978) delineated the Onib mafic-ultramafic complex and, in the western part of the area, they described a large number of serpentinized and/or heavily carbonatized ultramafic bodies that are interspersed within a prominent shear belt named by Kröner et al. (1987) as the Hamisana Shear Zone (HSZ). Stolojan et al. (1978) and Hussein et al. (1978) reported on the mineral potential of the Onib/Hamisana region and also proposed that the Onib complex is comparable to the Jebel Sol Hamed ophiolite described by Hussein (1977) and Fitches et al. (1983). In particular, in a summary of mineral assessment activities of the French/Sudanese Project, de Bretizel (1980) noted that the Onib complex is a probable ophiolitic remnant.

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Hussein et al. (1984), Kröner (1985), and Kröner et al. (1987) reappraised the Onib and other regions in the Red Sea Hills and recognized pillow basalts and sheeted basic dykes within the Onib mafic-ultramafic suite. These authors advocated that the Onib complex represents a well preserved, though tectonically dismembered, Pan-African ophiolite.

3. ANALYTICAL PROCEDURES Based on detailed petrography, representative suites of ultramafic rocks, gabbros, basic dykes and basaltic pillow lavas as well as chromitiferous ores were selected for major and trace element analysis. Each rock sample was pulverized (< 0.06 mm) using an agate vessel fitted into a Siebtechnik mill. The powders were then decomposed into stock solutions which were analysed for major elements using directly coupled (DC) plasma emission spectrometry. The rock powders were also compressed into duplicate pellets and were analysed for trace and light rare earth elements (LREE) using a Siemens SRS 200 X-ray fluorescence spectrometer. Analytical details are given in Laskowski and Kröner (1984). Polished and thin sections were prepared from selected mafic and ultramafic rocks as well as chromite ores. The sections were used for major element analysis of primary and secondary silicate mineral phases and chromite ores using a Camebax Microbeam electron microprobe fitted with four wavelength dispersive spectrometers, one energy dispersive spectrometer and the capability of electron microscope scanning. The accelerating voltage was 15 kV, and the beam current was 10 nA. All measurements were corrected by the ZAF procedure (Atomic Number, Absorption Fluorescence correction), and oxygen contents for the oxides were calculated by stoichiometry. The standard deviation for all components measured is about 1%. For whole-rock major element analyses the DC plasma emission spectrometer was calibrated with artificial standards and checked against international rock standards (NIM-G (granite), NIM-S (syenite), NIM-N (norite), NIM-D (diorite), NIM-P (pyroxenite), BHVO1 (basalt), DTS-1 (dunite), PCC-1 (proxenite), RGM-1 (rhyolite), GSP-1 (granodiorite), SDC-1 (mica schist), STM-1 (syenite), JB-1 (basalt), BR (basalt) and GA (granite)) (see also Krüger, 1982). The accuracy of the method was about 2%, whereas the precision was about 1%. For major elements determined on the electron microprobe, international rock standards (BCR-1 and G-2) were measured as unknowns. The accuracy of the measurements was mostly better than 1%. The trace element analyses by X-ray fluorescence spectrometry were calibrated against international standards. For details see Tables 1–6 in Hussein (2000).

4. OUTLINE OF THE GEOLOGY OF THE ONIB OPHIOLITE The Wadi Onib ophiolitic complex (Figs. 1–5) forms a large (ca. 80 km × 10 km), smoothly curved exposure and can be reached from Port Sudan via a number of desert

4. Outline of the Geology of the Onib Ophiolite

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Fig. 2. Geological map of the Wadi Onib ophiolite (after Hussein, 2000).

tracks that run westwards from an all-season coastal road along the Red Sea. Satellite imagery as well as geological mapping (Hussein et al., 1984; Kröner et al., 1987; Stolojan et al., 1978; Hussein et al., 1978) reveal that the Onib complex is enclosed within multiply deformed, weakly metamorphosed arc-related volcano-sedimentary sequences and that the contact between the two suites is a thrust fault everywhere. Though the basalt-andesite arc assemblages locally show rapid facies changes along a predominantly southwest to northeast bedding/foliation strike, they generally display comparable lithotypes. To the east of the Onib complex, there are widespread basaltic andesite/andesite lava flows in addition to restricted dacitic and/or quartz-crystal tuffs. The lavas are accompanied by abundant andesite-derived clastic sediments, tuffaceous greywackes, shales,

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Fig. 3. Distribution of ophiolitic lithologies along Wadi Onib (A) and Wadi Sudi (B). Locations of both Wadis are indicated in Fig. 2.

4. Outline of the Geology of the Onib Ophiolite

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Fig. 3. (Continued.)

polymictic intra-arc conglomerates (some of these contain rare ophiolitic clasts) as well as extensive limestone interbands that locally show stromatolitic structures. The conglomeratic and limestone beds in particular may be traced along strike for tens of kilometres (e.g., between wadis Onib and Sudi). To the southwest of Onib the arc assemblage includes sheared lavas, greywackes and siliceous sediments. Within the Hamisana shear zone to the west the arc sequences are mylonitized to the extent that, at places, it becomes difficult to identify protoliths in the field. The Onib ophiolite and arc-related volcano-sedimentary assemblages are intruded by widespread syntectonic granitoids of batholithic proportions (composite dioritic/granodioritic/granitic plutons) and/or by younger (post-tectonic) gabbroic/granitoid stocks.

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Fig. 4. Composite stratigraphic column for the Wadi Onib ophiolite (compiled from Kröner et al., 1987 and Hussein, 2000).

In places, the younger magmatic rocks form distinct ring structures (for details see Hussein, 2000).

5. OPHIOLITE LITHOLOGIES AND LITHOSTRATIGRAPHY The Onib complex displays an orderly succession that defines a Penrose-style (Anonymous, 1972) ophiolitic suite (Fig. 4). Though swelling, pinching out and shearing, thrusting and interslicing of magmatic layers prevail and an entirely intact ophiolitic sequence is not preserved, five ophiolitic lithostratigraphic units have been reconstructed through detailed

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Fig. 5. Schematic cross-sections showing disposition of lithologic components of the Onib ophiolite between Wadi Ader Aweb (1) via Wadi Sudi (2) and Onib (3) and as far as Wadi Hofra (4), northern Red Sea Hills, Sudan. A section across the Jebel Sol Hamed ophiolitic complex (5) farther northeast (see sketch map, top left) is shown for comparison (modified after Hussein, 2000).

field work carried out along major dry river beds (viz. Sudi and Onib) and/or side channels. The sequential ophiolitic stratigraphy begins, at the bottom, with a discontinuous, ca. 20 km × 2 km, extensively serpentinized and/or variably carbonatized basal peridotite. This unit is succeeded by an exceptionally thick (ca. 2–3 km) transitional zone (TZ) of interlayered cumulates including layered melano-to-mesogabbros, serpentinized dunitic and pyroxenitic peridotites in addition to well preserved olivine- and/or chromian spinel-bearing pyroxenites and wehrlitic transitions. At upper levels the TZ displays a gradual diminution of both mafic gabbroic and ultramafic rocks. Thereupon, a large (ca. 80 km × 2–6 km) gabbroic mass, dominated by crudely to well layered leucogabbros, appears in the magmatic sequence. At upper sections the layered gabbros intermingle with high-level isotropic gab-

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bro, hornblende gabbro and quartz-diorite (plagiogranite). These rocks were frequently injected by single, multiple and/or criss-cross microgabbroic, microdioritic and/or mafic to intermediate dykes. In particular, a sheeted mafic dyke complex that displays asymmetrical chilling of dyke margins occurs in lower Wadi Sudi. A block of mafic dyke-intruded basaltic pillow lava containing rare fragments of pelagic sediments (ribbon chert and/or graphitic to shaly carbonates) represents the highest stratigraphic section of the ophiolite sequence. The latter unit is tectonically juxtaposed against the basal peridotite of uppermost Wadi Onib and tributaries. 5.1. The Basal Ultramafic Unit The Onib basal unit is dominated by an extensively serpentinized peridotite that hardly displays any of the diagnostic features of metamorphic tectonites (e.g., gneissic foliation and linear fabrics due to subparallel orientation of stretched orthopyroxene porphyroblasts). It is best exposed around uppermost Wadi Sudi and environs where the unit and adjacent strongly sheared and/or mylonitized lavas, greywackes and siliceous sediments of the arc assemblage are in tectonic contact along a vertically to steeply eastward-dipping thrust fault. The serpentinized peridotites are generally fine-grained, smooth, massive, distinctly obdurate, dark greyish-black and/or greenish in colour. Whitish magnesite-vein infillings, reddish nets of iron stains in addition to medium brown colouration are common across these ultramafic rocks. Some 1.5 km downstream from the tectonic contact zone a mélange-like magmatic breccia and carbonatized serpentinites form part of the basal peridotitic suite. In particular, the breccia contains numerous disoriented blocks of up to several metres in diameter. Occurring within a weathered, fragile, yellowish-brown, mediumgrained pyroxenitic matrix, the blocks include greyish-brown, fine-grained serpentinized dunite and an exceptionally fresh and remarkably coarse-grained protogranular pyroxenite which shows pegmatoid textures at places. Hand specimens from the pyroxenitic blocks are steel-black to bluish-black on fresh surface and brown to yellowish on weathered surface. Longer dimensions of pyroxenes in these rocks can reach 3–5 cm in length and some 0.5 to 1.5 cm in width. We interpret this rock as a localized primary magmatic breccia. Farther north (upper Wadi Onib and adjacent areas) the basal unit is faulted against the pillow lava block. The contact zone is characterized by strong shearing and/or transformation of some rocks into greasy, brownish-white, talcose material. It is crossed by a few mafic dykes which seem to post-date the deformation of the pillow lavas. The Onib basal ultramafic mass locally shows clear magmatic contacts between fresh to mildly serpentinized interlayers. Magmatic interfingering is evident where dunitic schlieren penetrate pyroxenitic host rocks. 5.2. The Transition Zone (TZ) Transition zones in layered mafic-ultramafic complexes (Wilson, 1959) usually represent marker horizons for a gradual change from cumulate peridotites to cumulate gabbros. Various estimates show that they rarely exceed 750 m in thickness (Coleman, 1977). The

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Onib TZ, however, may thicken up to 2–3 km at places. In the main part it consists of a well preserved, lensoid pyroxenitic suite that grades into a crudely layered to a massive melano- to mesogabbroic assemblage containing frequent serpentinite layers and some restricted anorthositic lenses. The TZ ultramafic bodies vary in thickness and lateral extent from metres to kilometres in scale but average hundreds of metres in their long dimensions. Swelling out and/or sudden disappearance of the cumulate layers are common. In places, sharp magmatic contacts between ultramafic lenses and gabbroic host rocks occur on a metric scale as well. The pyroxenitic suite of the TZ is dominated by greyish-green, medium to coarse-grained olivine-rich pyroxenites and wehrlitic transitions, which form rugged hillocks and merge into the basal unit dominated by the serpentinized peridotite. In the same area the peridotitic rocks contain sporadic podiform (high-Cr) and disseminated (high-Al) chromite lenses and selvages that locally show size-grading and/or sharp magmatic layering (Fig. 6C); here the mineral stratification attests to and confirms the cumulative nature of the hosting pyroxenites and of the TZ. A pyroxenitic rock from the area is almost an ore-bearing host and is found to contain almost 40% Al2 O3 when chemically analysed. The widely distributed serpentinite bodies within the melano- to mesogabbros are finegrained, massive and range in colour from dark-brown to greyish-green. Their magmatic contacts with the gabbros are generally sharp though tectonic slices of serpentinite, detached from the basal peridotite unit, may belong to the TZ. Furthermore, it is not uncommon to observe gabbroic and serpentinitic rocks tightly folded together, a criterion that equally supports the primary nature of the TZ cumulates. Shear zones lacing deeper-level serpentinitic bodies of the TZ contain conspicuous, excessively carbonatized and/or silicified, yellowish-brown, massive, amorphous and inhomogeneous dyke-like bodies that range from centimetres to hundreds of metres in extent. Such silica-carbonate rocks are also known from the Jebel Sol Hamed ophiolitic complex near the Red Sea coast to the northeast (Hussein, 1977). Carbonatization (listwanitization) and/or silicification of the ultramafic rocks of the TZ and elsewhere are believed to be of metasomatic origin. Breakdown of serpentinite into carbonatized rock may have taken place through complicated processes involving release of carbon dioxide and its combination with oxygen and magnesium to form MgCO3 out of an essentially magnesium-aluminum-silicate serpentinitic structure. Alumina and silica will be removed as silicates and free quartz which is sometimes found as tiny aggregates or veinlets. Some of the iron, which originally existed in the olivine and/or pyroxene crystals, may be oxidized to form a pervasive reddish-brown colouration of these rocks. According to Buisson and Leblanc (1986) hydrothermal alteration at moderate temperatures (150–300 ◦ C) by Na- and Cl-brines derived from mantle material and interaction with seawater is perceived to be responsible for the development of carbonatized bodies from their ultramafic protoliths. The latter authors highlight that the listwanitization processes may be responsible for some gold mineralization (in quartz gangue), sulphides and arsenides known from the Bou Azzer complex, a Pan-African ophiolite in southern Morocco (Bodinier et al., 1984).

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5.3. The Gabbroic Unit Upwards in the magmatic sequence the TZ grades, through a gradual diminution of ultramafic layers, into a large (ca. 80 km × 6 km) gabbroic mass that forms the bulk of the Onib ophiolite complex. At deeper levels of the gabbro suite some dark, medium to coarsegrained, crudely layered, partly massive pyroxene-gabbros predominate that may thicken up to 1.5 km and may contain localized melagabbroic/anorthositic layers. The pyroxene gabbro was subjected to metamorphic recrystallization involving sea-floor (hydrothermal) metamorphism. Strong shearing has disrupted the continuity of individual gabbroic layers which cannot be followed for more than a few hundreds of metres along the roughly southwest to northeast strike of magmatic layering. In turn, the pyroxene-gabbros grade upwards into well layered, medium to coarsegrained, light grey to greenish-grey, plagioclase-rich leucogabbros that dominate the magmatic sequence farther east. In places, the leucogabbros contain rare serpentinitic lenses (tectonic slices?) and/or deformed pegmatoid wedges. They frequently display magmatic layering (Fig. 6E), gradual vertical change of cumulus mineral phases in overturned mineral graded stratification as well as combinations of continuous and sharp mineral grading with indications of cross stratification. The gross igneous stratification allows recognition of way-up criteria. This suggests that although there are northwest and southeast-dipping igneous laminations (variable orientations are due to large open folds deforming older, isoclinal folds and due to abundant low angle thrusts), the overall stratigraphic vergence is towards the east (i.e., the overall younging of the gabbroic layer is in that direction; see also Hussein et al., 1984). The layered gabbroic suite merges gradually upwards into a non-cumulate, isotropic gabbro which forms low hillocks to the SW and W of the Sudi well (Fig. 3B). Isotropic gabbro, however, is missing from the Wadi Onib area to the north, most probably due to the inconsistency of stratigraphic horizons across the entire ophiolitic sequence. The contact zone between the layered gabbroic pile and the isotropic gabbro is difficult to ascertain. High-level hornblende gabbro and quartz-diorite are associated with the isotropic

Fig. 6. (A) Photomicrograph (crossed nicols) of an Onib olivine-bearing pyroxenite (transitional zone of middle-upper Wadi Sudi) with abundant high-Al chromian spinel (dark subhedral crystals at centre). Length of photograph is approximately 2.5–3 mm. (B) Photomicrograph (crossed nicols, ca. 3.5 mm in length) of fresh Onib pyroxenite (uppermost Wadi Sudi) depicting twinning as well as deformed exsolution lamelle in clinopyroxene. (C) Mineral size-grading in chromite from Onib ophiolite (upper-middle Wadi Sudi). (D) Photomicrograph (crossed nicols) showing chromite in serpentinized peridotite (upper-middle Wadi Sudi). Note crude layering and/or pull-apart segmentation displayed by some of the chromite granules. Length of photo is about 2.5 mm. (E) Mineral grading in typically layered Onib ophiolitic gabbro (lower Onib gorge). (F) Tight isoclinal folding indicating high temperature ductile deformation in layered Wadi Onib ophiolitic gabbro (upper-middle Wadi Sudi). (G) Disposition of feeder (sheeted) dykes and lavas of Onib ophiolite (area of uppermost Wadi Onib). (H) Well-preserved Onib pillow lavas (uppermost Wadi Onib). Note tightly packed pillow lobes surrounded by chilled (fine-grained) interpillow matrix.

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gabbro. Restricted leucocratic, dyke-like bodies and thin plagioclase-rich veins also occur in places across this gabbroic lithotype. The leucocratic rocks contain minor plagiogranite bodies. The Onib ophiolitic gabbroic mass, in general, contains numerous mafic dykes, the frequency of which decreases westwards, a feature which implies a genetic link with the sheeted dyke complex farther to the east. 5.4. Sheeted Dyke Complex Sheeted doleritic dyke-in-dyke units are common within ophiolite complexes. They have been almost universally considered as convincing indicators of seafloor spreading at zones of plate accretion (Moores, 1982) and, consequently, provide evidence that ophiolitic sections represent ancient oceanic lithosphere. The sheeted dykes are generally characterized by asymmetric chilling of their margins. This is conceived to be the result of injections along a single fracture where each successive dyke is emplaced within the middle of the previously solidified dyke (Moores and Vine, 1971). A wide range of single, multiple, locally criss-crossing or swarmy doleritic, microgabbroic, microdioritic and porphyritic plagioclase andesite dykes and sills occur within the Wadi Onib complex. Emplaced at varying trends, they range in width from several cm to metres and extend in lengths over tens of metres. In general terms, all dykes experienced varying degrees of deformation (e.g., contortion and shearing) as well as low grade metamorphic alteration. Of special interest are dykes that form a sheeted block about 1.5 km west of the Sudi well (Figs. 2 and 3B). These are mafic dykes that invade each other and/or isotropic, microgabbroic to dioritic assemblages. The dykes have intrusive contacts, frequently carry screens of microgabbro, show doleritic textures and at places have chilled margins. The chilled margins do not exceed 10 cm in width, and asymmetric chilling is common. The sheeted dyke block is over one kilometre in length and widens up to 300 m. It merges into a transition from microgabbro to diorite which, in turn, grades into isotropic gabbro that passes westwards into the plagioclase-rich, well layered leucogabbroic mass outlined earlier. At the eastern side, the sheeted dyke complex is bounded by deformed arc-related lavas and volcaniclastic sediments through a sheared contact mostly obliterated by rock debris and superficial deposits. Strong shearing along this contact zone has transformed all rocks into schists. The sheeted dykes of Wadi Sudi are generally oriented north-northwest to south-southeast, approximately perpendicular to the principally southwest to northeast magmatic layering of the igneous cumulate rocks. However, due to post-emplacement deformation, this is by no means the only trend for the dykes which are frequently irregular in pattern. Along the western side of the ophiolite, specifically near the western tectonic contact between the basal ultramafic layer and the pillow lavas, there are also closely spaced basic dykes possibly of the sheeted dyke complex (Fig. 6G). The pillow lavas (upper Wadi Onib) were injected by such dykes which are subvertical. Over 13 east to west dykes can be seen within several metres across the stratigraphically upper level but structurally lowermost extrusive section. These magmatic injections range from 50 to 150 cm in width and mea-

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sure up to several or tens of metres in length. Their contacts with the pillow lavas are sharp and intrusive, but with negligible chilling along their margins. They are interpreted here as feeders for the lavas since both assemblages are similarly deformed. 5.5. Pillow Lavas and Sedimentary Cap Direct evidence for the existence of pillow lavas on top of the sheeted dyke complex is lacking. It may be that the pillow lavas of the ophiolite have been tectonically imbricated with the arc lavas and, therefore, the units have become indistinguishable. However, to the west there is a discontinuous, ca. 300–800 m wide volcanic strip that stretches SWNE parallel to the overall lithological layering of the ophiolite (Fig. 2). It is tectonically juxtaposed against the basal, serpentinized peridotite and/or cumulate gabbroic layers and is sharply intruded by batholithic granitoid masses farther west. The volcanic exposure comprises massive and pillowed basalt, breccias, sills, and many dykes, some of which are feeders as outlined earlier. These rocks are dark to greyishgreen, extensively sheared and amphibolitized locally where intruded by the batholithic granitoids. Near the contact zone, distinct, tightly interlocked pillow structures occur in several basaltic lava outcrops. Subcircular to lenticular pillows (Fig. 6H) dominate over other forms which take the shape of circular, sickle-shaped, roughly triangular and other structures. Pillow margins, or inter-pillow sheaths, are thin (3 mm to 1 cm in width) and are essentially formed of bluish-grey or greenish chloritic aggregates and/or quartz material. Individual pillows have voids and vugs, some of which are filled with siliceous and carbonate material. The latter material, filling the vesicles and hollow cores, may be used as an indirect criterion to estimate the depth of lava extrusion. Moore (1979) advocates that the primary control of vesicularity of subaqueously erupted basalt is the depth (and hydrostatic pressure) of eruption. Shallow eruption produces more (and larger) vesicles not only because exsolved volatiles expand with lower pressure, but also because exsolution of volatiles dissolved in the melt is more complete. In the Sol Hamed ophiolite, the analogue and probable extension of the Wadi Onib ophiolite, pillow lavas were estimated to have been extruded at a water depth of about 1–2 km (Price, 1984). In addition, the upper Onib sheared pillow lavas also contain restricted, fragmentary bluish to brown ribbon chert and black, graphitic, shaly carbonate lenses. The pelagic sediments are several centimetres to a metre long and up to 5 cm thick. They generally occur as thin infillings and have experienced intense shearing together with their host pillow lavas. It may also be that the sediments were ripped off from a much thicker cap that once covered the pillow lavas.

6. INTERNAL STRUCTURE AND DEFORMATION HISTORY OF THE ONIB OPHIOLITE Structures developed during ductile flow and late stage congealing of magma before and during final crystallization (Davis, 1984) are reasonably well preserved within the

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Onib deeper sections (basal unit, TZ and layered gabbros). The principal primary structures within the lower ophiolitic sections are the cumulate fabric and lithological layering. In the TZ pyroxenites, the mineral-size grading of chromite (Figs. 6C and D) is evidence for cumulate formation within a magma chamber. Irregular, anastomosing, magmatic interlayering of peridotite with pyroxenite/dunite schlieren occurs locally in upper Wadi Onib. Sinuous contacts within the schlieren constitute flow lines that facilitated the development of a high-temperature ductile fabric. Fitches et al. (1983) also recorded a penetrative flattening fabric within ultramafic and layered gabbroic rocks of the Jebel Sol Hamed ophiolite complex. They interpreted the structure as a high-temperature ductile fabric, assumed to be related to horizontal ductile shearing in the uppermost mantle. Moreover, the cyclic nature of the cumulate rocks of the TZ and contiguous layered gabbros can be conceptualized in the framework of magmatic differentiation. Pearce et al. (1984a) described such an order of lithologies in ophiolites which they interpreted as having formed in a supra-subduction zone (SSZ) setting. However, Shervais (2001) advocated that SSZ ophiolites display a consistent sequence of events during their formation and evolution; entailing birth, youth, maturity, death and, finally, resurrection—a progression implying that ophiolite formation is not a stochastic event but is a natural consequence of the SSZ tectonic setting. Evidence of secondary structures in the Onib complex is provided by abundant folds ranging in size from microscopic crenulations (lineations) through meso- to macroscopic structures. From field observations it has been possible to separate at least three phases of deformation. This is indicated by folds with varying orientations and amplitudes within discrete layers. High-temperature isoclinal folding is recorded within interlayered peridotites and gabbros (Fig. 6F). This is probably one of the earliest structures and may have formed whilst the rocks were still hot and partially consolidated. At places mm-thick bands of metamorphic mineral growths (foliation) can be seen crossing the isoclinally folded magmatic layering. The younger bands of metamorphic minerals (e.g., amphiboles) may have developed during greenschist- to lower amphibolite-facies metamorphism. The syn-igneous (?) folds are best preserved in the lower levels of the TZ (upper-middle Wadi Sudi). Their dimensions vary from a few to tens of centimetres. The folds are generally asymmetrical to parallel, tight, and are commonly accompanied by high-temperature ductile shearing. At places there is clear transposition of folded igneous layers due to strong stretching, and the folds are intrafolial. They normally disappear along strike of the igneous layers. Tight to open, asymmetric folds, plunging at 30◦ to 60◦ northeast, are well developed in the Onib TZ and overlying layered gabbro. These folds are cut by penetrative schistosites, some of which are oriented parallel to the generally southwest to northeast trending fold axes. In the middle of Wadi Onib, layering of the gabbro is bent into a clearly asymmetric fold with limbs separated by about 90◦ . Farther west (upper middle, left side of Wadi Onib—Fig. 3A) highly contorted serpentinitic peridotite interlayered with gabbro can be seen within a section of 300–400 m. The folds in this area are tight, a few to several metres in amplitude, and make up a repetitive synform/antiform structure. They are locally disrupted by steep shear zones along which dislocations of layers can be seen. The folds also affected some of the mafic dykes crossing the plutonic ophiolite lithologies. The asymmet-

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ric, tight-to-open folds that deform the igneous layering of the cumulates are believed to have developed during early phases of ophiolite obduction. At a few places early (?) shear zones cross the limbs of such folds, and the shear zones are locally annealed by mm- to cm-wide fillings of oxidized pyrite crystals (e.g., contorted gabbro some 5 km NW of Sudi well—Fig. 3B). The shear zones themselves are cut by later thrusts which are mainly shallow to subhorizontal and are apparently responsible for the interstacking and thickening of the ophiolite complex. Some late, gentle open folds are tens of metres to km in size and post-date all the structures outlined above. They occur in both the ophiolitic sequence and the adjacent VAA strata and may belong to syn-to-post-ophiolite emplacement deformation events.

7. PETROGRAPHY OF THE ONIB OPHIOLITE SEQUENCE Although the Onib complex was affected by serpentinization, metasomatic and ocean floor hydrothermal alteration in addition to regional metamorphic re-equilibration, detailed petrographic and geochemical studies made it possible to characterize several rock types that form the major lithological units outlined earlier. Details pertaining to the lithotypes of the plutonic and volcanic components are summarized below. 7.1. Dunite and Harzburgite Dunitic exposures within the basal ultramafic unit are inferred, indirectly, from the degree of carbonatization and/or the development of magnesite (vein-infillings and/or encrustations) in addition to a yellow-brown colouration on weathered surfaces. Harzburgitic rocks were not mapped as separate bodies in the Onib ophiolite. However, it is not uncommon to encounter partially serpentinized harzburgitic kernels that “swim” within strongly serpentinized assemblages. Such rocks are medium to coarse-grained in texture and range in colour from steel grey to black, dark brown, and pale green at places. Harzburgitic rocks in ophiolites are generally interpreted as representing refractory residues of a depleted mantle (Gass, 1980). They may also be essential components of a polycyclic igneous assemblage that shows distinct cumulate textures (e.g., size grading and parallelism of mineral phases such as chromite), kelyphytic growth of adcumulates as well as elongate, euhedral to subhedral and irregular, sharp grain-to-grain boundaries of some of the mineral components such as olivine and/or orthopyroxene. Serpentine derived from dunitic rocks has low relief, fine-grained texture and shows nets of mesh- and/or bladed as well as hour-glass structures (lizardite? and/or antigorite?). It constitutes some 95 to 99% of the rock, the rest of which includes scanty relicts of olivine and/or rare, pseudomorphosed orthopyroxene in addition to minor amounts of magnetite and/or chrome spinel. Some of the serpentine phases are crossed by silky serpentine and/or chloritic veinlets that probably developed during post-obduction regional metamorphism. Magnetite, in particular, occurs as a secondary product of serpentinization (black streaks, feathery granular aggregates and shapeless to irregular spots) or as rare, subhedral,

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smoothly outlined accessories that could be of primary origin. Its disposition helps to indirectly recognize outlines of precursor olivine and/or orthopyroxene crystals. Subhedral and/or polygonal grain-to-grain boundaries of olivine are frequently preserved by magnetite trails and trains; this suggests a cumulate origin for the Onib serpentinized dunites. The magnetite grains also form linear tracts of exsolution lamellae of pseudomorphosed orthopyroxene, a phenomenon suggesting high temperatures and slow cooling. Furthermore, the accessory chromite grains range in colour from black to various tints of coffee-brown to reddish, signifying alteration into ferrichrome. Pull-apart structures and/or criss-crossed cracks have formed within the spinel granules some of which contain inclusions of flaky plates of serpentine, a phenomenon indicating that olivine was probably occluded as inclusions. Elongation and tadpole structures shown by some of the chrome accessories, in particular, are due to strong ductile deformation. These are high-temperature fabrics and were also reported from ophiolite complexes elsewhere (e.g., see Burgath and Weiser, 1980; Nicolas and Al Azri, 1991; Li et al., 2002; Huang et al., 2004). The harzburgites, on the other hand, consist of strongly to completely serpentinized olivine (up to 85%) and orthopyroxene (10–15%) in addition to substantial percentages of magnetite and minor, but ubiquitous, accessory chromite granules. Though the chromite grains are predominantly equant, euhedral and/or subhedral, they display elongation as well as pull-apart (?) segmentation and development of tapering ends in places. Relict olivine and/or orthopyroxene crystals showing poikilitic and/or kelyphitic interrelations are evident. Olivine engulfed by orthopyroxene (or bastitic pseudomorphs) and, in places, contrasts between large and small serpentine flakes are interpreted as indicators of original textures. Rare clinopyroxene crystals were locally encountered among the dominating mineral phases. In general terms, the serpentinized harzburgitic rocks consist essentially of meshtextured lizardite?, bladed antigorite? and tabular, or fibro-lamellar, bastitic? aggregates. Where bladed antigorite? phases become dominant, this feature may be correlated with advanced stages of tectonic mobilization which normally leads to recrystallization of the early serpentine phases as proposed by Prichard (1979). Serpentinization of orthopyroxene usually takes place through consumption from the outer margins inwards until the crystals are pseudomorphosed or left as tiny patches (oikocrysts?) surrounded by serpentine flakes. In addition, micro-structures related to high-temperature ductile deformation are common (e.g., as segmentation and development of tapering ends of elongate chromite grains as well as smooth, S-shaped bending and/or curvature of orthopyroxene crystals). Furthermore, in one harzburgitic section a triple-phase structural relation is indicated whereby anomalously yellow and/or greenish serpentine fibrils (probably chrysotile? grown along slip movement directions) cross serpentine blades which themselves penetrate into orthopyroxene crystals. In the oceanic realm the orthopyroxenes were probably subjected to further alteration through hydrothermal solutions at elevated temperatures to produce serpentine phases (e.g., bastite and/or modifications?). During ophiolite obduction, however, tiny blades of serpentine recrystallized, and these were then crossed by late-stage chrysotile? serpentine that shows the least indications of deformation. The latter serpentine phases, in particular, may well be related to post-obduction regional metamor-

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phism which resulted in greenschist-facies conditions imprinted on the ophiolitic sequence and enclosing arc-related strata. 7.2. Pyroxenite and Transitions The Onib pyroxenitic assemblage includes exceptionally coarse-grained clinopyroxenite of the basal unit in addition to olivine-rich, spinel-bearing pyroxenite (Fig. 6A) and wehrlitic/websteritic pyroxenite and/or other transitions within the overlying cumulate TZ. The Sudi clinopyroxenite displays very coarse-grained, holocrystalline (hypodiomorphic granular) textures and consists predominantly of augitic clinopyroxene showing two-directional exsolution lamellae which are deformed (Fig. 6B) into zig-zag patterns. Clinopyroxene includes blebs of secondary amphibole. Bending and twisting of the minerals and their lamellae strongly suggest that polyphase deformation has affected the Sudi clinopyroxenite; at least one phase is thought to be related to ductile deformation, indicated by tight isoclinal folding of magmatic layering. Microprobe analyses of the exsolution lamellae within clinopyroxene indicate compositions close to chromian augite and/or ferroaugite (Deer et al., 1966). Only localized blebs within the augitic crystals show relatively high amounts of sodium, indicating local inversion of pyroxene into amphibole. Opaques within clinopyroxene are magnetite and, less commonly, chromian spinel. Within the lower sections of the TZ the most abundant pyroxenitic lithofacies is olivinerich spinel-bearing pyroxenite. This rock, which is inferred to be a cumulate, is medium to coarse-grained equigranular and consists of clinopyroxene in the main part (∼ 60–75%) as well as varying proportions of orthopyroxene and/or olivine (which is variably serpentinized) and accessory amounts of chromite and magnetite. Clinopyroxene and orthopyroxene reveal a common mosaic fabric. The pyroxene crystals are characterized by schiller structure and mutual intergrowths at places. Deformational effect are shown by granulation, cracking, twisting and/or minor dislocations of exsolution and twin lamellae, bending of entire crystals as well as wavy, oscillatory and fan extinction. Other textural features are relict embayments, occluded relations (where recrystallized pyroxene grains surround serpentine mantles that include clinopyroxene remnants) and deformed lamellae that are not associated with kelyphitic recrystallization. Optical and microprobe data from olivine-rich, spinel-bearing pyroxenites indicate a predominance of augite, diopside-augite, endiopside and, possibly, pigeonite and subcalcic augite at places (clinopyroxene) as well as enstatite bronzite and/or hypersthene. Olivine is either recrystallized into smaller grains and/or partially serpentinized. Wehrlitic transitions from olivine-rich spinel-bearing pyroxenite are not uncommon. Microscopic investigations of wehrlite indicate that it consist essentially of olivine (∼ 65%), a substantial amount of clinopyroxene (∼ 30%), with or without trace amounts of orthopyroxene. These mineral phases are accompanied by accessory amounts of spinel that may constitute up to 2% of the mode at places. Olivine is variably serpentinized. Whereas some olivine grains were replaced by serpentine along cleavage fractures and borders, others were totally consumed. In places it is possible to recognize olivine pseudomorphs which are enclosed in fresh, though strained, clinopyroxene within wehrlitic sec-

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tions. Spinel is varying in abundance and type from one wehrlite to another. Dark opaque subhedral (local euhedral) chromite grains and clots are common. Another accessory phase is characteristically greenish and translucent under plane-polarized light and isotropic under cross-polarized light. This is an aluminous spinel which is abundant within olivinepyroxenite in the SW section of the TZ. In addition to spinel, black magnetite dust and granules are commonly seen within fractures of serpentinized olivine or within partings of orthopyroxene. Other pyroxenite assemblages within the TZ include websterite and websteritic-olivineclinopyroxenite. From a modal point of view websterite consists of comparable proportions of ortho- and clinopyroxene. The websteritic olivine-clinopyroxenite contains orthopyroxene, clinopyroxene and olivine in decreasing abundance. 7.3. Podiform Chromite The deeper levels of the Onib TZ comprise sporadically distributed, disseminated or speckled (high-Al) and restricted massive (high-Cr) chromian spinel occurrences. Petrographic studies of the disseminated chromite under transmitted and reflected light reveal that the ore occurs in interlocked to disparate anhedral to subhedral, brownish to greenish grains with silicate phases filling the interspaces. The ore shows relict cumulate textures (including poikilitic relations, i.e., chromian spinel enclosed by pyroxene crystals) though it is unquestionably deformed (cracked and microbrecciated grains are evident at places). It varies from a low, accessory level up to 30% or more of the olivine-pyroxenitic host rock (see Fig. 6A). Martitization (haematitization of magnetite veinlets associated with greenish and/or brownish spinel—the colours are internal reflections of the latter mineral) is common. This is possibly due to the deformational effect and/or metamorphic transformation that has affected the primary mineralogy and fabric of the ophiolite. On the other hand, the high-Cr chromian spinel is far richer (forming up to 80% of the host serpentinite) than the disseminated (high-Al) variety and, under the ore microscope, clearly forms a chromitite. This is generally massive and is made up of crudely layered, compactly interlocked and/or interspaced subhedral to anhedral, dark-brown to amber coloured (plain polarized light) grains and aggregates. The chromite is intricately intergrown with serpentine flakes which contain tiny magnetite dust in places. Ferrichrome rims and/or martitized magnetite alteration zones within the chromite grains are common. A rather rare mineral associated with the chromite is some metallic-yellow, anisotropic suspected sulphide (hazelwoodite nickel and/or chalcopyrite?) which is noticeable across some polished chromitite sections. Platinoid granules (?) are also not ruled out from the suspect sulphides. The subparallel disposition (alternation of ore and silicate mineral sheaths), the existence of some euhedral grain forms and occluded textural relations with the serpentine phases principally indicate a cumulate origin for the massive chromite as also advocated for the disseminated variety. The Onib chromitite, however, is believed to have been subjected to a complex, postaccumulation deformational history. Occasional pull-apart segmentation structures (e.g., Fig. 6D), brecciation and invasion of some chromite grains by strings of serpentine and/or

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carbonate infillings as well as the presence of smoothly-bounded, curvilinear to irregular clots and one-side-tapering (development of tadpoles?) chromite grains of 1 to 2 mm in length, are common. The tadpoling of the chromite granules is interpreted as an indication of high-temperature ductile deformation. It is broadly comparable to flattening and elongation of chromite parallel to foliation and lineation reported by Li et al. (2002) from an Archaean ophiolitic peridotite in northern China. These authors attribute the phenomena to intensive high-temperature shear strain and that the textures probably record the plastic flow of the upper mantle, now mainly preserved in the core of tectonic blocks. Tadpole grains (i.e., boudinaging) of chromite can be brought about by solid state flow as documented and interpreted by Burgath and Weiser (1980) who studied primary features and genesis of podiform chromites from Greek ophiolites. These authors considered chromite stretching as the first step of boudinage which, on a large scale, can explain the formation of the dissected, lensoid podiform chromite deposits in the ophiolites of Greece. Holtzman (2000) investigated mantle chromitite pods in the Oman ophiolite and explained that at some time after chromite has crystallized, it begins to deform by fracture and possibly creep mechanisms in a matrix deforming mainly by dislocation creep. He summarized that among the only strain markers in peridotite, the chromitite pods record the kinematics of corner flow and rheological transition from asthenosphere to lithosphere. Thus, as chemical heterogeneities, they provide opportunities for insight into mechanical and petrologic processes of melt migration in the upper mantle. 7.4. Melano- and Mesocratic Gabbro These rocks are essentially pyroxene-gabbros with inferred gabbronoritic gradations. They are dark in colour and medium to coarse-grained in texture. Mineralogically they are simple, due to a notable absence of olivine as well as considerable recrystallization of pyroxene into uralitic amphibole and plagioclase into saussuritic material. For example, melanocratic gabbro from the middle to uppermost Wadi Onib consists of some 70% uralite, ∼ 25% saussurite and minor amounts of zoisite and chlorite as well as traces of magnetite and sphene. Mesocratic gabbro is much more abundant than the mela facies and contains up to 45% saussuritic aggregates, the main components of which are sericite, albite, carbonates and/or laths of ink-blue epidote (zoisite) that has straight extinction and residual twin lamellae. Whereas uralite (tremolite/actinolite?) occurs as fibrous blades which are non-uniformly bleached and show interpenetration, the saussurite is cloudy to turbid and rarely contains original plagioclase. From this it is clear that although the Onib gabbros are altered, they preserved their primary textures; this indicates greenschist to lower amphibolite-facies metamorphism. Such mineralogical and textural criteria are indicators of in-situ metamorphism (i.e., hydrothermal transformation of minerals taking place within the oceanic crust). This was also reported from neighbouring ophiolites, e.g., Wadi Ghadir, SE Desert of Egypt (El Bayoumi, 1980) and Gerf some 80 km to the north of Onib (Zimmer, 1989).

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Chapter 5: The Wadi Onib Mafic-Ultramafic Complex

7.5. Leucogabbro The bulk of the Onib gabbroic mass is made up of well layered, predominantly leucogabbroic varieties. Microscopic investigations revealed a typical cumulate nature. The rock shows considerable variations in texture and/or mineral assemblage, though the simple paragenesis saussurite/uralite is retained almost everywhere. However, the average rock from the layered plagioclase-rich leucogabbro consists of more than 60% saussuritic material, about 30% uralitic amphibole, and the rest of the mode comprises the usual ink-blue zoisite, sericite, albite patches, clear specks of quartz (?), carbonate aggregates, iron oxide rods and blebs as well as epidote clusters and apatite needles in decreasing abundance. Plagioclase is strongly altered into ink-blue zoisite and retains some twinning. Its An-content is about 60%. 7.6. Isotropic Gabbro Coalescing occurrences of isotropic gabbro, hornblende-gabbro and/or gabbro-diorite occur mainly to the west and southwest of the Sudi well (Fig. 3B). The rocks are dark to greenish-grey, massive, fine- to medium grained, essentially holocrystalline (hypodiomorphic to allotriomorphic granular) and show less deformation than the layered gabbroic cumulates. The average isotropic gabbro consists essentially of somewhat equal proportions (∼ 40%) of uralitic amphibole and saussurite with subordinate amounts of remnant clinopyroxene. Chlorite, zoisite, quartz, apatite and iron granules constitute accessory phases. The appearance of quartz and apatite may be due to metamorphism of plagioclase or may be explained as reflecting crystallization (as interstitial material) during late stages of magmatic activity when temperatures were significantly lower than those maintained during formation of the cumulates. The isotropic gabbro is the only gabbroic assemblage in the Onib area that contains up to 7% Qtz and between 3 and 26% Ab in its CIPW normative composition. Its An is conversely the lowest (∼ 37%) and Hy is ∼ 15%, similar to the other gabbroic facies analyzed (Hussein, 2000). The normative features are not far from the modal mineral characteristics, though alteration is ubiquitous. In general terms, the isotropic gabbro shows a shift towards a sodic plagioclase-bearing gabbroic/dioritic association that is frequently encountered along the marginal zones of the uppermost gabbroic sections of many ophiolitic complexes. 7.7. Plagiogranite Plagiogranite or trondhjemite (low colour index quartz-plagioclase rocks) is frequently found as irregular bodies within, or marginal to, the gabbroic suites of ophiolitic complexes. These rocks may occur as small dykes and veinlets amongst high-level gabbrodioritic complexes or as wedges within sheeted dykes. Plagiogranite is a late-stage differentiate of qtz-normative basaltic magma (Coleman and Peterman, 1975). Chemically, an ideal plagiogranite is impoverished in K2 O (< 1%) and Rb (< 5 ppm) but high in Sr when

7. Petrography of the Onib Ophiolite Sequence

185

compared to continental granophyres. Because plagiogranites are end members of the differentiation products of the ophiolitic suite they are enriched in SiO2 , Na2 O, FeO, and TiO2 . Pearce et al. (1984a) asserted that plagiogranites of all tectonic subgroups (normal ocean ridges, anomalous ocean ridges, back-arc basin ridges and supra-subduction zone ridges) have hornblende as the dominant ferromagnesian mineral, plot as quartz-diorite or tonalite in the Streckeisen (1976) diagram and may be metaluminous or peraluminous in character. In the Wadi Onib ophiolite, plagiogranites can easily be confused in the field with postophiolite arc-related batholithic granitoid occurrences. However, the batholithic granitoids are biotite and/or hornblende-bearing; they have larger dimensions and are, more or less, mafic dyke-intruded. They also have sharp contacts with the country rocks. The plagiogranite, on the other hand, contains little or no biotite and occurs as dyke-like and/or irregular bodies within the gabbroic domain. In addition, staining and petrographic studies of a number of thin sections from suspected plagiogranites have revealed no or only traces of alkali feldspar. Under the microscope the Onib plagiogranite is holocrystalline, allotriomorphic granular and consists essentially of plagioclase, quartz, hornblende and uralite. In addition, there is chlorite, epidote, sericite, zoisite, apatite and iron ore granules that occur in trace abundance. Plagioclase appears to be albite to oligoclase in composition. Modal albitic plagioclase is clearly reflected by some 40% of normative Albite (Hussein, 2000). Single zircons from a sample of plagiogranite collected ca. 5 km west-northwest of the Sudi well yielded a magmatic crystallisation age of 808 ± 14 Ma (Kröner et al., 1992). 7.8. Sheeted Dykes Petrographic studies and interpretations of mineral parageneses and fabrics of the Onib inter-ophiolite dykes indicate that they are extensively deformed and recrystallized through hydrothermal ocean-floor alteration as well as through regional metamorphism that accompanied ophiolite obduction and orogeny (Hussein, 2000). The ophiolite-related dykes are tholeiitic (normative hypersthene-bearing) and texturally vary between basalt, dolerite and microgabbro. Sheeted dykes are found west of Sudi well and within the pillow lavas which are tectonically juxtaposed against the basal ultramafic unit in uppermost Wadi Onib. Both occurrences have comparable mineral parageneses and textural relations. The Sudi dykes and Onib counterparts are greyish-green in colour, medium to fine-grained in texture and predominantly consist of saussuritized, partly zoned tabular to fragmental plagioclase crystals with subordinate actinolite/tremolite. These are associated with trace amounts of remnant pyroxene and ubiquitous alteration products, mainly ink-blue zoisite that accompanies cloudy, cryptocrystalline saussuritic material, chlorite, epidote, carbonate and albite. The sheeted dykes, though altered and deformed, still preserve subophitic textures (with plagioclase crystals providing the framework and pyroxene crystals and derivatives building the interstitial pore-filling assemblages). However, the dykes of Wadi Sudi have less zoisite than those from uppermost Wadi Onib. Whereas the former have about 19% normative hypersthene the latter, in general, contain less normative hypersthene. Additionally, the Sudi sheeted dykes are more deformed than those in Wadi Onib and show tendencies of upper

186

Chapter 5: The Wadi Onib Mafic-Ultramafic Complex

greenschist/lower amphibolite-facies grade of metamorphism. This is thought to be an indication of the different levels of injection of the dykes; the Sudi dykes are believed to be stratigraphically lower than those of the Onib area. This inference is also corroborated by field criteria (e.g., in uppermost Wadi Onib the dykes inject pillow lavas whereas at Sudi they mainly inject microgabbros). 7.9. Pillow Lavas The Wadi Onib lavas have been extensively deformed and metamorphosed. Though the mineral assemblages are broadly similar, there are variable textures. Thin sections from margins of individual pillows, for example, show cryptocrystalline greenish chlorite which has probably replaced a hyaloclastic matrix. On the other hand, thin sections from the inner portions of the usually subrounded to elongated pillows still display some aphyric to microporphyritic textures, in places forming subophitic intergrowths between plagioclase laths and/or phenocrysts and partially to completely uralitized pyroxene flakes and blades. In one section, abundant uralitic amphibole alternates with fragmental, cloudy plagioclase crystals and all form sub-parallel structures (a sort of metamorphic foliation). Another rock shows more plagioclase than amphibole. However, the plagioclase is highly charged with zoisite and, to a lesser extent, with sericite. These secondary minerals form roughly subparallel bands. The latter may be indications of residual flow structures rather than a metamorphic foliation. The predominance of plagioclase and alteration products in some of the Onib lavas may indicate a fractionation phenomenon (fractionation leads to differentiated melts). Nonetheless, the majority of the lavas show pillow structures, textures and mineralogies of submarine tholeiites. 7.10. Sequence of Crystallization and Metamorphism Field and microscopic observations as well as structural interpretations enable tentative inferences pertaining to the possible course of mineral crystallization within the magma chamber(s) generating the Onib ophiolite. Granularity, crystallinity, grain shape and size as well as fabric relations of mineral phases that constitute the plutonic components reveal a broad zonation whereby the ophiolite became more fractionated in the west and east respectively. Olivine, chrome spinel, pyroxene and calcic plagioclase are the most abundant phases of the deeper sections of the plutonic cumulate mass, mainly in the basal peridotite and the adjacent TZ. Higher up in the magmatic sequence, the layered gabbroic rocks become enriched in plagioclase. Finally, at the highest levels, some isotropic gabbros bear sodic plagioclase, primary amphibole and minor amounts of quartz. The latter becomes an essential component of the differentiation end product, the plagiogranite. Preserved textures in the peridotite/gabbro component point to cumulate processes. Oikocrysts of olivine and orthopyroxene swimming within serpentine meshes and/or blades, euhedral to subhedral chrome spinel grains engulfed by or engulfing pyroxene crystals at places (poikilitic relations), the presence of two generations of pyroxene and olivine (large and small), the binding of plagioclase tabular crystals by intercumulus pyroxenes

8. Geochemistry of the Wadi Onib Ophiolite

187

(subophitic to ophitic textural fabric) and many other microstructural criteria suggest that a complex evolution took place before the cumulate pile was built up. The exceptionally thick TZ contains abundant clinopyroxene. Both olivine and orthopyroxene are thought to have been trapped early in the basal peridotite. This inference is corroborated by the fact that olivine is not distinguished as a component of the gabbroic rocks which are found to be essentially pyroxene and/or plagioclase-rich with noritic and anorthositic affinities at places, the latter being adcumulate. However, due to the advanced stage of alteration in the cumulate component (serpentinization in the peridotite and saussuritization and/or uralitization of the gabbros) it was difficult to establish a single order of crystallization, though combinations of olivine-clinopyroxene-orthopyroxene-plagioclase, olivine-clinopyroxeneplagioclase-orthopyroxene and/or olivine-orthopyroxene-clinopyroxene-plagioclase sequences are thought to have been maintained. It is probable that a parental magma was generated through partial melting of mantle peridotite beneath a spreading ridge and evolved towards a basaltic and, in places, even andesitic composition through fractional crystallization. It has produced large quantities of olivine, spinel, clinopyroxene and calcic plagioclase within the TZ. Further fractionation has given rise to high-level gabbroic/dioritic emplacements and, finally, to the uppermost intrusive/extrusive carapace of dykes and pillow lavas. Laurent et al. (1980) asserted that the composition of a parental magma and its changes during differentiation are the main factors controlling the composition of the cumulus minerals and the order of crystallization. Furthermore, detailed petrographic studies indicate that the Onib ophiolite was subjected to varying degrees of metamorphic overprinting, partly due to hydrothermal circulation. Criteria that enable us to make such an inference are: (1) vertical zonation of metamorphic grades and termination of metamorphic effects with depth (namely the pyroxenites and/or wehrlites of the TZ are the best preserved amongst the cumulate rocks), (2) albitization and saussuritization of plagioclase, (3) uralitization of pyroxene, (4) development of lizarditic serpentine (possibly early phases), (5) alteration of ilmenitic accessories into sphene and magnetite into haematite, (6) preservation of igneous textures although the primary minerals are more or less replaced (e.g., zoisite is well developed and retains twin lamellae of precursor plagioclase phases), and (7) rarity of schistosity and linear fabric within the gabbros, though saussuritized and/or uralitized.

8. GEOCHEMISTRY OF THE WADI ONIB OPHIOLITE Whole-rock and mineral data for the Onib ophiolitic sequence are presented in Tables 1–3 and Figs. 7–11. Due to the cumulate nature of the Onib plutonic sequence, emphasis is placed on the geochemistry of the volcanic suite. Although the latter was subjected to sea-floor hydrothermal alteration and/or subsequent regional metamorphism most major and trace elements still reflect the composition of the Onib parental magmas (Hussein, 2000). Major element abundances, variations and ratios are used to establish broad-scale chemical groupings, whereas trace elements are used to highlight igneous processes and pertinent palaeotectonic environments.

ONH18 Gabbro 48.21 0.18 24.57 1.08 3.4 0.06 5.7 14.86 2.42 < 0.01 0.05 0.65 0.49 0.09 0.16

ONH41 Gabbro 45.65 0.08 20.71 1.48 3.88 0.07 10.66 15.59 0.66 < 0.01 0.01 1.89 1.88 0.11 0.01

Total

100.28

101.27

100.79

Rb Sr Nb Zr Y V Co Cr Ni Cu Zn

3.7 Ga Isua (Greenland) supracrustal sequence. Geology 24, 43– 46. Rudnick, R.L., 1995. Making continental crust. Nature 378, 571–578. Sajona, F.G., Maury, R.C., Bellon, H., Cotton, J., Defant, M.J., Pubellier, M., 1993. Initiation of subduction and generation of slab melts in western Mindanao, Philippines. Geology 21, 1007– 1010. Sajona, F.G., Maury, R., Bellon, H., Cotton, J., Defant, M., 1996. High field strength element enrichment of Pliocene-Pleistocene island arc basalts, Zamboanga Peninsula, Western Mindanao (Philippines). Journal of Petrology 37, 693–726. Schiano, P., Clocchiatti, R., Shimizu, N., Maury, R.C., Jochum, K.P., Hofmann, A.W., 1995. Hydrous, silicate-rich melts in the sub-arc mantle and their relationship with erupted lavas. Nature 377, 595–600. Shchipansky, A.A., Samsonov, A.V., Bibikova, E.V., Babarina, I.I., Konilov, A.N., Krylov, K.A., Slabunov, A.I., Bogina, M.M., 2004. 2.8 Ga boninite-hosting partial suprasubduction zone ophiolite sequences from the North Karelian greenstone belt, NE Baltic Shield, Russia. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 425–486. Smellie, J.L., Stone, P., Evans, J., 1995. Petrogenesis of boninites in the Ordovician Ballantrae Complex ophiolite, southwestern Scotland. Journal of Volcanology and Geothermal Research 69, 323– 342. Smithies, R.H., Champion, D.C., 2000. The Archean high-Mg diorite suite: Links to tonalitetrondhjemite-granodiorite magmatism and implications for early Archean crustal growth. Journal of Petrology 41, 1653–1671. Stern, R.J., Morris, J., Bloomer, S.H., Hawkins, J.W., 1991. The source of the subduction component in convergent margin magmas: trace element and radiogenic evidence from Eocene boninites, Mariana forearc. Geochimica et Cosmochimica Acta 55, 1467–1481. Stern, R.A., Syme, E.C., Bailes, A.H., Lucas, S.B., 1995. Paleoproterozoic (19–186 Ga) arc volcanism in the Flin Flon Belt, Trans-Hudson Orogen, Canada. Contributions to Mineralogy and Petrology 119, 117–141. Stern, C.R., Kilian, R., 1996. Role of the subducted slab, mantle wedge and continental crust in the generation of adakites from the Andean Austral Volcanic Zone. Contributions to Mineralogy and Petrology 123, 263–281. Stein, M., Hofmann, A.W., 1994. Mantle plumes and episodic crustal growth. Nature 372, 63–68.

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Sun, S.-s., Nesbitt, R.W., 1978. Geochemical regularities and genetic significance of ophiolitic basalts. Geology 6, 689–693. Sun, S.-s., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the Ocean Basins. Geological Society of London Special Publication 42, 313–345. Tatsumi, Y., Maruyama, S., 1989. Boninites and high-Mg andesites: tectonic and petrogenesis. In: Crawford, A.J. (Ed.), Boninites and Related Rocks. Unwin Hyman, London, pp. 50–71. Taylor, R.N., Nesbitt, R.W., 1988. Light rare earth enrichment of supra subduction-zone mantle: evidence from the Troodos ophiolite, Cyprus. Geology 16, 448–451. Taylor, R.N., Nesbitt, R.W., Vidal, P., Harmon, R., Auvray, B., Croudace, I.W., 1994. Mineralogy, chemistry, and genesis of the boninite series volcanics, Chichijima, Bonin Islands, Japan. Journal of Petrology 35, 577–617. Taylor, S.R., McLennan, S.M., 1995. The geochemical evolution of the continental crust. Reviews of Geophysics 33, 241–265. Teklay, M., Berhe, K., Reimold, W.U., Armstrong, R., Asmerom, Y., Watson, J., 2002. Geochemistry and geochronology of a Neoproterozoic low-K tholeiite-boninite association in Central Eritrea. Gondwana Research 5, 597–611. Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L., Sage, R.P., 1991. Wawa Subprovince. In: Thurston, P.C., Williams, H.R., Sutcliffe, H.R., Stott, G.M. (Eds.), Geology of Ontario. Ontario Geological Survey Special Volume 4 (1), 485–539. Wolde, B., Asres, Z., Desta, Z., Gonzalez, J., 1996. Neoproterozoic zirconium depleted boninite and tholeiite series rocks from Adola, southern Ethiopia. Precambrian Research 80, 261–279. Wyman, D.A., Kerrich, R., 1993. Archean shoshonitic lamprophyres of the Abitibi subprovince, Canada: Petrogenesis and tectonic setting. Journal of Petrology 34, 1067–1109. Wyman, D.A., 1999a. Paleoproterozoic boninites in an ophiolite-like setting, Trans-Hudson orogen, Canada. Geology 27, 455–458. Wyman, D.A., 1999b. A 2.7 Ga depleted tholeiite suite: evidence of plume-arc interaction in the Abitibi greenstone belt, Canada. Precambrian Research 97, 27–42. Wyman, D.A., Kerrich, R., Polat, A., 2002. Assembly of Archean cratonic mantle lithosphere and crust: plume-arc interaction in the Abitibi-Wawa subduction-accretion complex. Precambrian Research 115, 37–62. Xie, Q., Kerrich, R., Fan, J., 1993. HFSE/REE fractionations recorded in three komatiite-basalt sequences, Archean Abitibi greenstone belt: implications for plume sources and depths. Geochimica et Cosmochimica Acta 57, 4111–4118. Yogodzinski, G.M., Kay, R.W., Volynets, O.N., Koloskov, A.V., Kay, S.M., 1995. Magnesian andesite in the western Aleutian Komandorsky region: Implications for slab melting and processes in the mantle wedge. Geological Society of America Bulletin 107, 505–519.

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Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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Chapter 18

ARCHEAN GREENSTONE BELTS DO CONTAIN FRAGMENTS OF OPHIOLITES MAARTEN J. DE WIT CIGCES, Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa

1. INTRODUCTION Most Archean greenstone belts are severely tectonised so that reconstruction of their rock assemblages revealing original autochthonous relationships is a daunting task (de Wit and Ashwal, 1997; Kusky and Vearncombe, 1997). There are about 260 individual Archean greenstone belts worldwide. Of these about 40 (∼ 15%) have been studied in sufficient detail (and mapped at a scale of less than 1:100.000) to provide relatively reliable information about pre-2.5 Ga geological processes (de Wit and Ashwal, 1995, 1997). Greenstone belts represent some of the earliest records of Earth history, but they are not restricted to the Archean. For example, the large Neoproterozoic Arabian-Nubian shield has an Archean-like cratonic crust with at least 7 major greenstone belts, most of which comprise island arc-like successions and associated (but often dismembered) ophioliteassemblages (Berhe, 1997). Similarly, the Baltic shield contains greenstone belt sequences ranging in age from > 3.1 Ga (Mesoarchean) to 1.9 Ga (Mesoproterozoic). Some of the Mesoproterozoic greenstone belts share characteristics of many Archean greenstone belts (e.g., abundant komatiites), whilst others share characteristics of Phanerozoic ophiolites (Sorjonen-Ward et al., 1997; and this volume). There is no simple definition of a greenstone belt other than that they contain significant volumes of basaltic rocks metamorphosed at relatively low grades to yield “green” mineral assemblages (de Wit and Ashwal, 1995). A wide spectrum of tectonic environments is preserved within Archean greenstone belts, and many individual belts are mixtures of components from different tectonic environments and in particular from island arc terrains (de Wit and Ashwal, 1995, 1997; Kusky and Vearncombe, 1997). It is claimed nevertheless by some that oceanic crust-forming environments are not preserved amongst this mixture of tectonic regimes because in their views no rocks assemblages in Archean greenstone belt sequences exhibit sufficient features to warrant definitive classification as an ophiolite (Bickle et al., 1994; Hamilton, 1998). The difficulty in recognizing and even defining ophiolites has been acknowledged widely and is not addressed here (Anonymous, 1972; and this volume). In this short contribution I outline some probable and some possible ophiolite sequences that have been reported from a number of Archean greenstone belts DOI: 10.1016/S0166-2635(04)13018-1

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Chapter 18: Archean Greenstone Belts Do Contain Fragments of Ophiolites

Fig. 1. Global map of Archean greenstone belts from which ophiolite-like sequences have been reported and described. See Table 1 for references.

around the world (Fig. 1 and Table 1). I also comment on the likely tectonic implications of these examples to better resolve Archean processes. A vast number of geochemical analyses indicate that oceanic-like (MORB-type) basalts and alpine-type peridotites occur in a large number of the studied greenstone belts. To access this, the reader is referred to Sylvester et al. (1997), O’Hanley (1997), and Arndt et al. (1997) for overviews.

2. EVIDENCE FROM GEOLOGY 2.1. Ophiolite Complexes and Slivers in Mélanges of the 2.55–2.05 Ga Central Orogenic Belt of the North China Craton Most recently the first laterally extensive end-Archean ophiolite complex (the 50 × 5 km2 Dongwanzi ophiolite, 2505 Ma) has been documented in the North China Craton (Kusky et al., 2001; and this volume). The complex is surrounded by a mélange of ophiolitic fragments, ranging from deformed mantle peridotite with podiform chromitites to pillow lavas, set in sheared metasediments and amphibolites (Li et al., 2002, 2004). The Dongwanzi complex is one of a series of tectonic blocks with ophiolitic affinities that outcrop within the Zunhua tectonic zone of the 1600 km-long Central Orogenic Belt that is flanked to the east by the Qinglong foreland fold-and thrust-belt. The northern section of the Central Orogenic Belt also hosts the Quingjang greenstone belt (Yuehua and Wang, 1997).

Greenstone Belt /Craton Kalgoorlie Greenstone Belt, Yilgarn Craton, Kanowna Lake and Lake Cowan localities

Reference

Distinguishing features

Age

Alternative interpretations

Neoarchean 2675–2715 Ma

Proposed tectonic setting Extensional oceanic environment

Fripp and Jones, 1997

Basaltic-komatiitic sheeted dykes, sills with ophiolitic-type peridotite-gabbroic plutonic sequence. Pillow lavas, cherts. Contacts between lithologies are tectonic. Ductile magmatic fabrics in gabbros

Yellowknife Greenstone Belt, Slave Craton

Helmsteadt et al., 1986

Mafic sheeted dykes, microgabbros with contacts exposed. No ultra-mafics, no komatiites. MORB-like pillow sequences. Lower Kam Group (Chan Fm) locally resembles ophiolite sheeted duke complex. No komatiites Ophiolite allochthons, pillows lavas, dykes, gabbros, ultramafics

Neoarchean 2716–2686 Ma

Back-arc basin or proto-oceanic rift

Bickle et al., 1994

Point Lake Belt

Kusky, 1991

Neoarchean

Imbricated forearc thrust over continental crust

King and Helmsteadt, 1997

Cameron River Belt, Slave Craton

Kusky, 1990

Mafic pillow lavas, dyke-gabbro transition, sill complexes

Neoarchean

Possible oceanic crust of back-arc basin

2. Evidence from Geology

Table 1. Some Archean Greenstone Belts with ophiolite-like mafic-ultramafic assemblages

King and Helmsteadt, 1997; Corcoran et al., 2004 (continued on next page)

601

602

Table 1. (Continued) Reference

Distinguishing features

Age

Proposed tectonic setting Fore-arc supra-subduction zone or slivers of oceanic crust now in accretionary prism

Kusky et al., 2001; and this volume; Li et al., 2002; Yuehua and Wang, 1997

Complete “Penrose-type” ophiolite metamorphosed at amphibolite facies. Includes pillow lavas, sheeted dykes, gabbros and ultra-mafic complexes and harzburgitic tectonites. Part of the Central Orogenic Belt that includes the Shenyang greenstone belt and other ophiolite complexes. No komatiites

Neoarchean 2505 Ma

Cleaverville Greenstone Belt, NW Pilbara Craton

Otha et al., 1996; Barley, 1997

Series of duplexes of pillow basalts and diabases with MORB geochemistry overlain by cherts

3.1 Ga

Tectonic slides of upper oceanic crust

South Pass Greenstone Belt Wind River Range Wyoming Craton

Harper, 1985; Wilks and Harper, 1997

Metamorphosed pillow lavas, diabases, gabbros and ultra-mafics. Highly strained and disrupted by shear zones. Possible dismembered ophiolite

Mesoarchean 2800–2630 Ga

Possible oceanic crust of a back-arc basin

Alternative interpretations Zhai et al., 2002

(continued on next page)

Chapter 18: Archean Greenstone Belts Do Contain Fragments of Ophiolites

Greenstone Belt /Craton Dongwanzi Complex Zunkuan mélange of the Central Orogenic Belt North China Craton

Greenstone Belt /Craton Barberton Greenstone Belt Kaapvaal Craton

Reference

Distinguishing features

Age

Proposed tectonic setting 1. Fore-arc supra-subduction complex 2. Oceanic basin— back-arc basin 3. Fore-arc supra-subduction zone with small mélanges

de Wit et al., 1987; Brandl and de Wit, 1997; Grove et al., 1997; Dann, 2001

3 ophiolitic allochthons of different ages of which 2 host komatiites. Pillow lavas and sheet flows, sheeted intrusions (mostly sills), peridotites, wehrlites, dunites, chromitites. Overlain by cherts and turbidites

Mesoarchean 1. 3480–3470 Ma 2. 3440–3340 Ma 3. 3300–3220 Ma

Pietersburg Greenstone Belt Kaapvaal Craton

de Wit et al., 1992; Brandl and de Wit, 1997

Pillow lavas, diabases, gabbros, ultra-mafics. Occasional komatiites. Sequence is tectonically disrupted into a number of Neoarchean thrust slices

Mesoarchean to Neoarchean 3450–2700 Ma

Oceanic back-arc basin, now in accretionary wedge and foreland basin

Isua Supracrustal belt, Greenland North Atlantic Craton

Maruyama et al., 1991; Myers, 2001

Pillow lavas, gabbros, ultramafics, BIF turbidites, felsic volcanics, calc-silicates. Highly deformed and metamorphosed

Paleoarchean

Oceanic crustal allochthons in subduction-related accretion complex

Alternative interpretations Lowe and Byerly, 1999; Dann, 2001

2. Evidence from Geology

Table 1. (Continued)

Nutman, 1997

(continued on next page)

603

604

Table 1. (Continued) Reference

Distinguishing features

Age

Proposed tectonic setting Immature oceanic arc or forearc supra-subduction zone ophiolite

Shchipansky et al., 2004

Boninites, pillow lavas, sheeted dikes, gabbro

2.8 Ga

Olondo greenstone belt, Aldan Shield

Puchtel, 2004

Komatiitic and tholeiitic pillow basalts, sheeted sill complex, gabbro, cumulate ultramafic rocks

3.0 Ga

Supra-subduction zone ophiolite

Wutai Shan, North China craton

Li et al., 2004

Metabasalts, chert, black smoker chimneys, gabbro

2.5 Ga

Forearc ophiolite fragment

Alternative interpretations

Chapter 18: Archean Greenstone Belts Do Contain Fragments of Ophiolites

Greenstone Belt /Craton North Karelian terrane, Baltic Shield

2. Evidence from Geology

605

Thus, this orogenic belt contains mafic-ultramafic remnants that have traditionally been documented as an Archean greenstone belt, as well as tectonic slivers of ophiolites. From all accounts this is the first regional tectonic mélange zone comparable to modern accretionary prisms and trench complexes, until now assumed to be absent from the Archean (e.g., Hamilton, 1998). Although some of the field interpretations have been challenged (Zhai et al., 2002) there is little doubt that the components of a true ophiolite sequence (cf. Anonymous, 1972) are present over a great aerial extent of the Central Orogenic Belt and that now require mapping at the kind of scales (< 1:15,000) that will allow more detailed comparisons with modern ophiolites and possibly modern oceanic crust. 2.2. Sheeted Dykes of the Yellowknife Greenstone Belt, Slave Craton The Kam Group of the circa 2.7 Ga Yellowknife greenstone belt is probably best known for its extensive sheeted dyke complex that displays a gradual transition onto overlying tholeiitic pillow lavas and grades downward into isotropic fine-grained metagabbros (Helmsteadt et al., 1986; King and Helmsteadt, 1997; Bickle et al., 1994). The base of the sequence is not clearly observed, but likely consists of a shear zone at the top of the underlaying BIF/rhyolites and quartzites (e.g., the Dwyer Formation; personal observations; Kusky, 1987). This suggests that the upper part of an ophiolite sequence is exposed here. The sequence has many characteristics of the Mesozoic Rocas Verdes ophiolites (that also lack an exposed ultramafic base) in the southern Andes, which were formed in a back arc environment, and initially emplaced into the granitic roots of an active volcanic arc (Stern and de Wit, 2003). Other possible slices of dismembered ophiolitic sequences have been suggested to occur in at least two other greenstone belts of the Slave Craton (e.g., the Cameron River and the Point Lake Greenstone Belts; Kusky, 1990, 1991; King and Helmsteadt, 1997; Corcoran et al., 2004). 2.3. Sheeted Dykes and Gabbro-Peridotite Sequence of the Yilgarn Craton Fripp and Jones (1997) propose an ophiolitic origin for some of the greenstones of the Kalgoorlie terrane. Their detailed mapping (at scales of 1:1000 and 1:2500) in two wellexposed lake section of the southern part of the Kalgoorlie greenstone belt has revealed a complex of sheeted dykes and sills of both high-Mg and possibly dunitic (komatiite) compositions, as well as an extensive ophiolite-like peridotite to gabbroic plutonic sequence with igneous ductile-deformation fabrics. One cliff section on Island 2 of Lake Cowan displays a continuous 25 m-wide section of sheeted intrusive units with both one-way and two-way chilling. The ophiolitic units are believed to be part of regional greenstone assemblages throughout the Kalgoorlie terrane. Evidence for an earlier granitic basement on which these greenstones may have been deposited (and which may have been responsible for inherited continental chemical and mineralogical contamination of the mafic rocks) is thought to be unlikely. Seismic reflection data suggests that the greenstones are allochthonous. This in turn suggests that small contamination (of for example zircons) may have been derived from assimilation of associated turbidites. Fripp and Jones (1997, p. 435)

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Chapter 18: Archean Greenstone Belts Do Contain Fragments of Ophiolites

suggest that many of the basaltic and komatiitic rocks of the generally poorly exposed Kalgoorlie (and surrounding) terranes, may represent more sheeted intrusions than is recognized at present, and that many of the so-called ultramafic layered complexes may comprise components of Archean oceanic lower crust and the transition zone to the upper mantle. These suggestions certainly warrant a re-evaluation of the geology of many greenstone belts within the Yilgarn Craton. 2.4. Ophiolite Slivers of the Barberton Greenstone Belt, Kaapvaal Craton The Barberton greenstone belt comprises at least 3 major allochthonous thrust sheets with predominantly mafic-ultramafic rock assemblages (Fig. 2; Brandl and de Wit, 1997; de Wit et al., in preparation; but see Lowe, 1999, for a more autochthonous interpretation). Maficultramafic sections of the Barberton greenstone belt have been interpreted in the past as remnants of ophiolites (Annhaeusser et al., 1968) and in particular an ophiolite representative of oceanic or back-arc basin crust, named the Jamestown Ophiolite Complex (JOC; de Wit et al., 1987). New mapping and geochronology since then have cast significant doubts on this interpretation, in part because different components of the restored JOC sections are now known to comprise mafic-ultramafic sections of different ages. For example, some Alpine-type ultramafics (dunites and harzburgites, including dunitic tectonites) were restored as the lowest sequence of the JOC, despite the fact that they were separated from the upper sequence by a major shear zone (de Wit et al., 1987). Precise geochronology has since shown that these cumulates and tectonites belong to a maficultramafic sequence of a separate tectonic block (Kaapvalley allochthon, Fig. 2) flanking the NW margin of the greenstone belt that is ∼ 230 million years younger than the upper ophiolitic section with which they were previously linked (the Onverwacht allochthon, Fig. 2; de Wit et al., 1987; de Ronde and de Wit, 1994). In addition, the upper sequences of the JOC (that include the type Komati and Hooggenoeg Formations) do not contain a sheeted dyke section as was previously suggested by de Wit et al. (1987). Composite, and sometimes sheeted mafic-ultramafic intrusions, particularly massive tholeiites, wehrlites and pyroxenites are common in the lower parts of the section. de Wit et al. (1987) suggested that these were dykes, but they now appear to be mostly hypabyssal sills (Dann, 2000, 2001), although some extensive dykes are also present. The volume and relative age of the sills remains a matter of controversy (de Wit et al., 1987; Grove et al., 1997; Dann, 2000, 2001). Dann (2000, 2001) suggests that most of the sequence is extrusive and that the intrusives may be significantly younger than their host rocks. In contrast, de Wit et al. (1987) believe a significant proportion of the massive rocks are intrusives, near contemporaneous with the high-Mg basaltic pillow lavas. The chemistry of relict pyroxenes in some of the komatiites, when compared to those tested in experimental work, is compatible with such an intrusive interpretation (Parman et al., 1997, 2001, 2003; Parman and Grove, 2004). Recent geochemical and petrological work has shown that some of the associated komatiites have boninitic affinities and were probably generated during hydrous melting of depleted mantle in an arc-forearc environment (Grove et al., 1999; Parman et al., 2001, 2003; Grove and Parman, 2004; Parman and Grove, 2004). Thus the

2. Evidence from Geology

Fig. 2. Geological map of the Barberton Greenstone Belt, showing the 3 main allochthonous mafic-ultramafic sequences with ophiolite-like characteristics. Inset shows part of the Kromberg allochthon with the plutonic mafic-ultramafic sheeted sill intruding the mainly tholeiitic pillow lava-chert sequence. The sequence is about 2–3 km thick. 607

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Chapter 18: Archean Greenstone Belts Do Contain Fragments of Ophiolites

Fig. 3. Close up photo of part of the sheeted sills of the Kromberg Allochthon, showing one- and two-way chilling in tholeiitic intrusive units. Outcrop along the Komati River, located in Fig. 2.

lowermost mafic-ultramafic sequences of the Barberton greenstone belt probably comprise the upper ∼ 3 km of a 3.48–3.47 Ga supra-subduction complex, likely of ophiolitic nature. Overlying the supra-subduction complex (part of the Onverwacht allochthon) maficultramafic rocks comprise part of the Kromberg allochthon (Fig. 2). The type section of this sequence occurs along the Komatii River in the southeastern sector of the belt and the age of this sequence is bracketed between 3.45 and 3.36 Ga (de Ronde and de Wit, 1994; Lowe and Byerly, 1999). New mapping and geochemistry indicates that this sequence predominantly comprises a complex array of intrusive wehrlites, pyroxenites, dunites and massive tholeiites separating screens of tholeiitic pillow basalts and breccias, interbedded with minor cherts and BIF (Fig. 2 inset; de Ronde and de Wit, 1994; de Wit et al., in preparation). The pyroxenites and massive tholeiites are part of a vast intrusive sheeted-sill complex. Sheeted intrusions varying between several meters across to less than 10 cm thick, with one- and two-way chilled margins are well exposed midway along the section in the Komatii river (Figs. 2 and 3). These sheeted intrusions were initially also interpreted as sheeted dykes (de Wit et al., 1987) but the subsequent mapping suggests that they probably also represent sheeted sills. Neither the top nor the bottom of the mafic-ultramafic sequence is preserved in this section. Large dunite and peridotite massifs mostly with tectonic boundaries, have been interpreted as Alpine-type peridotite complexes and large thicknesses of tholeiitic pillow lavas are ubiq-

3. Evidence from Mantle Xenoliths

609

uitous throughout the sequence away from the type area (Barton, 1982; Paris, 1985; de Wit et al., 1987). Komatiites are relatively rare and the predominant chemistry of the extrusive/hypabyssal sequence is tholeiitic (de Wit et al., 1987; Lowe and Byerly, 1999). Thus, the Kromberg allochthon comprises a separate sequence that contains most of the components of an ophiolite (ss) including a tholeiitic sheeted complex albeit in the form of sills rather than dykes. In all, the Barberton greenstone belt contains at least three mafic-ultramafic allochthons each of which preserve the “Steinmann Trinity” and at least one of which conforms closer to the more rigorous definition of an ophiolite. Only continued mapping in this type Mesoarchean greenstone belt will reveal just how close. 3. EVIDENCE FROM MANTLE XENOLITHS A significant number of studies on Archean eclogite and peridotite xenoliths found in kimberlites and related rocks, and derived from mantle lithosphere underlying Archean cratons have been equated on geochemical grounds to subducted and metamorphosed oceanic crust. It is beyond the scope of this contribution to explore these findings further, but the interested reader should consult recent findings of Shirey et al. (2004) and references therein. Suffice it to state that the data from these xenoliths suggests that significant volumes of Archean accreted oceanic lithosphere and oceanic arcs are preserved in the mantle “keels” of most Archean cratons (Kusky, 1993; Saltzer et al., 2001; Foley et al., 2003). 4. DISCUSSION AND CONCLUSION Archean greenstone belts contain a large amount of tholeiitic pillow basalts (Hunter and Stowe, 1997; de Wit and Ashwal, 1995). Comparable volumes on the contemporary Earth are confined to its oceans and their arcs, plateaux and islands. “The world would have had to be astonishingly different in the Archean for such giant accumulations of pillow lavas to have formed on continents” (Burke, 1997). Studies of ophiolites have helped to resolve important questions about Phanerozoic oceanic crust. But the structure of Archean oceanic crust remains elusive. Archean ophiolites are rarely recognized for what they are, possibly because they do not necessarily fit a standard definition and/or expected model. Kusky et al. (2001) resolve that the restored crustal section of the Dongwanzi ophiolite (∼ 10 km) is substantially thicker than that of Phanerozoic ophiolite sequences. In contrast, de Wit et al. (1987) claimed that the Archean Jamestown ophiolite complex appeared to represent a relatively thin crustal section. Unresolved tectonic repetition does not allow a robust reconstruction of the various Barberton ophiolitic allochthons, but the most reliable restoration reveals a thickness of around 6–8 km. Both estimates appear to be in violation of simple models that predict much thicker oceanic crust on an Archean Earth with higher mantle temperatures, perhaps as thick as below Iceland (∼ 25 km) where a plume has elevated the mantle temperatures beneath the mid Atlantic ridge by some 200 ◦ C, approaching the

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Chapter 18: Archean Greenstone Belts Do Contain Fragments of Ophiolites

sort of temperatures expected of the upper Archean mantle (Bickle et al., 1994; Pollack, 1997; Grove and Parman, 2004). Yet spreading across Iceland is slow in contrast to the predicted fast spreading rates in Archean oceans (Abbott and Hofmann, 1984; Pollack, 1997; de Wit, 1998; Karson, 2001). It is not clear from modern oceanic environments what sort of crustal structure might evolve from voluminous magmatism at very fast spreading rates (Karson, 2001) and/or lower viscosities perhaps due to higher water content of Archean mantle (cf. Grove et al., 1997, 1999; Parman et al., 2001, 2003). Features specific to maficultramafic sequences of Archean greenstone belts should help resolve this. A most obvious difference between Paleoproterozoic-Recent ophiolites and that of the mafic-ultramafic crustal section of the Barberton greenstone belt sequence (and that of most Archean greenstone belt sequences) is the scarceness of gabbroic rocks in the latter (de Wit et al., 1987; de Wit and Ashwal, 1997). Gabbros reflect the transient storage of mantle melts before they transfer their evolved magma to form upper oceanic crust. The absence of significant volumes of gabbros indicates that Archean mafic mantle melts spend little time in magma chambers to construct a middle section of oceanic crust. This complements the generally high-Mg content of many greenstone belt basalts. High-Mg melts rarely bypass the magma storage chambers of Phanerozoic ophiolites and present oceanic crust (Elton, 1979). It is plausible that these characteristics of Archean basalts represent more rapid extensional processes of Archean tectonic environments. A similar suggestion has been put forward for the near-ubiquitous occurrence of the enigmatic ocelli (or variolites) in the mafic-ultramafic rocks of greenstone belts. These features are only rarely described from Phanerozoic ophiolites or oceanic crust, and must be providing an important message about Archean mantle melting and crystallization (de Wit and Ashwal, 1997). One suggestion is that they represent “frozen-in” differentiation processes (de Wit et al., 1987). This interpretation provides additional evidence that intermediate stages of magma storage in the form of gabbroic plutons may have been less frequent in the Archean than in the crustal framework of present oceans and island arcs. How these observations may be incorporated into models of Archean oceanic crust formation remains to be resolved. What is clear, however, is that Archean ophiolite-like sequences do occur in some Archean greenstone belts and their restored internal structures must be explored to further test and challenge theoretical models that imply oceanic crust is not preserved in Archean greenstone belts. Since at least 80% of known Archean greenstone belts have not been studied in sufficient details to partake in the discussions, its back to field-work we should go.

ACKNOWLEDGEMENTS Over the last 25 years a number of geologists familiar with Phanerozoic ophiolite sequences have visited the Barberton greenstone belt with me and provided constructive feedback on my interpretations of the rock sequences there as being ophiolite-like. I am grateful to them for sharing their critical views, in particular: C. Stern, R.A. Hart, W.S.F. Kidd, A.G. Smith, H. Helmsteadt, T.L. Grove, J.C. Dann, S. Bowring, A. Hynes, M. Searle,

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H. Furnes, H. Staudigal, N. Banerjee, K. Meuhlenbachs, B. Robbins (in chronological order of the visits). Although some were more convinced than others, I believe they all agreed that the sequences could well be ophiolitic, although this interpretation of their opinions remains mine. Over these years, funds to study this greenstone belt were provided through the FRD (now NRF), as well as a number of exploration companies, which is gratefully acknowledged.

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olite sequences from the North Karelian greenstone belt, NE Baltic Shield, Russia. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 425–486. Shirey, S.B., Richardson, S.H., Harris, J.W., 2004. Ages, parageneses, and compositions of diamonds and evolution of the Precambrian mantle lithosphere of southern Africa. South African Journal of Geology 107, 119–130. Sorjonen-Ward, P., Nironen, M., Luukkonen, E.J., 1997. The Baltic Shield: Greenstone associations in Finland. In: de Wit, M.J., Ashwal, L.D. (Eds.), Greenstone Belts. Oxford Univ. Press, UK, pp. 677–698. Stern, C.R., de Wit, M.J., 2003. The Rocas Verdes Ophiolites, southernmost South America: remnants of progressive stages of development of ocean-type crust in a continental margin back-arc basin. In: Dilek, Y., Robinson, P.T. (Eds.), Ophiolites in Earth History. Geological Society of London Special Publication 218. Sylvester, P.J., Harper, G.D., Byerly, G.R., Thurston, P.C., 1997. Volcanic aspects. In: de Wit, M.J., Ashwal, L.D. (Eds.), Greenstone Belts. Oxford Univ. Press, UK, pp. 55–90. Wilks, M.E., Harper, G.D., 1997. Wind River Range, Wyoming Craton. In: de Wit, M.J., Ashwal, L.D. (Eds.), Greenstone Belts. Oxford Univ. Press, UK, pp. 508–516. Yuehua, Y., Wang, W., 1997. North China Craton. In: de Wit, M.J., Ashwal, L.D. (Eds.), Greenstone Belts. Oxford Univ. Press, UK, pp. 730–735. Zhai, M., Zhao, G., Zhang, Q., 2002. Is the Dongwanzi Complex an Archean ophiolite? http:// www.sciencemag.org/cgi/content/full/295/5557/923a.

Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Published by Elsevier B.V.

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Chapter 19

NORTHERN PHILIPPINE OPHIOLITES: MODERN ANALOGUES TO PRECAMBRIAN OPHIOLITES? JOHN ENCARNACIÓN Department of Earth and Atmospheric Sciences, Saint Louis University, Saint Louis, MO 63103, USA

The northern Philippines is a possible modern analogue for some Precambrian greenstone belts. It has a ∼ 150 Myr history of multiple and overlapping periods of oceanic crust generation, arc volcanism, sedimentation, and deformation dominated by wrench tectonics. At least five ophiolite complexes of distinct ages make up most of the basement—all having a distinct suprasubduction zone signature. Arc plutons are predominantly of the diorite-tonalite series with minor alkali-feldspar bearing rocks. Sedimentary basins probably floored by oceanic crust are dominated by immature sediments and volcaniclastics and are locally up to ∼ 10 km thick. The whole arc and ophiolitic complex is in the process of being accreted to Eurasia, where it may be preserved in a broad “suture zone” between Eurasia and Australia and/or the Americas. 1. INTRODUCTION The Philippine islands constitute a mature island arc complex comprised mainly of ophiolites, island arc plutons, volcanics and volcaniclastics, and thick sedimentary basins filled with immature sediments. The term “island arc complex” is used because what is now the Philippines is a composite of more than one arc magmatic belt due to a long history of subduction whose polarity has probably changed more than once along the western and eastern side of the islands (in the current reference frame). However, the available evidence does not support the presence of any suture within the Philippines east of the Manila-Negros trench collision, at least in the northern part of the islands. Hence, the various arc volcanoplutonic belts on the islands were probably not juxtaposed by arc collisions, but rather were generated above the same suprasubduction zone setting by inward subduction from the western side and eastern sides. These magmatic belts are built on a basement of oceanic crust (ophiolites) that is itself composite. Based on good biostratigraphic ages on overlying sedimentary rocks and modern dating techniques (zircon U-Pb and 40 Ar-39 Ar dating), it has been shown that there are at least five generations of ophiolitic basement in the northern Philippines spanning an age range of ∼ 150 Myr. Based on areas that have been mapped in more detail, it appears that the ophiolites are not allochthonous slices that have been tectonically amalgamated. A model consistent with the field relations is one where DOI: 10.1016/S0166-2635(04)13019-3

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Chapter 19: Northern Philippine Ophiolites: Modern Analogues to Precambrian Ophiolites?

Table 1. Name and label in Fig. 1 a Lagonoy ophiolite b

Calaguas Island ophiolite

c

Dibut Bay ophiolite

d

Casiguran ophiolite

e

Montalban ophiolite

f

Zambales-Angat ophiolite

g

Itogon ophiolite

References David et al. (1997), Fernandez et al. (1994), Geary (1986), Karig (1983), Tamayo et al. (1998) Geary (1986), Geary et al. (1988), Geary and Kay (1989), Giese et al. (1986), Knittel (1989), Mitchell and Balce (1990) Billedo et al. (1996), Hashimoto et al. (1978), Tejada and Castillo (2002) Anonymous (1977), Billedo et al. (1995), Billedo et al. (1996), Tamayo et al. (2001) Arcilla (1991), Arcilla et al. (1989), Encarnación et al. (1993), Encarnación et al. (1999), Haeck (1987), Karig (1983) Abrajano et al. (1989), Arcilla (1991), Arcilla et al. (1989), Bachman et al. (1983), Encarnación et al. (1999), Encarnación et al. (1993), Evans (1985), Evans et al. (1991), Evans and Hawkins (1989), Florendo and Hawkins (1992), Fuller et al. (1991), Geary et al. (1989), Hawkins and Evans (1983), Karig (1983), Rossman et al. (1989), Schweller et al. (1984), Yumul (1996), Yumul et al. (1998) Anonymous (1977), Anonymous (1987), Balce et al. (1980), Bellon and Yumul (2000), Encarnación et al. (1993), Florendo (1994), Mitchell and Balce (1990)

List of ophiolites labeled on Fig. 1 and references pertaining to them.

each episode of oceanic crust generation occurs adjacent to or within older basement as forearc, backarc, or intraarc seafloor spreading type process. It has been pointed out that the Philippine island arc complex is “reminiscent of Precambrian greenstone belts” (e.g., Hall, 1996). In this paper, I outline some of the more salient features of the northern Philippines and its ophiolitic basement. I focus on the northern Philippines because it has been studied in more detail and is, therefore, more well-known. However, there are no known fundamental differences between the northern and southern Philippines, therefore the broad outline of the geology presented here is probably applicable to the south as well. For a more detailed discussion of each individual ophiolite terrane, the reader is referred to the original papers (see Table 1) and to a recent review paper focusing on the details of each individual ophiolite and their relationships to each other (Encarnación, in press). 2. REGIONAL TECTONIC SETTING The Philippines is located in Southeast Asia in the western Pacific at the juncture of the Philippine Sea plate and Eurasian plate (Rangin, 1991). It is a mature island arc that is in the process of being accreted to the Eurasian margin by subduction-related convergence along the east-dipping Manila and Negros trenches. Subduction along the western side of

2. Regional Tectonic Setting

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Fig. 1. Present tectonic setting of the Philippines. Rectangle outlines the location of Fig. 2. PF— Philippine Fault; MN—Mindanao; M—Mindoro; P—Panay; LZ—Luzon; NPB—North Palawan Block.

the northern Philippines is generating the Luzon volcanic arc, which consists of the stratovolcanoes in the southern Zambales range (where Mt. Pintubo is located) and volcanic centers running northward along the east side of the Central Cordillera to small islands north of Luzon and into Taiwan. The Luzon arc has collided with the Eurasian margin in Taiwan and in the central Philippines along the islands of Mindoro and Panay (Fig. 1) (McCabe et al., 1982). Presumably as a result of this collision, subduction may be waning on the west side of the Philippines and convergence between the Philippine Sea plate and the Eurasian plate may be increasingly taken up along the west dipping Philippine trench and East Luzon trough (Cardwell et al., 1980; Hamilton, 1979; Lewis and Hayes, 1983;

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Chapter 19: Northern Philippine Ophiolites: Modern Analogues to Precambrian Ophiolites?

Fig. 2. General geology of the northern Philippines. See Fig. 1 for location. Main ophiolite exposure: a—Lagonoy ophiolite; b—Calaguas Islands ophiolite; c—Dibut Bay ophiolite; d— Casiguran ophiolite; e—Montalban ophiolite; f—Zambales and Angat ophiolites; g—Itogon ophiolite. CTVB—Central Valley Basin; CAVB—Cagayan Valley Basin; MB—Marinduque Basin; CSC—Cordon syenite complex. Adapted from (Anonymous, 1963, 1977, 1981, 1991; Letouzey and Sage, 1988).

3. Major Geologic Elements of the Northern Philippines and Their Characteristics

619

Rangin, 1991). An active arc in southeast Luzon has formed due to subduction along the Philippine trench; no volcanic arc has yet developed from underthrusting at the East Luzon trough. The latter appears to be exploiting an old subduction zone because there is a mature accretionary prism and Eocene plutons and arc volcanics in the northern Sierra Madre as shown in Fig. 2 (Lewis and Hayes, 1983). Oblique northwest convergence of the Philippine Sea plate is being accommodated partly by sinistral wrench faulting along the Philippine Fault system (Aurelio et al., 1991). The Marinduque basin (Fig. 2) is a pull apart basin associated with wrench faulting and is floored by young oceanic crust (Sarewitz and Lewis, 1991). It formed by seafloor spreading type processes reflected in symmetric magnetic anomalies. It is a good actualistic example of formation of younger ophiolitic basement within older ophiolitic and arc basement.

3. MAJOR GEOLOGIC ELEMENTS OF THE NORTHERN PHILIPPINES AND THEIR CHARACTERISTICS Ophiolites are scattered throughout the northern Philippines as shown in Fig. 2. Many of them have been disrupted mainly by wrench faulting. All of the ophiolites variably preserve residual mantle peridotite (in the form of partly serpentinized harzburgite), gabbroic rocks and ultramafic cumulates, sheeted dikes or dike swarms, pillow lavas and pillow breccias. Modern dating by the zircon U-Pb and 40 Ar-39 Ar techniques coupled with traditional biostratigraphic work has shown that there are at least five generations of ophiolite preserved in the northern Philippines. The oldest is the Lagonoy ophiolite in southeast Luzon (Fig. 2), which has a minimum age of Jurassic (Geary et al., 1988). It comprises the oldest basement known in the Philippines east of the Manila trench-Negros trench convergent margin. It has a distinct suprasubduction zone signature and may be a primitive island arc. The Calaguas Islands, Dibut Bay, Casiguran, and Montalban ophiolites are constrained as pre-Late Cretaceous or Early Cretaceous (see Table 1 for references). Together they may comprise a large section of Early Cretaceous oceanic crust that may have formed as backarc basin lithosphere associated with the Jurassic Lagonoy ophiolite/primitive arc. An interesting relationship established by careful mapping and biostratigraphy in the Northern Sierra Madre is the occurrence of a sequence of pillow lavas of Late Cretaceous age with a distinct arc signature overlying the Early Cretaceous back-arc basin-like pillow lavas of the Casiguran ophiolite (Billedo et al., 1996) (Fig. 2). The Eocene Zambales ophiolite, in western Luzon (Fig. 2) has been the most intensely studied and is the largest exposed ophiolite in the Philippines. It has a distinct suprasubduction zone signature with boninites, MORB-like crust and transitional arc like crust (Hawkins and Evans, 1983). The Zambales ophiolite may be a part of a much larger Eocene ophiolitic basement that most likely extends eastward underneath the Central Valley of Luzon and into the Southern Sierra Madre where it crops out again as the Eocene Angat ophiolite. Fig. 3 is a cross section across central Luzon that shows the probable geometry of the basement and the known geology in the Zambales range and western side of the Southern Sierra Madre. The Eocene ophiolitic basement may also extend northward into

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Chapter 19: Northern Philippine Ophiolites: Modern Analogues to Precambrian Ophiolites?

Fig. 3. East-West section across Luzon from the Zambales Mountains, across the Central Valley Basin and into the Southern Sierra Madre (compare with Fig. 1). It is inferred that most of the section is underlain by an ophiolitic basement that crops out as the Zambales ophiolite in the west and the Angat ophiolite in the east. Section is based on data from several sources (Arcilla et al., 1989; Bachman et al., 1983; Encarnación et al., 1993; Haeck, 1987; Hawkins and Evans, 1983).

northern Luzon. Fig. 4 shows the inferred extend of the various ophiolitic basement in Luzon. A detailed discussion for the basis of this model is discussed elsewhere (Encarnación, in press). The Itogon ophiolite in the Central Cordillera is Oligocene in age and is thought to have been generated during intraarc rifting (Florendo, 1994). Oligocene and Miocene alkaline plutons and volcanics in the Cordon Syenite Complex (CSC, Fig. 2), are thought to have been related to the same intraarc rifting event. Some of the lavas and dikes associated with the Itogon complex are similar to magmatic rocks in the Sumisu Rift (Florendo, 1994). Geochemical data for the various ophiolites is of variable quality and coverage. As mentioned earlier, the Zambales ophiolite has been studied the most and has a wide variety of data available. Although clearly a single slab of oceanic lithosphere, it contains MORBlike, transitional arc tholeiite, and boninitic rocks typical of a suprasubduction zone setting. A detailed discussion of all the available geochemical data from all the ophiolites in the northern Philippines can be found in Encarnación (in press). Not surprisingly, all of the ophiolites have evidence for a suprasubduction zone origin (e.g., Hawkins and Florendo, 1992) primarily in the form of variable enrichment of large ion lithophile elements and depletion in high field strength elements. Overlying and intruding many of the ophiolites are arc volcanics/volcaniclastics and arc plutons, respectively (Figs. 2 and 5). Abundant Cretaceous arc volcanics and volcaniclastics of andesitic and dacitic composition are the first major manifestation of arc volcanism in the northern Philippines. Eocene and Oligocene volcanics are also widespread and abundant. Many of these arc rocks are metamorphosed to lower greenschist facies assemblages. They are largely submarine and along with the volcanic sections of the ophiolites have been mapped as “Cretaceous-Paleogene metavolcanics” in the older maps and literature. The large batholiths exposed in the Central Cordillera and Northern Sierra Madre and to

4. Analog for Archean and Proterozoic Systems?

621

Fig. 4. Possible extent of ophiolitic basement in the northern Philippines. Dashed lines—outline of northern Philippines. Solid lines—inferred location of contacts between ophiolites of various ages. Compare with Fig. 2. The extrapolation of ophiolitic basement beyond the main exposures shown in Fig. 2 is based on isolated smaller exposures, structural, isotopic, and seismic data (Encarnación, in press). The contacts between the ophiolites are probably “intrusive” (i.e., younger ophiolites form by extension adjacent to, or within older ophiolite basement), although possibly modified by later wrench faulting (Encarnación, in press; Encarnación et al., 1993; Karig, 1983; Karig et al., 1986). The dominant ophiolitic basement is intruded by arc plutons, partly covered by volcanics and volcaniclastics, and downfaulted or downwarped into deep sedimentary basins that are filled with sediments and volcaniclastics.

a lesser extent, the Southern Sierra Madre, are largely hornblende diorites, quartz diorites and tonalites. These plutons are poorly studied and few of them have any published chemical analyses. Although dominated by diorite-tonalite series plutons, minor syenites, granodiorites, monzonites and lamprophyres have also been described. Most of these more alkali rich rocks are found in the southern Central Cordillera.

4. ANALOG FOR ARCHEAN AND PROTEROZOIC SYSTEMS? The key characteristics of the northern Philippines that may be relevant to Precambrian ophiolites and greenstone belts are summarized in Fig. 5. Petrologically, the crust is dominated by ophiolitic lithologies that would be classified as “greenstones”. Much of the arc volcanics and volcaniclastics that overly the ophiolitic basement may be difficult to differentiate from the ophiolitic volcanic section and indeed they may certainly be transitional in space and time. In some areas, a significant portion of the crust is made up of

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Chapter 19: Northern Philippine Ophiolites: Modern Analogues to Precambrian Ophiolites?

Fig. 5. Cartoon illustrating some of the key features of northern Philippine ophiolites and related rocks. Figure corresponds roughly to a section from the Manila trench off of the southern Zambales Range northeast to the Baguio region and east to the East Luzon trough. Ages of oceanic crust are based on zircon U-Pb ages and biostratigraphic ages (Billedo et al., 1996; Encarnación et al., 1993). Ages of intrusives are zircon U-Pb ages and K-Ar ages (Anonymous, 1977; Bellon and Yumul, 2000; Encarnación et al., 1993; Florendo, 1994; Wolfe, 1981). Note that the ∼ 70 Ma pillow lavas are unconformable on the ∼ 110 Ma ophiolite basement (Billedo et al., 1996).

diorite-tonalite intrusions as in many greenstone belts (e.g., de Wit and Ashwal, 1997). The thick ophiolite-floored sedimentary basins are also another key feature and because of their thickness, are likely to be preserved in the geologic record. The span of geologic activity preserved in the Philippines, covering 150 Myr is comparable to the span of activity recorded in some greenstone belts (de Wit and Ashwal, 1997; Kusky and Polat, 1999). Ophiolite generation, arc magmatism, and sedimentary deposition all spanned this ∼ 150 Myr period and occurred in various places in northern Luzon at various times. If this complex package were preserved in an ancient mountain belt, any attempt to understand the geology in terms of a “layer cake” stratigraphy would lead to erroneous conclusions. Interpretation of structural relationships of volcano-sedimentary belts in Archean greenstone terranes might benefit from a more realistic, albeit more complex, model based on modern analogues (e.g., Kusky and Vearncombe, 1997). In order to use the Philippine ophiolitic terranes as analogues for Precambrian greenstone belts, we need to project current plate motions and kinematics forward in time and predict what the geometry of the major units will be upon final accretion to the Eurasian margin. The Philippines has already partly collided with Eurasia along the Mindoro and Panay area. Convergence continues with eastward subduction of the South China Sea basin and Sulu Sea basin. Once these two small marginal basins are completely consumed by subduction, the Philippine arc complex will have been completely accreted to the Eurasian margin. The final stages of collision may likely be a bit more complex because westward subduction along the Philippine trench has started and the central part of the Philippines is already locked along the North Palawan Block margin. Hence, the final stages of collision

References

623

will probably require opening up of extensional basins, with possible formation of oceanic crust to accommodate heterogeneous strain. For example, Pubellier et al. (1996) suggested that the Philippines is undergoing ‘escape tectonics’ which allow the archipelago to deform. The formation of the oceanic Marinduque basin (Sarewitz and Lewis, 1991) is partly related to this intradocking deformation phase. As the westward subduction on the east side of the Philippines develops further, the northern Philippines will be intruded by yet another suite of plutons and will be blanketed by young arc volcanics. If closure of the Pacific ocean continues, the whole assemblage will be incorporated in a broad suture zone between Eurasia and Australia and/or the Americas. It will then be presumably overprinted by compressional structures and perhaps finally intruded by a post-orogenic intrusive suite. Its overall petrological, structural, and sedimentological characteristics would then be quite evocative of some Precambrian granite-greenstone belts.

REFERENCES Abrajano, T.A., Pasteris, J.D., Bacuta, G.C., 1989. Zambales ophiolite, Philippines: I. Geology and petrology of the critical zone of the Acoje massif. Tectonophysics 168, 65–100. Anonymous, 1963. Geological Map of the Philippines. Philippine Bureau of Mines, Manila. Anonymous, 1977. Report on geological survey of northeastern Luzon, consolidated report. Japan International Cooperation Agency-Metal Mining Agency of Japan, Tokyo, p. 106. Anonymous, 1981. Geology and Mineral Resources of the Philippines. Bureau of Mines and Geosciences, Ministry of Natural Resources, Manila, p. 406. Anonymous, 1987. Geology and mineralization in the Baguio area, northern Luzon. 5. United Nations Department of Technical Cooperation for Development, Manila. Anonymous, 1991. Report on the Mineral Exploration, Mineral Deposits and Tectonics of Two Contrasting Geologic Environments in the Republic of the Philippines, Terminal Report. Japan International Cooperation Agency, Metal Mining Agency of Japan, Mines & Geosciences Bureau, R.P., Tokyo. Arcilla, C.A., 1991. Lithologic, age, and structural study of the Angat Ophiolite, Luzon, Philippines. M.Sc. thesis. University of Illinois, Chicago, p. 107. Arcilla, C.A., Ruelo, H.B., Umbal, J., 1989. The Angat ophiolite, Luzon, Philippines: lithology, structure, and problems in age interpretation. Tectonophysics 168, 127–135. Aurelio, M.A., Barrier, E., Rangin, C., Müller, C., 1991. The Philippine Fault in the Late Cenozoic tectonic evolution of the Bondoc-Masbate-N. Leyte area, Central Philippines. Journal of Southeast Asian Earth Sciences 6, 221–238. Bachman, S.B., Lewis, S.D., Schweller, W.J., 1983. Evolution of a forearc basin, Luzon Central Valley, Philippines. The American Association of Petroleum Geologists Bulletin 67, 1143–1162. Balce, G.R., Encina, R.Y., Momongan, A., Lara, E., 1980. Geology of the Baguio region and its implications on the tectonic development of the Luzon Central Cordillera. Geology and Palaeontology of Southeast Asia 21, 265–287. Bellon, H.a.Y., Yumul Jr., G.P., 2000. Mio-Pliocene magmatism in the Baguio mining district (Luzon, Philippines): age clues to its geodynamic setting. Comptes Rendus de l’Academie des Sciences Serie II Fascicule A—Sciences de la Terre et des Planetes 331, 295–302.

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Billedo, E., et al., 1995. The geology and structures of San Ildefonso Peninsula, Casiguran, Aurora Province: Their relationship to the Northern Sierra Madre orogeny. Journal of the Geological Society of the Philippines 50, 37–60. Billedo, E., Stephan, J.F., Delteil, J., Bellon, H., Sajona, F., Feraud, G., 1996. The pre-Tertiary ophiolitic complex of northeastern Luzon and the Polillo group of islands. Philippines. Journal of the Geological Society of the Philippines 51, 95–114. Cardwell, R.K., Isacks, B.L., Karig, D.E., 1980. The spatial distribution of Earthquakes, focal mechanism solutions, and subducted lithosphere in the Philippine and northeastern Indonesian Islands. In: Hayes, D.E. (Ed.), The Tectonic and Geologic Evolution of Southeast Asian Seas and Islands. In: Geophysical Monographs, vol. 23. American Geophysical Union, Washington, DC, pp. 1–35. David, S.J., et al., 1997. Geology and tectonic history of Southeastern Luzon, Philippines. Journal of Asian Earth Sciences 15, 435–452. de Wit, M., Ashwal, L.D., 1997. Greenstone Belts. In: Oxford Monographs on Geology and Geophysics, vol. 35, p. 809. Encarnación, J., in press. Multiple ophiolite generation preserved in the northern Philippines and the growth of an island arc complex. Tectonophysics (special issue on “Ophiolites and Continental Margins of the Pacific Rim and Caribbean Region”). Encarnación, J., Mukasa, S.B., Evans, C., 1999. Subduction components and the generation of arclike melts in the Zambales ophiolite, Philippines: Pb, Sr, Nd isotopic constraints. Chemical Geology 156, 343–357. Encarnación, J.P., Mukasa, S.B., Obille, E.J., 1993. Zircon U-Pb geochronology of the Zambales and Angat ophiolites, Luzon, Philippines: Evidence for Eocene arc-back arc pair. Journal of Geophysical Research 98, 19,991–20,004. Evans, C.A., 1985. Magmatic ‘metasomatism’ in peridotites from the Zambales ophiolite. Geology 13, 166–169. Evans, C.A., Castaneda, G., Franco, H., 1991. Geochemical complexities preserved in the volcanic rocks of the Zambales ophiolite, Philippines. Journal of Geophysical Research 96, 16251–16262. Evans, C.A., Hawkins, J.W., 1989. Compositional heterogeneities in upper mantle peridotites from the Zambales range ophiolite, Luzon, Philippines. Tectonophysics 168, 23–41. Fernandez, M.V., Revilla, A.P., David, S.J., 1994. Notes on the Cretaceous carbonates in Catanduanes Island and Caramoan Peninsula. Journal of the Geological Society of the Philippines 49, 241–261. Florendo, F.F., 1994. Tertiary intra-arc rifting in the northern Luzon terrane, Philippines. Tectonics 13, 623–640. Florendo, F., Hawkins, J.W., 1992. Comparisons of the geochemistry of volcanic rocks of the Zambales ophiolite, northern Luzon, Philippines, implications for tectonic setting. Acta Geologica Taiwanica 30, 172–176. Fuller, M., Haston, R., Lin, J., Richter, B., Schmidtke, E., Almasco, J., 1991. Tertiary paleomagnetism of regions around the South China Sea. Journal of Southeast Asian Earth Sciences 6, 161–184. Geary, E.E., 1986. Tectonic significance of basement complexes and ophiolites in the northern Philippines: results of geological, geochronological and geochemical investigations. Ph.D. thesis. Cornell University, Ithaca, p. 221. Geary, E.E., Harrison, T.M., Heizler, M., 1988. Diverse ages and origins of basement complexes, Luzon, Philippines. Geology 16, 341–344. Geary, E.E., Kay, R.W., 1989. Identification of an Early Cretaceous ophiolite in the Camarines NorteCalaguas Islands basement complex, eastern Luzon, Philippines. Tectonophysics 168, 109–126.

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Geary, E.E., Kay, R.W., Reyonolds, J.C., Kay, S.M., 1989. Geochemistry of the mafic rocks from the Coto Block, Zambales ophiolite, Philippines, trace element evidence for two stages of crustal growth. Tectonophysics 168, 43–63. Giese, U., Knittel, U., Kramm, U., 1986. The Paracale Intrusion: geologic setting and petrogenesis of a trondhjemite intrusion in the Philippines island arc. Journal of Southeast Asian Earth Sciences 1, 235–245. Haeck, G.H., 1987. The geologic and tectonic history of the central portion of the Southern Sierra Madre, Luzon, Philippines. Ph.D. thesis. Cornell University, Ithaca. Hall, R., 1996. Reconstructing Cenozoic SE Asia. In: Hall, R., Blundell, D. (Eds.), Tectonic Evolution of Southeast Asia. Geological Society Special Publications 106, 153–184. Hamilton, W., 1979. Tectonics of the Indonesian Region. U.S. Geological Survey Professional Paper 1078. Hashimoto, W., Aoki, N., David, P.P., Balce, G.R., Alcantara, M., 1978. Discovery of Nummulites from the Lubingan crystalline schist exposed east of Bongabon, Nueva Ecija, Philippines and its significance on the geologic development of the Philippines. Geology and Palaeontology of Southeast Asia 19, 57–63. Hawkins, J.W., Evans, C.A., 1983. Geology of the Zamabales range, Luzon, Philippines Islands: ophiolite derived from an island arc-back arc basin pair. In: Hayes, D.E. (Ed.), Tectonics and Geologic Evolution of Southeast Asian Seas and Islands. American Geophysical Union Geophysical Monographs 27 (2), 95–123. Hawkins, J.W., Florendo, F., 1992. Supra-subduction zone magmatism: implications for the origin of Philippine ophiolites. Acta Geologica Taiwanica 30, 163–171. Karig, D.E., 1983. Accreted terranes in the northern part of the Philippine archipelago. Tectonics 2, 852–855. Karig, D.E., Sarewitz, D.R., Haeck, G.D., 1986. Role of strike-slip faulting in the evolution of allochthonous terranes in the Philippines. Geology 14, 852–855. Knittel, U., 1989. Comment on “Diverse ages and origins of basement complexes, Luzon, Philippines”. Geology 17, 669. Kusky, T., Polat, A., 1999. Growth of granite-greenstone terranes at convergent margins, and stabilization of Archean cratons. Tectonophysics 305, 43–73. Kusky, T., Vearncombe, J.R., 1997. In: de Wit, M., Ashwal, L.D. (Eds.), Greenstone Belts. In: Oxford Monographs on Geology and Geophysics, vol. 35, pp. 95–127. Letouzey, J., Sage, L., 1988. Geological and Structural Map of Eastern Asia. American Association of Petroleum Geologists, Tulsa. Lewis, S.D., Hayes, D.E., 1983. The tectonics of northward propagating subduction along eastern Luzon, Philippine islands. In: Hayes, D.E. (Ed.), The Tectonic and Geologic Evolution of Southeast Asian Seas and Islands. American Geophysical Union Geophysical Monograph 27 (2), 57–78. McCabe, R., Almasco, J., Diegor, W., 1982. Geologic and paleomagnetic evidence for a possible Miocene collision in western Panay, central Philippines. Geology 10, 325–329. Mitchell, A.H.G., Balce, G.R., 1990. Geological features of some epithermal gold systems, Philippines. Journal of Geochemical Exploration 35, 241–296. Pubellier, M., Quebral, R., Aurelio, M., Rangin, C., 1996. Docking and post-docking escape tectonics in the southern Philippines: Tectonic Evolution of Southeast Asia. In: Hall, R., Blundell, D. (Eds.), Tectonic Evolution of Southeast Asia. Geological Society of London Special Publication 6, 511– 523. Rangin, C., 1991. The Philippine mobile belt: a complex plate boundary. Journal of Southeast Asian Earth Sciences 6, 209–220.

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Rossman, D.L., Castaneda, G.C., Bacuta, G.C., 1989. Geology of the Zambales ophiolite, Luzon, Philippines. Tectonophysics 168, 1–22. Sarewitz, D.R., Lewis, S.D., 1991. The Marinduque intra-arc basin, Philippines: Basin genesis and in situ ophiolite development in a strike-slip setting. Geological Society of America Bulletin 103, 187–203. Schweller, W.J., Roth, P.H., Karig, D.E., Bachman, S.B., 1984. Sedimentation history and biostratigraphy of ophiolite-related Tertiary sediments, Luzon, Philippines. Geological Society of America Bulletin 95, 1333–1342. Tamayo, A.R.J., et al., 2001. Preliminary geochemical and mineral data from the Isabela-Aurora ophiolite, northeastern Luzon, Philippines. InterRidge News 10 (2), 50–53. Tamayo, A.R.J., Yumul, G.P.J., Santos, R.A., Jumawan, F., Rodolfo, K.S., 1998. Petrology and mineral chemistry of a back-arc upper mantle suite: example from the Camarines Norte ophiolite complex, south Luzon. Journal of the Geological Society of the Philppines 53 (1–2), 1–23. Tejada, M.L.G., Castillo, P.R., 2002. In search of common ground: Geochemical study of ancient oceanic crust in eastern Philippines. In: Goldschmidt 2002. Cambridge Publications, Davos, Switzerland, p. A767. Wolfe, J.A., 1981. Philippine geochronology. Journal of the Geological Society of the Philippines 20, 1–30. Yumul, G.P.J., 1996. Varying mantle sources of supra-subduction zone ophiolites: REE evidence from the Zambales ophiolite complex, Luzon, Philippines. Tectonophysics 262, 243–262. Yumul, G.P.J., Dimalanta, C.B., Faustino, D.V., de Jesus, J.V., 1998. Upper mantle-lower crust dikes of the Zambales ophiolite complex (Philippines): distinct short-lived, subduction-related magmatism. Journal of Volcanology and Geothermal Research 84, 287–309.

Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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THE RESURRECTION PENINSULA OPHIOLITE, MÉLANGE AND ACCRETED FLYSCH BELTS OF SOUTHERN ALASKA AS AN ANALOG FOR TRENCH-FOREARC SYSTEMS IN PRECAMBRIAN OROGENS TIMOTHY M. KUSKYa , ROSE GANLEYa, JENNIFER LYTWYNb AND ALI POLATc a Department

of Earth and Atmospheric Sciences, Saint Louis University, St. Louis, MO 63103, USA b Department of Geosciences, University of Houston, Houston, TX 77058, USA c Department of Geology, University of Windsor, Ontario, Canada

Southern Alaska’s Mesozoic-Cenozoic Chugach-Prince William terrane is an unusual forearc in that it contains belts of graywacke-dominated flysch, mélange, and ophiolitic fragments all intruded by a suite of tonalite-trondhjemite-granodiorite plutons, and large parts of the accretionary prism are metamorphosed to the greenschist, amphibolite, or granulite facies. The overall structural geometry, abundance and types of rocks and rock suites present, the petrogenetic relationships between rock suites, and the metamorphic style are all strongly reminiscent of Archean granite-greenstone terranes. As such, the southern Alaska forearc represents one of the world’s best modern analogs to early stages in the evolution of Archean granite-greenstone terranes. In this contribution, we examine the regional geology of the flysch, mélange, and accreted ophiolites, as well as details of the geology of the 57 ± 1 Ma Resurrection Peninsula ophiolite of southern Alaska’s Chugach terrane as a remarkable analog to some Archean greenstone belts. The Resurrection ophiolite formed in a near-trench environment as the Kula-Farallon ridge was being subducted beneath North America. The magmatic sequence includes pillow lavas, sheeted dikes, gabbros, trondhjemites, and a poorly-exposed ultramafic section. The lavas show mid-ocean ridge basalt and arc-like geochemical signatures, interpreted to reflect compositionally diverse melts derived from near-fractional melting of a variably depleted mantle source, mixed with variable amounts of assimilated continentally-derived flysch. A sedimentary sequence overlying the ophiolite preserves a continuous record of turbidite sedimentation deposited on the ophiolite as it was transported to North America and emplaced in the Chugach accretionary prism. The top of the sedimentary section is truncated by the Fox Island shear zone, a 1-km thick, greenschist-facies, west-over-east thrust related to the emplacement of the ophiolite into the accretionary wedge. The Fox Island shear zone is intruded by a 53.4 ± 0.9 Ma granite, showing that the ophiolite formed, was transported to DOI: 10.1016/S0166-2635(04)13020-X

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the North American continent, overthrust by a major accretionary prism-related thrust, and intruded by granite all within 3.6 ± 1.4 Ma. Geological relationships in the southern Alaska forearc are instructive, in that if similar relationships were found in an Archean granite-greenstone terrane, they would probably currently be interpreted to reflect calc-alkaline mafic-felsic volcanic-plutonic complexes intruded and erupted through a complex metasedimentary sequence. As such, the belt would probably be interpreted as an arc sequence. Many Precambrian forearc ophiolites and accretionary prisms may have gone unrecognized because the processes of forearc ophiolite emplacement and intrusion by near-trench magmas at triple junctions has been poorly documented. The near absence of subduction mélanges in Archean terranes has been used to argue that accretionary wedges are not present in Archean granite-greenstone terranes. Models for Archean plate tectonics predict that the planet may have had a greater plate boundary length, with more, smaller, faster moving plates than at present. One consequence of such a model is that subducting plates would have tended to be thickly sedimented, with their proximity to active collisions along numerous plate boundaries. Subduction of these thickly sedimented plates would tend to produce Archean accretionary wedges dominated by relatively coherent terranes and few mélanges, consistent with the geological record preserved in Archean granite greenstone terranes. 1. INTRODUCTION AND REGIONAL GEOLOGY Southern Alaska is composed of a series of belts of Paleozoic, Mesozoic, and Cenozoic age (Fig. 1a) (Plafker et al., 1994). The Wrangellia composite terrane north of the Border Ranges fault system (Fig. 1a) includes intraoceanic Paleozoic and Mesozoic magmaticarc assemblages of the Peninsular, Wrangellia, and Alexander terranes, locally overlain by thick carbonate successions (Plafker et al., 1989, 1994). The Southern Margin composite terrane to the south is an accretionary complex formed as a result of late Mesozoic and Cenozoic subduction and offscraping of sediments derived largely from the Wrangellia terranes to the north (Plafker et al., 1994). Included in this composite terrane are the Chugach, Prince William, and Yakutat terranes of the Prince William Sound area, the Ghost Rocks Formation on Kodiak Island, and Chugach metamorphic complex in the eastern Chugach Mountains (Fig. 1). The Aleutian thrust fault system in the south separates the Southern Margin composite terrane from the Pacific plate (Plafker et al., 1994). The Chugach-Prince William terrane (Fig. 1a) is a complexly deformed Mesozoic and Cenozoic accretionary prism (Plafker et al., 1977; Tysdal et al., 1977; Tysdal and Case, 1979; Nelson et al., 1985; Bradley et al., 1994, 1999a (with marginal notes, http://wrgis.wr.usgs.gov/open-file/of99-18/); Pavlis and Sisson, 1995; Kusky et al., 1997a, 1997b; Kusky and Bradley, 1999) that is up to 100 km wide along the Gulf of Alaska and extends along strike for over 2000 km along the Gulf of Alaska continental margin from Baranof Island in southeastern Alaska to the Sanak and Shumagin Islands in the southwest. The inboard part of the Chugach accretionary prism, the McHugh Complex (Fig. 1b), is a Permian-Cretaceous mélange, whereas the further outboard

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Fig. 1. (a) Map of southern Alaska showing locations of the Chugach, Prince William, and Wrangellian terranes and other tectonic elements discussed in text. (b) Map of the Kenai Peninsula showing locations of the Resurrection Peninsula and Knight Island ophiolites.

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part is a belt of Upper Cretaceous deformed flysch of the Valdez Group representing deep-sea fan turbidites (Fig. 1b) (Nelson and Nelson, 1993; Bradley and Kusky, 1992; Plafker et al., 1994; Kusky et al., 1997a; Kusky and Bradley, 1999). Outboard of the Valdez Group is a belt of somewhat less deformed Paleocene to Eocene flysch of the Orca Group of the Prince William terrane (Fig. 1b) (Plafker, 1969; Tysdal and Case, 1979; Jones et al., 1981; Plafker et al., 1985a, 1985b; Nelson et al., 1987). The geology of the Gulf of Alaska region includes the two most outboard units of the accretionary wedge deposits exposed above sea level, the Valdez and Orca Groups. These units are similar in lithology as both are composed of graywacke, siltstone, and shale, and local conglomerates and volcanic rocks. The dominant depositional mode was by gravity flow as turbidity currents moving westward along the trench axis (Winkler, 1976; Nilsen and Zuffa, 1982), although some east and northeastward directed paleocurrents have been documented as well (Plafker et al., 1994; Kusky et al., 1997b; D. Bradley, personal communication, 1999a, 1999b). In general, both the Valdez and Orca Groups become richer in quartz, potassium feldspar, and plagioclase, and poorer in volcanic lithic fragments to the east along the margin (Dumoulin, 1987). These rocks are interpreted to represent Late Cretaceous to Paleocene slope, fan, basin plain, and trench-fill turbidites (Nilsen and Zuffa, 1982) or deep-sea sedimentary fans later transported during accretion (Clendenin, 1991; Plafker et al., 1994). Both the Valdez and Orca Groups are variably folded with slatey cleavage developed in the finer grained rocks (Tysdal et al., 1977). Metamorphism from zeolite to low-greenschist facies is widespread but upper greenschist to amphibolite facies occurs locally near shear zones and as contact metamorphism around intrusions (Tysdal and Case, 1979; Moore et al., 1983; Plafker et al., 1989; Sisson et al., 1989; Bol and Gibbons, 1992). The Contact fault (Fig. 1b) is mapped as a major terrane boundary separating the Late Cretaceous rocks of the Valdez Group from Paleocene to Eocene rocks of the Orca Group (Tysdal and Case, 1979; Plafker et al., 1985a, 1985b). In eastern Prince William Sound the Contact fault separates poly-deformed, strongly foliated, middle greenschist facies rocks of the Valdez Group from less-deformed and metamorphosed rocks of the Orca Group. However, in the western Gulf of Alaska, the location of the boundary is obscured by the absence of an abrupt change in composition and metamorphic grade (Dumoulin, 1987, 1988; Gilbert et al., 1992). Sparse fossil control on ages along a belt approximately 60 km wide through the region add further ambiguity in locating a position for the Contact fault (Jones and Clark, 1973; Tysdal et al., 1977; Plafker et al., 1985a, 1985b; Nelson et al., 1985; Bol and Roeske, 1993; Kusky and Young, 1999). Several partial and dismembered ophiolite sequences outcrop in the western Gulf of Alaska and Prince William Sound, the most notable of which are the Knight Island (Fig. 2) and Resurrection Peninsula ophiolites (Fig. 3). The Paleogene Resurrection and Knight Island ophiolites are relatively undeformed and are the most complete ophiolite sequences in Alaska (Nelson et al., 1987; Plafker et al., 1994). The Resurrection Peninsula ophiolite is located on the Resurrection Peninsula south of Seward, and the Knight Island ophiolite is located 85 km east of Seward on Knight Island (Fig. 1b). Gravity data (Case et al., 1966; Saltus et al., 1999 (ftp://greenwood.cr.usgs.gov/pubs/open-file-reports/ofr-97-0520/

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Fig. 2. Geologic map of the Resurrection Peninsula ophiolite and overlying sedimentary sequence. Modified after Nelson et al. (1987), Tysdal and Case (1979), and the authors’ mapping.

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Fig. 3. Geologic map of the Knight Island ophiolite (after Nelson et al., 1985) showing the main ophiolitic units and sample locations.

data) suggest that the exposures of ophiolites in the region may be continuous below sea level from south of Knight Island to Glacier Island northward (Fig. 1b), and then several tens of kilometers east, and extend to approximately 10 km depth (Crowe et al., 1992; Nelson and Nelson, 1993) (Fig. 1b). Minor ultramafic rocks occur as pods and lenses of harzburgite and dunite in the mafic sections of the ophiolites (Tysdal et al., 1977; Tysdal and Case, 1979; Nelson et al., 1987; Bol et al., 1992; Crowe et al., 1992; Barker et al., 1992; Bradley and Kusky, 1992; Nelson and Nelson, 1993; Plafker et al., 1994). Both ophiolites are similar chemically and petrographically (Tysdal et al., 1977; Nelson and Nelson, 1993; Lytwyn et al., 1997). Other outcrop belts of mafic rocks in

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the Valdez and Orca Groups occur as tectonic slices of pillow lavas, sheeted dikes, gabbros, and interbedded tuffaceous units within the flysch, but lack exposures of welldeveloped layered gabbro and basal peridotite sections characteristic of deeper levels of full ophiolite sections (Nelson et al., 1987). Pillow basalt, minor diabase, and gabbro form 15–20 percent of the exposed rocks of the Orca group, which has been metamorphosed to zeolite facies (with some local greenschist facies) during emplacement, compaction, and dewatering of the accretionary prism (Goldfarb et al., 1986; Barker et al., 1992).

2. THE SANAK-BARANOF MAGMATIC BELT Plutons of the Sanak-Baranof belt intrude a complexly deformed, Mesozoic and Cenozoic accretionary prism—the Chugach-Prince William composite terrane—along the seaward margin of the Peninsular-Wrangellia-Alexander composite terrane (Fig. 1). The inboard part of the prism is a mélange of variably metamorphosed basalt, chert, argillite, graywacke, plus minor limestone and ultramafic rocks (Triassic to mid-Cretaceous McHugh Complex and equivalents). Farther outboard are belts of strongly deformed Upper Cretaceous and Lower Tertiary flysch, assigned to the Valdez and Orca Groups, respectively. The Orca Group includes several belts of mafic and ultramafic rocks that have been interpreted as having formed at a spreading ridge just before ridge subduction (Bol et al., 1992). Penetrative deformation in the accretionary prism (thrust imbrication, folding, mélange formation) and regional metamorphism (typically prehnite-pumpelleyite to greenschist facies) occurred during and shortly after offscraping and underplating during the Cretaceous and Early Tertiary. Near-trench plutons were emplaced into the already deformed accretionary prism. Paleocene to Eocene plutons of the Sanak-Baranof belt outcrop discontinuously along the entire 2200 km length of the Chugach-Prince William terrane (Bradley et al., 2003; Kusky et al., 2003; Sisson et al., 2003). The plutons are mainly granodiorite, granite, and tonalite (Hudson, 1983). Some of the plutons are elongate parallel to structural grain of the accretionary prism, although they are discordant at their ends. Some are enormous—the Kodiak batholith, for example, is approximately 1650 square kilometers in area. Smaller intermediate to silicic dikes are plentiful in some regions, such as the Anchorage and Seldovia quadrangles (Winkler et al., 1984; Winkler, 1992; Bradley and Kusky, 1992). They crosscut everything but a few late, brittle faults in the accretionary prism.

3. FOREARC OPHIOLITES OF THE SOUTHERN ALASKA CONVERGENT MARGIN The Resurrection Peninsula and Knight Island ophiolites, structurally emplaced into the forearc accretionary prism of the Chugach and Prince William terranes of southern Alaska,

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are interpreted as remnants of the Kula-Farallon ridge that was subducted beneath the North American continental margin in the Early Tertiary (Bol et al., 1992; Lytwyn et al., 1997; Kusky and Young, 1999). Subduction of the Kula-Farallon ridge and large poleward migration of the Chugach-Prince William composite terrane are fundamental events in the interpretation of the history of plate motions across the northeast Pacific Basin for the Tertiary (Hillhouse and Gromme, 1977; Byrne, 1979; Stone et al., 1982; Plumley et al., 1983; Hillhouse et al., 1985; Engebretson et al., 1985; Stock and Molnar, 1988; Lonsdale, 1988; Atwater, 1989). The formation and emplacement of the Resurrection Peninsula ophiolite may serve as a model for other ophiolites emplaced in forearc trench-ridge-trench settings of all ages, including the Precambrian, when ridge-trench encounters were probably more common than at present. Additionally, ophiolites emplaced in this and similar triple junction settings are some of the few examples of ophiolitic crust that was clearly formed at an oceanic spreading center, and not in arc- or forearc-related environments. Ophiolite emplacement in the forearc is not the only consequence of Tertairy ridge subduction in southern Alaska; other effects of ridge subduction include near-trench magmatism (Marshak and Karig, 1977; Moore et al., 1983; Bradley et al., 1993, 1999b, 2003; Harris et al., 1996; Lytwyn et al., 1997; Kusky et al., 2003), anomalous forearc deformation (Kusky et al., 1997a), high-temperature, low-pressure metamorphism (Sisson and Pavlis, 1993; Pavlis and Sisson, 1995), gold mineralization (Bradley et al., 1994; Haeussler et al., 2003) and a cessation of arc magmatism (e.g., DeLong and Fox, 1977; Bradley et al., 2003). Many of the above events are diachronous along strike. This is highlighted especially well by dating of the near-trench magmatic rocks, which show a west to east age progression from 61 to 50 Ma, consistent with emplacement during passage of the Kula-Farallon ridge as it migrated beneath the southern Alaska margin (Bradley et al., 1993, 1999b; Kusky et al., 1997a). The 57 ± 1 Ma (Nelson et al., 1989) age of the Resurrection Peninsula ophiolite is broadly coeval with this suite of Eocene near-trench granitoids in the Gulf of Alaska region, suggesting that it too is related to passage of the Kula-Farallon-North America triple junction. The close association of the Resurrection Peninsula ophiolite to these events indicates the importance of understanding the process of its emplacement as it relates to the development of the Chugach/Prince William accretionary wedge, and consequently, to our understanding of ridge-trench encounters in general. This understanding may find general applicability for many Precambrian forearc ophiolites. Paleomagnetic data from mafic rocks of the Resurrection Peninsula ophiolite suggest that the ophiolite and the composite Chugach-Prince William terrane, into which it was emplaced soon after formation, originated along the Wrangellian margin of North America, far to the south of their present position, perhaps as far as the present latitude of northern Washington (Bol et al., 1992; Bradley et al., 1993). Assimilation of flysch-like sediments by volcanics and dikes, in addition to turbidites interbedded with pillow basalts, indicate that the Resurrection Peninsula ophiolite formed in close proximity to the trench axis where it was soon subducted (Lytwyn et al., 1997). During its transport history, the Resurrection Peninsula ophiolite thus passed through a series of changing geologic environments related

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to both migration down the ridge flank and away from the ridge as well as its increasing proximity to the North American continent. This study addresses general questions related to ridge-trench encounters and ophiolite emplacement with time, using the Resurrection ophiolite as an analog for some Precambrian ophiolites emplaced at convergent margins. Other issues including increased volcanism in sediments related to injection of mid-ocean ridge basalts (MORB) at the base of the prism, and the emplacement of MORB/melted sediment hybrid magmas in structures associated with ophiolite emplacement are also described, and their implications for studies of Precambrian ophiolites are discussed. 3.1. Geology of the Resurrection Peninsula Ophiolite The Resurrection Peninsula ophiolite (Fig. 2) consists of a west-dipping sequence of sedimentary rocks, pillow basalts, sheeted dikes, and massive and layered gabbro (Fig. 4). The ophiolite is metamorphosed to greenschist facies (Nelson et al., 1989; Lytwyn et al., 1997; Kusky and Young, 1999). Tysdal et al. (1977) mapped the Resurrection Peninsula ophiolite as intruding the Valdez Group flysch and inferred a Late Cretaceous age for the ophiolite based on a few widely scattered fossils from the Valdez Group. Nelson et al. (1989) reported a U/Pb (zircon) age of 57 ± 1 Ma for a plagiogranite that intrudes the dike complex and is also cut by dikes (Fig. 4) and thus serves to date active spreading of the ophiolite during Early Eocene. The Resurrection Peninsula ophiolite is therefore contemporaneous with the Orca Group. Ultramafic rocks occur as pods of serpentinized peridotite and pyroxenite in the gabbroic section, and as tectonic blocks in flysch lying in fault contact with the ophiolite, the flysch having been thrust in a west-southwest direction over the ophiolite (Nelson et al., 1987; Bol et al., 1992). The mostly serpentinized ultramafic rocks of Resurrection Peninsula were originally clinopyroxenite, dunite, and peridotite (Nelson et al., 1987). The gabbroic rocks of the Resurrection Peninsula occur as a mafic pluton of at least 25 km2 in area that makes up the eastern side of the peninsula (Tysdal et al., 1977; Nelson et al., 1987, 1989). The structurally lowest part of the gabbro is a layered unit in the east containing locally well-developed, west-dipping magmatic mineral layering of alternating light and dark layers of pyroxene and feldspar. The layered unit is separated from a massive unit to the west by a block of Valdez flysch and interbedded volcanic and sedimentary rocks (Fig. 2). To the west, the massive gabbro contains an increasing percentage of dikes which further grade into the sheeted dike unit occupying the central section of the peninsula (Tysdal et al., 1977; Nelson et al., 1985). The dikes of the sheeted dike unit (Fig. 2) generally strike north along the central part of the peninsula, although the strike is more westward in the southern part of the peninsula (Tysdal et al., 1977; Nelson and Nelson, 1993; Bol et al., 1992). Most dikes are vertical, with some local variation and crosscutting of preexisting dikes at low angles (Fig. 4b). Dikes range from 0.1–5 m thick but average 0.3–1 m, increase in percentage upward, with both one-way and two-way chill margins observed on individual dikes in approximately equal proportions (Nelson et al., 1987; Bol et al., 1992). The dikes are variably aphanitic,

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Fig. 4. Photographs of: (a) deformed argillites and sandstone beds of the Humpy Cove Formation overlying the Resurrection Peninsula ophiolite, (b) basaltic pillow lavas of the Resurrection Peninsula ophiolite, (c) Sheeted dike complex from Killer Bay, (d) Mafic dike of sheeted complex cutting trondhjemite dated at 57 ± 1 Ma (Nelson et al., 1989).

porphyritic, and diabasic (Nelson et al., 1987). Both the gabbro and the sheeted dikes are intruded by 57 ± 1 Ma old plagiogranite (Nelson et al., 1987). To the west, a gradational zone 10 m thick separates the sheeted dikes and pillow basalts (Bol et al., 1992). Diabasic dikes exhibit intersertal, intergranular, and subophitic textures. The mineralogy is dominated by anhedral to euhedral plagioclase occurring as randomly oriented laths, some extensively embayed. Subophitic augite is common whereas pristine olivine is rare. Alteration minerals include (in order of decreasing abundance) chlorite (after olivine), actinolite, epidote (after plagioclase), and carbonate. Pillow basalts, minor amounts of massive basalt flows, and broken pillow breccias (Figs. 2, 3, and 4) form a unit 1000 m thick on the western flank of the peninsula (Tysdal et al., 1977; Nelson et al., 1987). Pillows average 0.5 m in diameter, strike north, and dip about 30◦ to 45◦ to the west, with interpillow spaces filled locally with red

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and green chert (Nelson et al., 1987). Pillow basalts of the Resurrection and Knight Island ophiolites contain interbedded terrigeneous flysch deposits (Fig. 4), suggesting continuous sedimentation during formation of the ophiolites at a source of high sediment input such as a continental margin (Tysdal et al., 1977; Nelson et al., 1987; Bol et al., 1992). Sheeted dike complexes containing chilled margins within the two ophiolites suites, however, suggest that the ophiolites formed in an extensional setting such as a spreading ridge, which indicates a close proximity of a spreading center with a continental margin (Tysdal et al., 1977; Coleman, 1977; Nelson et al., 1987; Nicholas, 1989; Nelson and Nelson, 1993; Bol et al., 1992). Pillow basalts include aphanitic, porphyritic, and hypocrystalline varieties. Subhedral to euhedral plagioclase phenocrysts are enclosed within a fine-grained quenched groundmass or devitrified glass containing abundant needle-like, randomly oriented plagioclase microlites. Some plagioclase phenocrysts are embayed and many crystals are normally zoned. Some show corroded, altered cores, but most are remarkably unaltered. Clinopyroxene occurs rarely as phenocrysts and more commonly as microphenocrysts in hypocrystalline basalts and large seriate grains in holocrystalline samples. Olivine is rare. Although several samples of pillow basalt appear virtually unaltered, others are slightly metamorphosed as evidenced by chlorite (after clinopyroxene) and/or amygdules filled with secondary chlorite, zeolites (natrolite?), and carbonate. Rare veins of prehnite also are found. Investigations of paleomagnetic data by Bol et al. (1992) on rocks of the Resurrection Peninsula ophiolite suggest that it (and presumably the Chugach terrane) was 13 ± 9◦ south of its present position with respect to North America, at 57 Ma when the ophiolite formed along the Kula-Farallon ridge. The 57 Ma age of the ophiolite, in conjunction with geologic relationships led Bol et al. (1992) to conclude that the Resurrection Peninsula ophiolite represents an accreted fragment of the Kula-Farallon ridge that was emplaced shortly after its formation. 3.2. The Knight Island Ophiolite The Knight Island Ophiolite (Figs. 1 and 3) is a fault-bounded complex within isoclinally folded flysch of the Orca Group. Pillow basalt and sheeted dikes, variably affected by greenschist-facies metamorphism, dominate the complex, while ultramafic rocks are found only as boulders along the beach and as xenoliths within the sheeted dike complex (Richter, 1965; Tysdal et al., 1977; Nelson et al., 1985, 1989; Bol et al., 1992; Crowe et al., 1992; Nelson and Nelson, 1993). The pillow and massive basalts form a unit over 5000 m thick with interpillow spaces filled with breccia, agglomerate, and minor red and green chert (Tysdal et al., 1977; Nelson and Nelson, 1993). Individual sheeted dikes up to several meters thick locally intrude Orca flysch, while small gabbroic plutons intrude both the dikes and adjacent sedimentary rocks (Tysdal et al., 1977). The alteration minerals found in pillow lavas and dikes are consistent with zeolite to lower greenschist-facies hydrothermal metamorphism recognized within the upper sections of both ophiolites (Wiltse, 1973; Tysdal et al., 1977; Nelson et al., 1989; Lytwyn et al., 1997).

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3.3. Valdez/Orca Metabasalts East of Resurrection Peninsula and Knight Island, a belt of metabasaltic and metadiabasic rocks (Fig. 1) is discontinuously exposed from Valdez to east of Yakutat, Alaska (Lull and Plafker, 1990). These metabasalts and metadiabases consist of pillow lavas, breccias, and dikes intruding and interbedded with mostly Valdez Group accretionary metasediments (Lull and Plafker, 1990) and were therefore generally regarded as Late Cretaceous (84– 66 Ma) in age. Recent discovery of geochemically similar rocks in the Orca Group, however, demonstrate that at least some of these metabasalts may be Eocene in age. We collectively refer to these rocks as Valdez/Orca metabasalts and later present geochemical analyses and modeling, suggesting a link with the Resurrection/Knight Island ophiolites and the time of ridge subduction.

4. GEOCHEMISTRY OF THE RESURRECTION PENINSULA AND KNIGHT ISLAND PILLOW BASALTS AND SHEETED DIKES 4.1. Effects of Greenschist-Facies Metamorphism Sample preparation for analysis of major, minor, and trace elements by inductively coupled plasma (ICP) spectrometry was as described by Lytwyn et al. (1997). Static hydrothermal alteration of the pillow lavas and sheeted dikes, as evidenced by zeolite to lower greenschist-facies metamorphism in both ophiolites, could result in mobility of certain elements such as Na, K, Ca, Sr, Ba, and Rb. Although groundmass alteration and devitrification is evident in some analyzed samples, phenocryst and microphenocryst phases are generally unaltered. The extent of clinopyroxene and plagioclase alteration in diabases exhibiting somewhat greater metamorphism is still less than 15 percent. Elements with low (less than three) ionic potentials (ratios of atomic number/ionic radius) such as Sr, K, Ba, and Rb are considered mobile in aqueous fluids relative to those having larger ionic potentials (greater than three) such as REE and high field strength elements (e.g., Pearce and Peate, 1995). Sr is typically enriched (displays positive anomalies on extended spiderdiagrams) in extensively altered metabasites (Pearce and Cann, 1973), whereas most of the ophiolitic samples, although enriched in Ba, exhibit negative Sr anomalies (Fig. 5) which is more consistent with plagioclase fractionation. Hydrothermal alteration (Fig. 5) causes Mg uptake in basalt (Mottl, 1983). Resurrection/Knight Island samples (e.g., K-13 and RP-390B) with relatively high MgO (8.25 and 9.02 wt%), however, also have the lowest abundances of incompatible trace elements (Fig. 5) which further supports the idea that their geochemistry is largely primary and little affected by hydrothermal processes. 4.2. Compositional Range and Tectonic Setting Pillow lavas and sheeted dikes from the Resurrection Peninsula and Knight Island ophiolites range from basalt to basaltic andesite (Table 1). Plots of major, minor, and trace

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Fig. 5. Spiderdiagrams of (A) Resurrection Peninsula, (B) Knight Island, and (C) Valdez/Orca samples from this study normalized to the chondritic values of Anders and Grevesse (1989). Fig. 2c also includes metabasalt sample APR164C from the Valdez Group (Lull and Plafker, 1990). Elements are arranged from left to right in order of increasing compatibility in solid mantle phases.

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Table 1. Inductively coupled plasma whole rock analysis of ophiolitic pillow lavas and sheeted dikes RP86B233B (Dike) 51.93 1.64 14.41 11.58 0.23 6.75 8.10 3.04 1.17 0.16 99.01 2.34 122 115 40 39 329 303 NA NA 7.0 18.5 13.6 4.04 1.33 5.4 6.4 4.2 3.82 0.60 10.41

RP86B246B (Pillow) 53.22 1.99 14.02 10.60 0.21 5.92 8.07 5.18 0.24 0.22 99.67 1.59 161 122 53 39 61 343 NA NA 8.3 22.2 17.2 5.15 1.72 6.9 8.2 5.4 4.97 0.77 9.53

RP86B322B (Pillow) 51.36 1.37 14.43 10.39 0.19 7.21 10.47 3.23 0.44 0.15 99.24 1.67 103 226 34 42 112 303 NA NA 5.1 13.7 10.7 3.21 1.22 4.5 5.4 3.6 3.24 0.50 9.34

RP86B361B (Pillow) 55.71 1.86 15.89 9.68 0.13 5.50 6.30 4.23 0.20 0.20 99.70 2.50 167 163 46 37 133 298 NA NA 9.8 24.4 16.8 4.92 1.48 6.3 7.4 4.8 4.40 0.69 8.70

RP86B365B (Pillow) 51.50 1.58 15.36 9.82 0.16 8.30 6.40 4.22 1.06 0.19 98.59 2.86 146 95 46 36 434 275 NA NA 9.4 23.1 16.2 4.58 1.46 5.9 6.9 4.5 4.17 0.63 8.83

RP86B375B (Pillow) 48.90 1.58 18.77 10.34 0.22 5.54 9.71 4.46 0.58 0.22 100.32 5.79 154 164 37 34 324 300 193 59 7.5 19.8 14.3 3.81 1.13 4.8 5.9 3.8 3.57 0.51 9.29

RP86B390B (Pillow) 50.51 1.31 16.13 9.51 0.15 8.25 11.64 2.00 0.10 0.15 99.75 2.09 103 142 30 35 53 251 NA NA 6.7 16.8 11.3 3.21 1.09 4.1 4.8 3.1 2.81 0.44 8.55

RP86B408B (Pillow) 54.17 1.23 15.30 9.63 0.15 8.24 8.96 2.38 0.33 0.20 100.59 2.88 111 177 30 33 156 236 NA NA 7.4 18.0 12.0 3.36 0.99 4.2 4.8 3.1 2.81 0.42 8.66

RP86B420B (Pillow) 51.30 1.32 16.22 9.60 0.16 7.16 8.45 3.68 1.35 0.16 99.40 2.59 124 181 34 37 521 258 NA NA 7.0 17.7 12.4 3.61 1.24 4.6 5.5 3.5 3.19 0.48 8.63

RP86B451B (Pillow) 53.11 1.48 15.74 10.33 0.17 6.09 5.61 5.36 1.00 0.18 99.07 1.94 146 119 42 33 270 288 NA NA 9.1 22.8 15.1 4.14 1.32 5.5 6.4 4.2 3.94 0.62 9.29

RP86BK-2 471B (Pillow) (Dike) 52.66 53.94 2.10 0.99 13.49 15.44 13.78 8.37 0.23 0.13 6.00 7.12 7.50 8.35 3.32 5.10 0.80 0.15 0.25 0.15 100.13 99.74 2.45 2.25 172 87 118 78 58 27 41 33 263 24 359 233 NA NA NA NA 8.2 4.1 23.2 11.2 18.8 7.9 5.58 2.51 1.88 0.83 7.5 3.5 9.0 4.2 5.9 2.8 5.40 2.60 0.83 0.42 12.39 7.52 (continued on next page)

Chapter 20: The Resurrection Peninsula Ophiolite

SiO2 TiO2 Al2 O3 Fe2 O3 * MnO MgO CaO Na2 O K2 O P2 O5 Sum LOI (%) Zr Sr Y Sc Ba V Cr Ni La Ce Nd Sm Eu Gd Dy Er Yb Lu FeO*

RP86B207B (Pillow) 50.23 1.44 16.69 8.55 0.17 7.55 9.85 2.99 1.61 0.14 99.22 2.91 104 123 34 46 237 264 NA NA 5.4 15.3 11.5 3.40 1.29 4.6 5.6 3.6 3.33 0.51 7.69

SiO2 TiO2 Al2 O3 Fe2 O3 * MnO MgO CaO Na2 O K2 O P2 O5 Sum LOI (%) Zr Sr Y Sc Ba V Cr Ni La Ce Nd Sm Eu Gd Dy Er Yb Lu FeO*

K-4 (Pillow) 51.17 1.21 15.47 9.85 0.18 7.99 10.52 3.60 0.17 0.13 100.29 1.74 100 175 31 39 81 285 NA NA 4.0 11.5 9.2 2.83 1.06 3.9 4.9 3.2 2.98 0.46 8.86

K-8 (Pillow) 50.40 1.19 16.18 10.38 0.16 8.42 10.53 3.44 0.14 0.09 100.93 2.22 72 139 30 43 75 262 NA NA 3.4 9.9 8.0 2.55 0.99 3.8 4.6 3.1 2.88 0.46 9.33

K-11 (Pillow) 50.15 1.20 16.14 9.87 0.15 7.87 12.06 2.15 0.05 0.12 99.76 1.65 85 118 30 38 24 244 NA NA 3.8 11.1 8.8 2.75 1.03 3.9 4.7 3.1 2.85 0.44 8.87

K-13 (Pillow) 52.02 0.79 15.88 7.18 0.14 9.02 11.78 2.33 0.06 0.09 99.29 2.44 71 152 20 28 50 159 NA NA 3.4 9.6 6.8 2.08 0.70 2.7 3.2 2.0 1.90 0.31 6.45

K-19 (Pillow) 53.84 1.13 15.62 9.37 0.13 7.92 7.25 5.05 0.06 0.11 100.48 2.73 88 95 29 32 36 245 264 195 4.0 12.3 9.5 2.61 1.04 3.7 4.5 3.0 2.84 0.41 8.42

K-21 (Pillow) 51.23 1.13 15.70 9.34 0.14 7.07 11.21 4.00 0.08 0.12 100.02 3.02 84 144 25 32 35 236 324 121 3.3 10.8 8.4 2.15 0.85 3.2 3.9 2.6 2.40 0.34 8.40

K-24 (Pillow) 48.25 1.29 17.06 9.59 0.17 8.47 11.70 2.55 0.07 0.13 99.28 2.46 91 153 29 35 46 259 421 164 2.7 10.3 9.5 2.65 0.96 3.8 4.7 3.0 2.84 0.40 8.62

K-26 (Pillow) 50.81 1.21 16.30 9.45 0.18 8.43 12.01 2.79 0.10 0.13 101.41 1.83 77 126 29 35 28 232 NA NA 3.8 10.9 8.7 2.69 0.97 3.8 4.6 3.1 2.74 0.43 8.86

K-35 (Pillow) 53.83 1.87 15.45 11.51 0.17 4.93 5.26 5.34 0.30 0.26 98.92 2.42 198 118 53 30 143 329 NA NA 7.6 22.7 17.8 5.20 1.37 6.9 8.2 5.5 5.12 0.80 10.35

K-37 (Pillow) 40.44 1.96 18.37 13.59 0.20 4.99 14.74 3.78 0.07 0.32 98.46 8.39 123 158 46 45 67 381 NA NA 7.3 17.3 13.8 4.17 1.31 5.8 7.0 4.7 4.19 0.65 12.22

K-42 (Pillow) 50.09 1.21 16.29 10.17 0.15 9.68 10.88 2.04 0.20 0.12 100.83 3.79 92 135 26 33 90 254 356 178 3.6 11.5 9.2 2.54 1.01 3.5 4.2 2.7 2.56 0.37 9.14

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Samples from Resurrection Peninsula indicated by RP and Knight Island by K. NA, not analyzed.

K-30 (Pillow) 50.23 1.75 15.69 11.31 0.15 6.31 10.79 3.35 0.05 0.19 99.82 3.17 133 115 43 37 12 301 NA NA 5.6 16.3 13.3 4.12 1.47 5.6 6.8 4.5 4.12 0.65 10.17

4. Geochemistry of the Resurrection Peninsula and Knight Island Pillow Basalts and Sheeted Dikes

Table 1. (Continued)

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Fig. 6. Samples plotted on binary diagrams of (A) FeO* /MgO versus SiO2 showing the calc-alkaline (CA) and tholeiitic (Th) fields of Miyashiro (1974) and (B) K2 O versus SiO2 showing the low-K (tholeiitic), medium-K (calc-alkaline), and high-K (calc-alkaline) subdivisions of Le Maitre et al. (1989) and Rickwood (1989). Symbols in these and subsequent diagrams are as follows: solid circles, Resurrection Peninsula pillow lavas and sheeted dikes from this study; open circles, Resurrection Peninsula pillow lavas and sheeted dikes from Crowe et al. (1992); solid triangles, Knight Island pillow lavas from this study; open triangles, Knight Island pillow lavas and sheeted dikes analyzed by Crowe et al. (1992) and Nelson and Nelson (1993); pluses, Valdez/Orca metabasalts from this study; crosses, Valdez metabasalts from Lull and Plafker (1990) and Crowe et al. (1992). Field of basalts and andesites from the Taitao Ophiolite, South America, from Kaeding et al. (1990).

4. Geochemistry of the Resurrection Peninsula and Knight Island Pillow Basalts and Sheeted Dikes

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Fig. 7. Samples plotted on various discriminant diagrams using same symbols as Fig. 5. Open squares represent greenstones (Hill, 1979) from the Ghost Rocks Formation. The abbreviations are as follows: OFB, ocean floor basalts; LKT, low-K tholeiites; CAB, calc-alkaline basalts; WPB, within-plate basalts; MORB, mid-ocean ridge basalts; IAT, island-arc tholeiites; OIT, ocean-island tholeiites; OIA, ocean-island alkalic basalts. Discriminant diagrams are from the following sources: Ti/Zr/Y and Ti versus Zr from Pearce and Cann (1973), TiO2 /MnO/P2 O5 from Mullen (1983), and Zr/Y versus Zr from Pearce and Norry (1979).

elements show only modest compositional differences between the two ophiolites (e.g., Figs. 5–9) with the Resurrection samples somewhat more enriched in alkalis and light rare earth elements (LREE—Fig. 10). Binary plots of major and minor elements and ratios reveal significant compositional scatter in the ophiolitic data when utilizing SiO2 (Fig. 6) and MgO (Fig. 8) as measures of differentiation (discussed below). Chondrite-normalized spiderdiagrams (Fig. 5) show that Resurrection and Knight Island samples generally have flat to slightly LREE-enriched patterns and overall REE abundances ranging between 10X and 40X chondritic. Samples also display overall negative Sr, Eu, and Ti anomalies (relative to adjacent REE), positive Zr anomalies, and negative

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Chapter 20: The Resurrection Peninsula Ophiolite

Fig. 8. Geochemical binary plots of Resurrection Peninsula and Knight Island pillow lavas and sheeted dikes on coordinates of (a) SiO2 , (b) TiO2 and (c) CaO/Al2 O3 versus MgO. Data symbols same as Fig. 5. The field of Valdez/Orca metabasalts (lighter shaded) defined from data by Lull and Plafker (1990) and Crowe et al. (1992), while the field of Ghost Rocks greenstones (darker shaded), Kodiak Island, is based on the data of Hill (1979). Liquid lines of descent for perfect fractional crystallization (PFX) of selected parental liquids RP-365B, RP-390B, and K-13 at pressures of 1 bar and 10 kbar were calculated using an algorithm after Weaver and Langmuir (1990). The compositional range of Resurrection and Knight Island samples cannot be related through fractionation of a common parental magma.

to positive Ba anomalies (Fig. 5). When normalized to normal mid-ocean ridge basalts (N-MORB) as in Fig. 13, Resurrection and Knight Island lavas and dikes show relative enrichments in highly incompatible elements (Th, K, Rb, and Ba) (Crowe et al., 1992; Nelson and Nelson, 1993) and display slight LREE enrichments similar to transitional (Ttype) mid-ocean ridge basalts (normalization values from Sun and McDonough, 1989). Major and minor element plots of FeO* /MgO and K2 O versus SiO2 indicate that Resurrection and Knight Island lavas and dikes include both tholeiitic and calc-alkaline varieties

5. Valdez/Orca Metabasalts

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Fig. 9. Comparisons of samples discussed in the text with calculated MORB parental magmas (after Niu and Batiza, 1991) on coordinates of (a) SiO2 /FeO* versus MgO and (b) TiO2 versus SiO2 /FeO* . Heavy lines represent the range of calculated parental magma compositions produced through polybaric melting of rising peridotitic material beginning at different pressures of melting (8–20 kbar) and continuing to around 4 kbar pressures. Lines that are near orthogonal to the melting trends connect equal values of F (degrees of melting) at intervals of 0.05. Liquid lines of descent (thin lines) for selected parental liquids were determined with an algorithm after Weaver and Langmuir (1990). The selected parental magmas represent F = 0.24 following initial melting at 18 kbar, F = 0.18 following initial melting at 14 kbar, and F = 0.20 following initial melting at 10 kbar. Data sources for the fields of Ghost Rocks greenstones and Valdez/Orca metabasalts same as Figs. 3 and 4.

(Fig. 5) as is also indicated on the AFM (A = Na2 O + K2 O, F = FeO + Fe2 O3 , M = MgO) diagram (Crowe et al., 1992; Nelson and Nelson, 1993). Some discriminant diagrams, utilizing trace elements (e.g., Ti, P, Zr, Y, REE, etc.) considered relatively immobile during low-temperature alteration and metamorphism up to greenschist facies (Pearce and Cann, 1973; Pearce and Norry, 1979; Pearce et al, 1981; Pearce, 1982; Mullen, 1983), define Resurrection and Knight volcanics and dikes as mid-ocean ridge basalts (MORB) and ocean floor tholeiites (Fig. 4). The Th-Hf-Ta discriminant diagrams, however, identify Resurrection and Knight Island lavas and dikes as calc-alkaline basalts and island-arc tholeiites (Crowe et al., 1992; Nelson and Nelson, 1993).

5. VALDEZ/ORCA METABASALTS Valdez/Orca metabasalts analyzed in this study (Table 2) and Lull and Plafker (1990) are generally more depleted in incompatible trace elements (REE, Ti, Zr, Y, etc.) than lavas and dikes from the Resurrection and Knight Island ophiolites (Fig. 5). Valdez/Orca basalts include both tholeiitic and calc-alkaline varieties (Figs. 6A, 6B, and 7).

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Fig. 10. N-MORB normalized spiderdiagrams comparing Knight Island lavas and dikes (shaded field) with a basaltic glass (D42-4) from the Chile Ridge (Klein and Karsten, 1995) and a calc-alkaline basalt from the southern volcanic zone (SVZ) of the Andes (Hickey et al., 1986). Knight Island data from Nelson and Nelson (1993) were used in order to compare Ta, Th, Nb, Cs, Rb, and U not analyzed in this study. All the spiderdiagrams patterns are broadly similar in terms of enrichments in highly incompatible elements (e.g., Cs and Ba) and negative anomalies in certain HFS elements (e.g., Th and Ta). Normalization values for N-MORB from Sun and McDonough (1989).

6. SIGNIFICANCE OF MAGMATIC ROCKS Geochemical analysis of the igneous rocks of the Resurrection and Knight Island ophiolites as well as related igneous rocks of the Valdez and Orca Groups show that they are chemically similar to mid-ocean ridges or primitive island arcs (Crowe et al., 1992;

6. Significance of Magmatic Rocks

647

Fig. 10. (Continued.)

Nelson and Nelson, 1993; Lytwyn et al., 1997). All the authors, however, have realized the arc-like chemistry of these rocks to be incompatible with the field evidence in the Gulf of Alaska, as no arc remnants are exposed outboard of the Gulf of Alaska. This observation is critical for the interpretation of ancient sequences, where the accretionary prism has experienced a collision and original relationships are not as clear. The arc-like chemistry of the most-outboard accreted ophiolites serves as a lesson to those who would use geochemistry alone to attempt to discriminate tectonic environments. Crowe et al. (1992) suggested sediment contamination to explain the variable volcanic rock compositions in an accretionary setting, having identified subduction of transforms and ridges as instances where MORB type magmas might be contaminated by sediments. Nelson and Nelson (1993) believing a contractional forearc setting is incompatible with a strong MORB-like signature and the presence of sheeted dikes, suggested that the migration of magma along a partially subducted ridge to an unsubducted part of the ridge can explain the mixing of sediments and MORB magma. Lytwyn et al. (1997) determined that mafic rocks of the Resurrection Peninsula and Knight Island have similar petrographic and geochemical compositions, which indicates a similar origin and tectonic history. The mafic rocks formed at a distance from the paleotrench have a strong MORB geochemical signature, whereas contemporaneous magmas derived from the same mantle melting column but intruded through the accretionary prism have more calc-alkaline signatures, reflecting near-fractional melting of parental magmas and assimilation of sedimentary material from the accretionary prism. These refractory arctype magmas came from shallow melts that are unable to mix with deeper less refractory magmas once ridge subduction occurred, suggesting ophiolite genesis at a spreading ridge near a subduction zone (Lytwyn et al., 1997, 2000; Kusky and Young, 1999; Bradley et al., 2003). With a greater number of plates and an even greater number of triple junctions in the Archean, calc-alkaline ophiolites generated at ridges near subduction zones were probably much more common in the earlier record.

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Table 2. Inductively coupled plasma whole rock analysis of Valdez and Orca metabasalts Latitude Longitude SiO2 TiO2 Al2 O3 Fe2 O3 * MnO MgO CaO Na2 O K2 O P 2 O5 Sum Zr Sr Y Sc Ba V Cr Ni La Ce Nd Sm Eu Gd Dy Er Yb Lu FeO*

82AMH-52A 60◦ 58 40 145◦ 53 37 50.20 0.47 15.08 9.33 0.22 10.98 12.00 1.54 0.12 0.05 99.99 29 51 15 42 23 248 594 179 0.68 2.2 1.8 0.62 0.26 1.5 2.2 1.7 1.64 0.26 8.40

82AMH-40C 60◦ 50 57 146◦ 34 32 51.20 0.79 15.52 9.14 0.14 9.20 12.61 2.02 0.11 0.08 100.81 49 86 24 43 26 292 307 99 2.10 5.9 4.7 1.48 0.57 2.6 3.4 2.5 2.50 0.41 8.22

82AMH-43A 60◦ 50 39 146◦ 31 50 50.88 0.88 15.03 10.82 0.17 8.58 11.82 1.66 0.10 0.08 100.02 52 75 25 44 27 289 283 96 1.6 5.6 5.1 1.70 0.65 2.9 3.8 2.7 2.67 0.43 9.73

7. SEDIMENTARY ROCKS OF THE RESURRECTION PENINSULA Sedimentary rocks associated with the Resurrection ophiolite include typical Valdeztype flysch rocks west of the Fox Island shear zone, and an unusual group of thinly bedded flysch (Humpy Cove Formation of the Orca Group, Kusky and Young, 1999) east of the Fox Island shear zone (Fig. 2).

7. Sedimentary Rocks of the Resurrection Peninsula

649

7.1. Valdez Group, West Shore of Resurrection Bay Generally north-striking beds of the Valdez Group flysch along the western shore of Resurrection Bay comprise massive coarse-grained sandstone, containing carbonate nodules, with interbedded 1–3 cm thick siltstone and shale couplets. In the central parts of the bay several massive units of black mud-matrix debris flows or olistostromes with blocks of graywacke up to 0.5 m in length out crop intermittently. Deformation within the thick mudstone units decreases to the north where numerous quartz veins dominate the outcrop. Folded quartz-veins and boudins show west-over-east asymmetry. 7.2. Humpy Cove Formation of the Orca Group Sedimentary rocks of the Resurrection Peninsula deposited above the ophiolite include variably metamorphosed graywacke, siltstone, and shale. Because these rocks rest conformably over a 57 Ma ophiolite, and are cut by 53.4 Ma plutonic rocks, they are temporally correlated with the Orca Group. However, because of their unusual lithologic nature and unique tectonic setting the name “Humpy Cove Formation” of the Orca Group was proposed for these rocks (Kusky and Young, 1999), including description of a formal type section extending along the north shore of Humpy Cove and extending through the north shore of Fox Island (Fig. 2). Detailed descriptions and chemical analyses of these rocks are provided by Kusky and Young (1999). The turbidites resting conformably over the ophiolite typically comprise monotonously bedded 1–5 cm thick graded argillite/shale couplets with 10–20 cm thick beds of siltstone and thicker (0.2–1 m) sandstone/graywacke beds. Most of the shales are metamorphosed slates up to greenschist facies, but for the purposes of the sedimentological description the term “shale” is retained. Beds strike to the north-northeast and dip west. Small-scale sedimentary structures such as flute marks, cross-laminations, ripples, and parallel-laminations are only rarely present, precluding designation of Bouma facies. Diagenetic carbonate concretions, approximately 5–20 cm in diameter are abundant in the graywacke of Thumb Cove, forming bedding-parallel horizons within the units (Young, 1997). These nodules form elongate, darkly weathering pits, and weather to a chalky brown color and typically appear slightly flattened where the exposure is three-dimensional. Local depositional sequences show variations in the proportion of coarse- to finegrained horizons. These variations include an increasing thickness of sandstone beds to 10 m or more, without a corresponding increase in finer-grained layers, and rhythmically interbedded, 1–5 cm thick siltstone and shale couplets in sequences ranging to approximately 10 m thick. A slatey cleavage is typically developed in the fine-grained rocks. Where black shales are present, the layers tend to be thin interbedded shale and siltstone couplets with only minor beds of sandstone, if present at all. The black shales typically have a fissile fabric. Volcanic tuffs interbedded with turbidites form volumetrically minor lenses that pinch and swell, making them difficult to use as marker beds across widely separated outcrops. The proportion of volcanic beds increases up-section where they form a notable portion of the rocks of Fox Island (Fig. 2).

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Except in the Fox Island shear zone, bedding is fairly coherent, with localized deformation of bedding typically expressed as open to isoclinal folding, slumping confined to individual strata, and minor shortening or extension of more competent beds. Pervasive quartz veins filling bedding-parallel cleavage planes also show isoclinal folding in places. 7.3. Age and Duration of Sedimentation of the Humpy Cove Formation The 57 ± 1 Ma age of the Resurrection Peninsula ophiolite represents a maximum age for turbidites deposited conformably over the ophiolite. A 40 Ar/39Ar cooling age on biotite of 53.4±0.9 Ma for the Hive Island stock (Fig. 2) that intrudes along a continuation of the Fox Island shear zone south of Fox Island (Bradley et al., 1999b) provides a constraint on the minimum age for sediments structurally below the shear zone. This age is a biotite cooling age for the pluton; consequently, the pluton may have been emplaced well before 53.4 ± 0.9 Ma. Thus we estimate that the sediments above the ophiolite, up to the Fox Island shear zone, took approximately 3.6 Ma to accumulate, from 57 to 53.4 Ma (PaleoceneEocene). This estimate for the time duration of sedimentation involves some additional assumptions. First, the 57 ± 1 Ma age for the ophiolite is from a trondhjemite that may have intruded late in the seafloor spreading history of the ophiolite, and could conceivably even be coeval with some of the tuffaceous rocks in the Humpy Cove Formation. However, the trondhjemite intrudes the sheeted dike complex and is also cut by dikes of the sheeted complex, so it presumably represents an accurate age for the ophiolite. There also may be an unknown time delay between the age of the ophiolite and initiation of sedimentation of the Humpy Cove Formation, although this too is unlikely since flysch of the lowermost Humpy Cove Formation is interbedded with pillow lavas of the ophiolite. The upper age bracket on the age of the Humpy Cove Formation is provided by the 53.4 ± 0.4 Ma 40 Ar/39 Ar age of an intruding pluton; since this is a cooling age, the age of intrusion is older. In the area, there is typically about a 2 Ma time difference between U/Pb and 40 Ar/39 Ar ages for plutonic rocks of this age (Bradley et al., 1999b). Taking these assumptions and possible sources of error into account, the interval bracketed for deposition of the Humpy Cove Formation could be as long as 58 to 53 Ma (5 m.y.), or as short as 56 to 53.8 Ma (2.2 m.y.). Accordingly, we adopt 3.6 ± 1.4 Ma as the best estimate for the duration of sedimentation of the Humpy Cove Formation.

8. FOX ISLAND SHEAR ZONE The Fox Island Shear Zone is a greater-than 1-km wide north-striking ductile shear zone that juxtaposes deformed and metamorphosed sedimentary and volcanic rocks of the Valdez Group to the west over lower-grade metasedimentary Orca Group rocks lying above the ophiolite to the east (Fig. 2). The eastern unit on Fox Island comprises interbedded sedimentary and volcanic rocks of similar metamorphic grade to rocks of the Resurrection

8. Fox Island Shear Zone

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Peninsula (Fig. 2). Sedimentary rocks are primarily blue-gray to black shales interbedded with thin siltstone and silty sandstone layers with locally massive sandstone units of 1–2 m thickness. Tuffs form individual units ranging from approximately 0.5 to several meters thick. The tuffs are lenticular with finely laminated light green to light brown patches. The fine-grained groundmass contains plagioclase phenocrysts, quartz augen, and elongate lenses a few centimeters long that appear to be flattened volcanic breccia or pumice. Where these lenses are present, they are either light green or brown and the groundmass is brown or light green. Some of the quartz augen (less than one centimeter) have rotated asymmetric tails indicating a west over east sense of shear. Very fine-grained black shales dominate outcrops on the northeastern side of the island. These rocks are interbedded with thin beds of sandy siltstone and typically have a fissile fabric that shows flow along disrupted and sheared bedding planes. Thin zones of folding and boudinage occur throughout the section. Stratigraphically above the black shales, volcanic tuffs increase in both the number of beds and in thickness of individual layers. A few volcanic flows or sills are present near the southern end of the section. In cliffside exposures, alternating tuff, mudstone, and graywacke beds are coherent and planar. Shear zones that are typically several tens of centimeter wide cut the northeast side of the island. Siltstone beds are pulled apart and isoclinally folded. Bedding on either side of these zones is relatively undeformed and planar. Isoclinally folded layers include doubly plunging folds indicating rotation of fold axes into the extensional direction. Motion along these shear zones is west over east with rotation of the fold axial surface down to the northwest. Similar orientations of beds and foliations on Fox Island as well as folds for all of Resurrection Bay indicate are compatible with west over east thrusting (e.g., Kusky et al., 1997b; Kusky and Young, 1999). A different style of deformation is evident approximately 500 m south, near the sand spit that juts out from the eastern side of Fox Island (Fig. 2). Alternating 1-cm thick siltstone and 2-cm thick mudstone beds are chaotically folded but do not exhibit the fissile fabric or flow along bedding surfaces that are observed in the zone of folding to the north. Quartz veins filling cleavage planes are axial planar to small folded quartz veins that cut the outcrop. Those perpendicular to cleavage are flattened, whereas those oriented along cleavage planes are extended, indicating strong flattening. Minor sandy beds are boudinaged, whereas the tops and bottoms of folded horizons appear welded or annealed, suggesting that soft-sediment deformation may have preceded tectonic flattening in this zone. Central and Western Fox Island consists predominantly of metavolcanic and metasedimentary rocks within the Fox Island shear zone (Fig. 11) that records ductile, west-overeast thrusting. Rocks of the central and western units are more deformed and have a higher metamorphic grade than rocks to the east. Mafic and pelitic rocks exhibit a schistosity close to the Fox Island shear zone (Fig. 11). Secondary minerals in sedimentary rocks include biotite (giving a purple hue to the siltstones and mudstones), muscovite, and lesser amounts of chlorite. However, due to the wide range of temperatures and pressures over which these minerals are stable, we cannot specify a unique temperature and pressure under which these

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Chapter 20: The Resurrection Peninsula Ophiolite

Fig. 11. Photographs of high strain fabrics and porphyroclasts indicating west-over-east thrusting in the Fox Island shear zone. (a) Sigma-shaped asymmetric porphyroclasts of quartz indicating a west-over-east sense of shear (top of the photo moved to the upper left). (b) Doubly-plunging fold in graywacke/argillite in a shear zone on the eastern side of Fox Island. (c) Staircase trajectory in sigmoidal-shaped quartz layers in high-grade psammite of the western belt in the Fox Island shear zone, also indicating west-over-east shearing (top of the photo moved to the left).

8. Fox Island Shear Zone

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rocks formed from these data. In comparison, eastern unit sedimentary rocks contain minor biotite, muscovite, and calcite. Locally abundant chlorite in the volcanic rocks provides a greenish tint and greasy feel to the rocks. One-to-three centimeter chlorite and amphibole porphyroclasts locally produce a schistose fabric. Recrystallised quartz-augen are present in all rock types. The central and western units are distinguished from one another on the basis of lithology and degree of deformation. The central unit is characterized by thick sandstone sequences, whereas the western unit is dominated by volcanic tuffs and sills. Deformation increases across the central and western units to the west. The units are separated by a mylonitic shear zone that roughly follows the changing lithology from the sandstone-rich to the volcanic-rich unit and is mapped as the thrust surface in figures showing the shear zone. Deformation within the central unit is characterized by abundant quartz veins filling bedding-parallel cleavage planes. Secondary foliations typically are isoclinally folded, whereas porphyroclasts, such as sigmoidal augen, provide excellent kinematic indicators (Fig. 11). Disruption of bedding occurs as boudinage of sandy and silty layers with finer grained rocks exhibiting flow fabrics around the boudins. Minor shortening of beds is evident as small-scale ramps duplexing competent layers. All fabrics exhibiting ductile flow features indicate west-over-east kinematics (Fig. 11). As is observed on Resurrection Peninsula, minor but ubiquitous brittle late down-to-thewest normal fabrics are locally superimposed on rocks of the central unit. This deformation is manifested as quartz-filled tension gashes in siltstone and sandstone, and as kink banding in finer grained rocks. Metamorphic rocks of the western unit were highly deformed along the ductile thrust fault that separates the central from the western portions of the island. Disrupted bedding is nearly ubiquitous as multiple duplexed packages are ramped one upon the other making identification of individual beds nearly impossible. Thin quartz veins filling planar foliations are typically isoclinally folded whereas thicker quartz veins are elongated in sigmoidal augen with strongly elongated shear-sense indicators (Fig. 11). Mylonitic C-S fabrics also indicate west-over-east shearing. Thrusting produced alternating layers of volcanic and pelitic rocks at all scales. Rusty weathering, broken beds of graywacke form lozenges in thrust-imbricated packages 2– 3 m in length and 0.5–1 m thick. These duplexed thrust packages are separated by 2 to 3-cm thick shear zones expressed by finer grain size and a closely spaced foliation. Similar deformation is observed in duplexed thrust packages at the scale of tens of meters made up of many smaller ramped packages. Motion along the Fox Island shear zone produced some large-scale structures. In Sunny Cove the Fox Island shear zone is associated with upright, steeply west-plunging isoclinal folds. Very large (∼ 10 m) laterally extensive boudins are visible in the cliff faces along the south side of the island. Some of these boudins can be traced laterally through a portion of the island and are observed where the beds reappear on the opposite side.

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9. INTERPRETATION OF SEDIMENTARY FACIES OF THE HUMPY COVE FORMATION Flysch of the Resurrection Peninsula may be expected to show features within the graded beds that illustrate the type of flow regime present when they were deposited. However, primary depositional features such as cross-bedding, parallel-laminations, ripple marks, flute casts, and load casts are only rarely observed, precluding assignment of typical Bouma Ta-Te classification schemes, or submarine fan facies models (Mutti and Ricci Lucchi, 1978; Walker, 1979, 1992; Howell and Normark, 1982; Miall, 1984). Examination of units where deformation appears relatively minor indicates that these features remain absent. Depositional features that may be expected even in a low flow regime apparently were not preserved during initial deposition. Sandstone units in Thumb Cove increase in thickness above the contact with the ophiolite, forming a coarsening-upward sequence. This is expected in settings such as that interpreted here for the Humpy Cove Formation, in which the depositional site is conveyed from a ridge to an abyssal or outer trench slope environment, into a zone of active terrigeneous sedimentation in a trench. Channel features are present only at the top of the Humpy Cove Formation, with erosional features at the base of the sandstone bodies absent, indicating that the lateral extent of the sandstones is not controlled by channels, but by a general waning of the transport process (e.g., Mutti and Ricci Lucchi, 1978; Howell and Normark, 1982; Shanmugam and Moiola, 1985). Thin-bedded sandy siltstones and shales interbedded with pillow basalts of the ophiolite suggest deposition away from the main sink of coarse-grained terrigeneous sediments. These sandy siltstones were deposited during active magmatism at the ridge when the ophiolite was at a distance from the trench, and therefore represent a distal depositional environment. Deposition along the active ridge axis suggests that the turbidity current must have greatly reduced its velocity during its encounter with the topographically high ridge, permitting only the finer-grained suspended particles to travel up the ridge crest. Therefore sedimentation is expected to occur as fine-grained, thinly-bedded turbidites.

10. SEDIMENTATION AT A TRIPLE JUNCTION DURING A RIDGE-TRENCH INTERACTION Subduction of the Kula-Farallon ridge beneath the North American continental margin is well documented (Marshak and Karig, 1977; Sisson and Hollister, 1988; Lonsdale, 1988; Bol et al., 1992; Bradley et al., 1993, 2003; Sisson et al., 1994; Lytwyn et al., 1997; Kusky et al., 1997a, 1997b; Sisson et al., 2003). Subduction of the Kula-Farallon ridge accompanied by eastward migration of the trench-ridge trench triple junction along the trench axis (Bradley et al., 1993, 1999b) formed a topographic high that moved east along the trench axis and which may have acted as a barrier that inhibited sediment transport (Fig. 12). As sediment transport was dominantly (though not universally) westward along

10. Sedimentation at a Triple Junction during a Ridge-Trench Interaction 655

Fig. 12. Schematic diagram showing sedimentation patterns around postulated tectonic setting for the formation and history of the Resurrection Peninsula ophiolite. The ophiolite formed along the Kula-Farallon spreading center at 57 ± 1 Ma and was transported toward and emplaced into the Chugach accretionary prism. Continentally-derived flysch filtered along the ridge axis became interbedded with pillow lavas, and some may have escaped out along transform offsets to add a terrigeneous component to those sediments deposited on top of the ophiolite as it was transported toward North America (Humpy Cove Formation). As the ophiolite entered the trench, turbidite sedimentation became dominated by thicker, more massive turbidites with a stronger biogenic component, suggesting higher marine productivity (upwelling or shallow water?) in the source area. CCD is carbonate compensation depth.

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the trench axis (Nilsen and Zuffa, 1982; Plafker et al., 1994; Kusky et al., 1997b), several things may have occurred during the eastward migration of the ridge (Fig. 12). Uplift and erosion of the accretionary prism over the migrating ridge would redeposit previously deposited sediments on either side of the migrating ridge (e.g., DeLong and Fox, 1977; Forsythe and Nelson, 1985; Nelson et al., 1993). Sediment being transported westward along the trench axis would encounter the ridge axis and begin to slow, potentially depositing coarser material than would otherwise be deposited if there were no barrier to sediment transport. Where sediments may have jumped the ridge, the effect may be analogous to channel overbank deposits allowing only the finer grained sediments to spill over (Fig. 12). Sediment sources perpendicular to the trench axis would supply additional input to the main sedimentary feeder channels (Nilsen and Zuffa, 1982), although sediment input on the western side of the ridge axis may still be expected to travel west along the trench to be deposited away from the ridge. A transition from black shales and siltstones to turbidites with fewer black shales and then massive sandstones is recorded in the stratigraphic profiles of the Humpy Cove Formation. This change in lithology occurs by approximately 1400 m above basalt in all three stratigraphic sequences. Most chemical profiles also exhibit large concentration changes at 1400 m with decreases in terrigeneous elements and increases in biogenic elements that is best illustrated in the Sc and Sr chemical profiles (Kusky and Young, 1999). The close association between variations in chemical stratigraphy and changing lithologies between the three sequences at 1400 m suggests that the same event is recorded in each of the sequences. The transition from a regime dominated by thin-bedded turbidites to one dominated by thicker sandstone beds is here interpreted as a transition from distal to more proximal turbidite facies associated with transport of the ophiolitic basement to the Humpy Cove Formation toward the continent (Fig. 12). Turbidites interbedded with pillow lavas, and those just above the pillow basalts in Humpy Cove, are all thinly bedded, suggesting that they are the result of low flow velocity or distal turbidity currents. The transition to thicker-bedded sandstone turbidites stratigraphically up section is interpreted as increasing proximity to the continental margin where an increased volume of sediment was deposited. Early in the history of the ophiolite, distal turbidity currents produced a few pulses of thin-bedded siltstones, each approximately 10 m thick along the still active ridge. Since these also show a strong terrigeneous component, it is likely that they represent distal turbidites transported along the axial rift valley of the Kula-Farallon spreading center (Fig. 12). Sources could include both sediments transported down submarine canyons cutting through the accretionary prism, and turbidites that initially flowed along the trench, with sufficient velocity to jump and flow along the ridge (Fig. 12). After the ophiolite was conveyed away from the ridge, a 1000-m thick sequence of thin-bedded turbidites containing black shales, siltstones, and thin graywacke beds accumulated above the pillow basalts in Humpy Cove. Continued transport, both away from the ridge and toward the continent, gradually resulted in a depositional regime of increasingly proximal turbidite sedimentation characterized by lighter-colored shale beds with a greater proportion of sandstones

11. A Neogene Analog: Chile Rise and Taito Ophiolite

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making up individual turbidites. This period of sedimentation may mark the transition of the location of the ophiolite from the outer trench slope to the floor of the trench where axially transported turbidites rapidly accumulated (Fig. 12). 11. A NEOGENE ANALOG: CHILE RISE AND TAITO OPHIOLITE The emplacement of the Taito ophiolite during subduction of the Chile Rise beneath the South American margin represents a Neogene example of ophiolite emplacement during ridge subduction (Herron et al., 1981; Schweller et al., 1981; Forsythe and Nelson, 1985; Cande and Leslie, 1986; Forsythe et al., 1986; Forsythe and Prior, 1992; Nelson and Forsythe, 1989; Bangs et al., 1992; Nelson et al., 1993; Lagabrielle et al., 1994; Sisson et al., 1994). Comparison with the Resurrection Peninsula ophiolite illustrates differences in sedimentation resulting from emplacement under different conditions: the Chile Rise strikes nearly parallel to the trench, whereas the Kula-Farallon ridge had a more oblique strike to the trench. The Taito ophiolite was emplaced at 3 Ma during passage of the Chile Rise along the south American margin near the Gulfo de Penas basin (Nelson et al., 1993). The Taito sedimentary cover comprises thin marine sedimentary rocks overlying conglomeratecontaining clasts of the pre-Jurassic metamorphic basement structurally beneath the ophiolite. The absence of a thick sedimentary cover sequence is interpreted to be the result of ophiolite emplacement at the location where the Chile Rise came into contact with the continent. Passage of the ridge beneath the accretionary prism produced near trench plutonic rocks and localized uplift that exposed the prism above sea level shedding sediments to either side of the trench prior to obduction. The buoyant, topographically high ridge, maintained a very shallow-dipping angle of subduction. The close proximity of the Taito ophiolite to the point of contact between the Chile Rise and the continent suggests that the ophiolite was emplaced at very shallow levels in the accretionary prism. In contrast, the thick sequence of sedimentary rocks over the Resurrection Peninsula ophiolite and the ductile shearing along the Fox Island shear zone suggests that this ophiolite was emplaced into the accretionary prism at depth and exposed at the surface only after uplift and erosion. This agrees with the idea that the ophiolite was transported a maximum of approximately 90 km from the ridge crest during its 3.6 ± 1.4 Ma transport history.

12. IMPLICATIONS FOR UNDERSTANDING RIDGE-TRENCH ENCOUNTERS IN THE PRECAMBRIAN The southern Alaska convergent margin consists of belts of accreted PaleozoicCenozoic flysch, mélange, and ophiolites. Paleogene near-trench plutons intruded the margin diachronously from 61 Ma in the west to 50 Ma in the east during migration of a trench-ridge-trench triple junction. Recognizing these plutons as products of ridge subduction has several implications for forearc evolution and interpretation of linear belts of

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plutons in ancient orogens. Forearcs are not necessarily places characterized exclusively by high-P, low-T metamorphic series and a lack of plutonism, but may contain high-T, low-P metamorphism in association with belts of plutonic rocks if the forearc was affected by ridge subduction (Fig. 13). Similarly, belts of magmatic rocks in ancient orogens may not necessarily represent individual arc terranes, but could be a paired arc/forearc system that experienced ridge subduction (Fig. 13) (Kusky et al., 2003). The record of ridge subduction events varies considerably depending on plate geometry and rates of triple junction and slab window migration (e.g., Sisson et al., 2003). However, some of the hallmark signatures of ridge subduction in forearcs include along strike diachronous intrusion of TTG (tonalitic, trondhjemitic, granodiorite), to granitic plutons, high-T metamorphism, diachronous gold mineralization, and belts of anomalous complex faulting (Fig. 13). Structural, thermal and magmatic aspects of the Chugach terrane are similar to the geology of Archean granite-greenstone terranes. In both, deformation is locally melt-dominated, and plutons follow a low-K series from diorite to trondhjemite. Metamorphism is of a high-T, low-P series. Most Archean granite-greenstone terranes acquired their first-order structural and metamorphic characteristics at convergent plate margins, where large accretionary wedges similar in aspect to the Chugach, Makran, and Altaids grew through offscraping and accretion of oceanic plateaux, oceanic crustal fragments, juvenile island arcs, rifted continental margins, and pelagic and terrigeneous sediments. Some suites of TTG in these terrains appear to have been generated during ridge subduction events, suggesting that ridge subduction is an important process in continental growth. Ridge subduction was likely an important process in the Archean, when the total number of plates was higher, and the number of ridge-trench encounters was greater. The southern Alaska margin serves as a relatively modern example of processes important in Archean forearc evolution and continental growth.

13. FLYSCH AND MÉLANGE DOMINATED ACCRETIONARY PRISMS: PHANEROZOIC AND PRECAMBRIAN EXAMPLES Many contemporary and ancient accretionary prisms are dominated by two fundamentally different types of structural belts-mélange and relatively coherent flysch terrains (e.g., Fig. 1b). However, mélange terranes are very rare in Archean convergent margin sequences, and only slightly more common in Early Proterozoic orogens. Controls on the accretion of flysch and mélange terranes at convergent margins are poorly understood. The Chugach terrane of southern Alaska represents the Mesozoic outboard accretionary margin of the Wrangellian composite terrane. It consists of two major lithotectonic units, including Triassic-Cretaceous mélange of the McHugh Complex, and Late Cretaceous flysch of the Valdez Group. Mapping on the Kenai Peninsula, and in the eastern Chugach Mountains has clarified the nature of the boundary between the mélange and flysch terrains. In most places, the contact is a gradational or interleaved boundary between chaotically deformed mélange of argillite, chert, greenstone, and graywacke of the McHugh Complex, and a less chaotically deformed mélange of argillite and graywacke of the Valdez Group. The latter

13. Flysch and Mélange Dominated Accretionary Prisms: Phanerozoic and Precambrian Examples

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Fig. 13. Three-dimensional sketch indicating the relative position and timing of ridge subduction, formation of the slab window, generation of the Nuka-Aialik, and Harris Bay plutons and their intrusion into the Resurrection ophiolite, and activation of the late out of sequence thrust faults including the Fox Island shear zone.

is known as the Iceworm mélange (Kusky et al., 1997a, 1997b), and is interpreted it as a contractional fault zone (Chugach Bay thrust) along which the Valdez Group was thrust beneath the McHugh Complex. The McHugh Complex had already been deformed and metamorphosed to prehnite-pumpellyite facies prior to formation of the Iceworm mélange.

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(a)

(b) Fig. 14. Schematic diagram showing possible different accretion styles for (a) a thinly and (b) a thickly sedimented oceanic plate. These differences may account for the paucity of mélanges in Archean terranes.

We interpret accretion of the mélanges of the McHugh Complex as a normal response to subduction of oceanic crust bearing a thin pelagic sedimentary veneer (Fig. 14), whereas accretion of the coherent thrust packages of the Valdez Group is interpreted as a normal wedge response to subduction of oceanic crust bearing a thick pile of loosely consolidated submarine fan and trench sediments (Fig. 14). Thrusting of the McHugh Complex over the Valdez Group and the formation of the Chugach Bay thrust and Iceworm mélange may represent a dramatic critical taper adjustment to these changing subduction regimes.

13. Flysch and Mélange Dominated Accretionary Prisms: Phanerozoic and Precambrian Examples

661

Davis et al. (1983) defined a popular Coloumb wedge model for the mechanics of accretionary prisms in which the stress is everywhere as large as possible, with the wedge thickening toward the rear because of the increased strength of rock in this direction. For a wedge of given strength (determined by rock rheology, basal resistance, pore fluid pressure, etc.), the forces that resist motion of the wedge are balanced by the forces that control the shape of the wedge (e.g., the thickness of the wedge determines the cross-sectional area across which the principal horizontal stress can act). In this way, the combination of surface slope plus basal decollement dip (critical taper) is maintained by deformation of the wedge, as material is added to the toe or base of the wedge. Platt (1986) analyzed the dynamics of accretionary wedges with negligible yield strength and demonstrated that underplating can oversteepen the surface slope, causing extension in the upper part of the rear of the wedge. This regains the equilibrium critical taper and causes uplift of metamorphic rocks in the rear of the accretionary wedge. These models are directly applicable to the Chugach accretionary wedge in that the McHugh complex is interpreted as material off-scraped intermittently during Triassic-Early Cretaceous subduction, with the establishment of a critically tapered wedge (Fig. 14). Attempted subduction of the thick sedimentary veneer of the Valdez Group during Maastrichtian-Paleocene time dramatically upset the critical taper by both rapid frontal accretion and increased underplating (Fig. 14b). The wedge’s response to this episode of rapid accretion was to deform internally to regain the critical taper. Rapid frontal accretion would cause the wedge to thicken to maintain the critical taper, and in the case of the Chugach terrane much of this thickening was accommodated on the Chugach Bay thrust, emplacing the older part of the wedge (McHugh Complex) over the younger part (Valdez Group). Formation of the Iceworm mélange may have been aided by fluids released from dewatering of the recently accreted Valdez Group escaping upward along the lower boundary of the relatively impermeable McHugh Complex. The Chugach Bay thrust may also have formed where it did because the protoliths of the Iceworm mélange represent a muddy facies of the Valdez Group with considerably more pore fluids that in other parts of the Valdez Group. Elevated pore fluid pressure within part of the accretionary wedge decreases the effective strength of the rock, providing a horizon for enhanced localized deformation. Increased underplating of the wedge during accretion of the Valdez Group caused enhanced uplift of the rear of the accretionary wedge, probably with associated extensional faulting, (e.g., Platt, 1986), providing a mechanism for the uplift and exposure of the metamorphosed McHugh Complex, and its emplacement over the Valdez Group on the Chugach Bay thrust (Fig. 14). Differences in accretion style analogous to the McHugh/Valdez dichotomy are suggested to be characteristic of subduction of thinly versus thickly sedimented oceanic plates, and we cite here a few additional examples to emphasize the generality of this model. Late Cretaceous uplift and erosion of the Mongolia-Okhotsk intracontinental collision zone in Asia shed a large amount of sediments to the south and east (Klimetz, 1983; Nanayama, 1992), which triggered rapid frontal accretion and growth of accretionary complexes along the northwest Pacific rim (Kimura, 1994). Kimura (1994) also documented that rapid underplating by attempted subduction of thick sedimentary caps on the downgoing plates caused uplift of metamorphic rocks of the Kamuikotan complex of Hokkaido

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and the Susunai complex of Sakhalin. Taira et al. (1988) have documented rapid growth of the Nankai accretionary wedge in response to the Izu collision. Tertiary uplift and erosion of the Himalayan mountain range from the India-Asia collision caused rapid growth of the Makran and Sunda accretionary prisms from sediments supplied rapidly to the trench by the Indus and Ganges Rivers (Graham et al., 1975; Burbank et al., 1996; Sengör and Natal’in, 1996). Similarities are also seen between the model presented here and the ongoing collision of the Yakutat terrane with the southern Alaska margin. This collision has caused uplift and erosion of the Chugach and St. Elias Mountains, deposition of a thick sequence of Miocene and younger clastic sediments in the Aleutian trench, and rapid accretion at the toe of the Aleutian accretionary prism (Plafker et al., 1994). Climate, as well as collisions, may play a significant role in episodic growth and structural style of accretionary prisms (e.g., Hay, 1996; von Huene and Scholl, 1991). Bangs and Cande (1997) have shown how a thickly-sedimented segment of the Chile trench south of the Chile triple junction in the area strongly influenced by glacial climate corresponds to rapid growth and accretion in the accretionary prism. The subducting Chile ridge acts as a topographic barrier to sediments shed into the trench by glacial erosion in the south, and growth of the prism north of the triple junction is less rapid than that to the south of the triple junction. Periods of rapid accretion were also noted during intervals when glaciation extended farther north and erosion accelerated trench sedimentation north of the triple junction. Subduction of the Chile ridge causes significant erosion of the accretionary complex, making the pre-ridge subduction record south of the northward migrating triple junction difficult to interpret. We suggest that, in general, subduction of slabs with a thin veneer of sedimentary cover may lead to subduction erosion (e.g., von Huene and Scholl, 1991) or to the formation of mélanges in accretionary wedges, whereas subduction of thickly-sedimented slabs may generally result in large-scale accretion of relatively coherent packages separated by zones of shearing and type I mélange formation. Using a constant slip-rate model, subduction of an oceanic slab with a thin sedimentary veneer will result in significantly higher shear strains in the boundary layer between the upper and lower plates than would subduction of an oceanic slab with a thick sedimentary cover. For example, if we assume one kilometer of thrusting per million years and we accommodate this strain in a 100-m thick sedimentary section, then shear strains of ten will be typical for the sedimentary cover and will likely involve significant mixing between layers. If the subducting oceanic plate has a 1-km thick sedimentary cover, then shear strains will be about one, and we may expect strains that are lower than in the mélange and perhaps more typical of a foreland fold-thrust belt. This model may also explain a fundamental temporal change in the style of subduction zone tectonism. Many workers have noted the near absence of subduction mélanges from Archean terranes, even those interpreted as ancient accretionary wedges (e.g., Kusky, 1989, 1990, 1991). Models for Archean plate tectonics predict that the planet may have had a greater plate boundary length, with more, smaller, faster moving plates than at present (e.g., Pollack, 1997). One consequence of such a model is that subducting plates would have tended to be thickly sedimented, with their proximity to active collisions along numerous plate boundaries. Subduction of these thickly sedimented plates would tend to

14. Ridge Subduction Model for Evolution Archean Greenstone Belts, TTG Magmas, and VMS Deposits

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produce Archean accretionary wedges dominated by relatively coherent terranes and few mélanges, consistent with the geological record preserved in Archean granite greenstone terranes (e.g., de Wit and Ashwal, 1997).

14. RIDGE SUBDUCTION MODEL FOR EVOLUTION ARCHEAN GREENSTONE BELTS, TTG MAGMAS, AND VMS DEPOSITS Archean greenstone belts are 1- to 100-km scale terranes of linear to curvilinear volcanic and sedimentary supracrustal sequences mostly with tectonic boundaries. These greenstone belts typically record poly-phase deformation, greenschist to amphibolite facies metamorphism, multiple phases of granitoid intrusion, deposition of giant volcanogenic massive sulphide (VMS), banded iron formation, and gold deposits, consistent with complex geodynamic interactions between various geological processes operation at Archean subduction zones (Kusky and Polat, 1999). The formerly contiguous late Archean Abitibi-Wawa subprovince, Superior Province, is the largest Archean greenstone-granitoid terrane globally, exposed over a strike length of greater than 1000 km, and was built by the amalgamation of volcanic and sedimentary rocks and intrusion of granitoids with diverse geochemical compositions (Stott, 1997; Polat and Kerrich, 2001). The formation of the Wawa-Abitibi greenstone-granitoid terrane occurred over a relatively short period, from 2.75 to 2.65 Ga, diachronously from north to south by lateral spread of subduction-accretion complexes (Percival et al., 1994; Calvert and Ludden, 1999; Polat and Kerrich, 2001). Many geological features of the Wawa-Abitibi greenstone-granitoid terrane can be better explain by complex interactions between intra-oceanic arcs, and mantle plumes and subducted ridges. For example, the subduction of ridge systems followed by opening asthenospheric windows in the late Archean Wawa-Abitibi subduction zone may have provided optimized thermal conditions for the formation of temporally and spatially associated ultramafic to felsic volcanic rocks (e.g., boninites, picrites, adakites, magnesian andesites, rhyolites, etc.), intrusion of TTG plutons, amphibolite metamorphism, and the generation of VMS, BIF and lode gold deposits. 14.1. Ridge Subduction, T TG Petrogenesis and Continental Growth It is generally accepted that Archean continental crust grew by accretionary and magmatic processes taking place at convergent plate boundaries (Taylor and McLennan, 1995; Kusky and Polat, 1999; Polat and Kerrich, 2001). Many continental growth models suggest that 60–70 percent of the present-day continental crust was formed by the end of the Archan (see Taylor and McLennan, 1995; Rudnick, 1995). Polat and Kerrich (2000) showed from field relationships and geochemical considerations that late Archean continental crust in the southern Superior Province grew by mixing of oceanic plateau and subduction derived components (granitoids, tholeiitic to calc-alkaline bimodal volcanic rocks, trench turbidites) at convergent plate boundaries. Simple mixing calcula-

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tions suggest that 6–12 percent oceanic plateau, mixing with a 88–94 percent arc magmas, is required to produce late Archean continental crust in the southern Superior Province. Several geochemical studies have suggested that the growth of the Archean continental crust took place in convergent margin settings mainly by intrusion of TTG, which constitute the larger portions of the Archean continental crust (greater than 70 percent). These TTG magmas are thought to result from reworking of basaltic crust which was either subducted or accreted (Moorbath, 1977; Taylor and McLennan, 1985, 1988; Tarney and Jones, 1994; Polat and Kerrich, 2000); TTG are genetically linked to eclogites through reworking of basaltic crust (Rollinson, 1997; Rapp et al., 2003). However, the reworking mechanism(s) by which basaltic crust is transferred to TTG and an eclogitic residue in Archean convergent margins has not been fully understood. The origin of Archean TTG is mainly attributed to melting of subducted oceanic crust (Martin, 1999; Drummond et al., 1996; Rapp et al., 1999). Melting of subducted slabs alone, however, cannot explain the presence of the eclogitic residue in the continental lithosphere because the denser refractory eclogitic slabs would have been recycled into the deep mantle (MacDonough, 1991). To explain the presence of eclogites and TTG suits in the continental lithosphere, here we propose that partial melting of accreted and/or underplated oceanic plateaus and normal oceanic crust in the accretionary wedges and/or on the base forearc by upwelling of an asthenospheric window, following a ridge subduction episode, played a major role in the generation and preservation of the Archean continental crust. The partial melting of accreted basaltic crust, which was metasomatized by previous slab-derived melts and/or fluids, occurred under amphibolite to eclogite metamorphic conditions. Diaprically upwelling tonalite, trondhjemite, granodiorite, diorite, mozonite, and sanukitoid melts derived from partial melting of basaltic crust on the base of thickened accretionary complexes formed the continental crust. The complementary eclogitic residues contributed to the sub-continental lithospheric mantle and further growth of the lithospheric mantle resulting from underplating mantle plumes (see Wyman et al., 2002). TTG suites that intruded Archean oceanic island arc accretionary complexes formed the nuclei of intra-oceanic island arcs, subsequent accretion oceanic island arcs and oceanic plateaus, and oceanward migration of the arc-trench system resulted in the lateral growth of Archean continental crust in subduction zones. 14.2. Ridge Subduction and VMS Deposits There are several present-day examples of VMS deposits resulting from ridge subduction. The Bransfield Strait of the Antarctic Peninsula, for example, is currently undergoing extension due to subduction of the Antarctic-Phoenix spreading centre (Lawver et al., 1995). There is an extensive hydrothermal activity in Bransfield Strait, resulting in deposition of VMS (Lawver et al., 1995). Development of a mantle window following the ridge subduction may have triggered the extension of the forearc oceanic lithosphere by thermal erosion. An upwelling of mantle window beneath the extending forearc oceanic crust could have triggered intensive hydrothermal activity and deposition of the VMS in

15. Conclusions

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Phanerozoic orogenic belts. Archean VMS deposits likely to have been formed by similar geodynamic processes. On the basis of striking similarities between the geological characteristics of Archean greenstone belt and those of the Cenozoic Alaskan subductionaccretion complex (see Crowe et al., 1992; Nelson and Nelson, 1993; Kusky et al., 2003; Sisson et al., 2003), we propose that many Archean VMS deposits formed during ridgesubduction events.

15. CONCLUSIONS Recognizing the Sanak-Baranof belt as a product of ridge subduction has several implications for forearc evolution and the interpretation of linear belts of plutons in ancient mountain belts. Forearcs are not necessarily exclusively places characterized exclusively by high-P, low-T metamorphic series and a lack of plutonism, but may contain high-T, lowP metamorphism in association with belts of plutonic rocks if the forearc was affected by ridge subduction. Similarly, belts of magmatic rocks in ancient mountain belts may not necessarily represent individual accreted island arc terranes, but could be a paired arc/forearc system that experienced a ridge subduction event. The record of ridge subduction events may vary considerably in individual examples depending on the plate geometry and rates of triple junction and slab window migration. However, some of the hallmark signatures of ridge subduction in forearcs include ophiolite emplacement, time-transgressive intrusion of tonalitic, trondhjemitic, granodiorite, to granitic plutons, high temperature metamorphism, and diachronous gold mineralization, and belts of anomalous complex faulting. Faults can control pluton emplacement, both as inactive zones of structural weakness, and as active dilational bends (Kusky et al., 2003). Structural, thermal and magmatic aspects of the Chugach terrane are similar to the geology of Archean greenstone-granodiorite terranes (Pavlis et al., 1988; Kusky, 1989; Barker et al., 1992; Kusky and Polat, 1999). In both, deformation is locally meltdominated, and plutons follow a low-K series from diorite to trondhjemite (Pavlis et al., 1988). Metamorphism is of a high-temperature low-pressure series. Ridge subduction was most likely an important process in the Archean, when the total number of plates was higher, and the number of ridge-trench encounters was greater. The southern Alaska margin then may serve as a relatively “modern” example of processes that were likely to have been important in Archean forearc evolution and continental growth. The Resurrection Peninsula and Knight Island ophiolitic lavas/dikes exhibit both calcalkaline and tholeiitic geochemistry possibly reflecting fractionation from multiple, compositionally diverse parental magmas coupled with variable assimilation of flysch metasedimentary rock. Parental magmas may be modeled as near-complete mixtures of geochemically diverse, near-instantaneous melts derived from a polybaric mantle column through melting of variably depleted mantle source (Lytwyn et al., 1997). Compositional variation among the parental magmas may primarily reflect different degrees of overall mantle melting. Slight alkali and LREE enrichments in Resurrection Peninsula and Knight Island lavas/dikes may indicate assimilation of flysch by invading N-MORB, suggesting

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formation along an oceanic (possibly Kula-Farallon) ridge system in close proximity to a trench. In contrast, refractory metabasalts from the Valdez/Orca accretionary complex possibly formed through inefficient mixing of near-instantaneous melts extracted mostly from shallower, more depleted regions of the mantle melting column (Lytwyn et al., 1997). Basalts from the Resurrection and Knight Island ophiolites, Ghost Rocks Formation, southern Kenai Peninsula and possibly Valdez/Orca accretionary complex may be related to high-angle subduction of the Kula-Farallon ridge system. Mantle convection beneath the mid-ocean ridge portion permits near-instantaneous melts from shallower, more refractory mantle sources to efficiently mix with melts from deeper, more fertile sources to produce MORB-like parental magmas. Ridge subduction, however, limits the top of the melting column to greater depths beneath the overiding plate and leads to inefficient pooling (mixing) of near-instantaneous melts as reflected in refractory basalts along the forearc. Basalts generated along a subducting spreading center, and erupted within or seaward of the forearc accretionary prism, may acquire a calc-alkaline overprint through assimilation of sediments and crustal rocks by upward-migrating melts along the asthenospheric window/accretionary prism interface. Resurrection Peninsula sediments from Thumb Cove, Humpy Cove, and Fox Island, located directly above the Resurrection Peninsula ophiolite, show a progression from thin turbidite deposition characterized by black shale, siltstone, and predominantly thin (10– 20 cm) sandstone beds, to thick massive turbidites characterized by shale and siltstone couplets with predominantly thick (greater than 50 cm) sandstone beds. This transition is interpreted as a shift from distal, deeper water turbidites deposited nearest the ophiolite as it migrated away from the ridge, to more proximal turbidite deposition as the ophiolite was transported into the trench and was emplaced in the accretionary prism forming along the North American continental margin. The Fox Island shear zone is a west-over-east ductile thrust that places older turbidites of the Valdez Group above younger turbidites deposited on the Resurrection Peninsula ophiolite. The shear zone is the upper bounding fault along which the ophiolite was emplaced, cutting sediments deposited above the ophiolite. It developed after passage of the Kula-Farallon ridge when MORB and sediments mixed at the base of the prism and created the granitic melts which intrude the shear zone a few kilometers south of Fox Island. Metamorphism and ductile flow fabrics are interpreted as the result of shearing at depths of less than 20 km (Kusky et al., 2003) aided by the heat of nearby plutons.

ACKNOWLEDGEMENTS This work was supported by NSF Grants EAR-9304647, and EAR- 9706699 awarded to T. Kusky, and by the U.S. Geological Survey. We thank D. Bradley, S. Nelson, S. Bloomer, S. Dolan, and K. Dietrich for stimulating discussions and assistance in the field.

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Plate: Crustal Formation and Andean Convergence. Memoirs of the Geological Society of America 154, 323–350. Sengör, A.M.C., Natal’in, B.A., 1996. Paleotectonics of Asia: fragments of a synthesis. In: Yin, A., Harrison, T.M. (Eds.), The Tectonic Evolution of Asia. Cambridge Univ. Press, pp. 486–640. Shanmugam, G., Moiola, R.J., 1985. Submarine fans and related turbidite sequences. In: Bouma, A.H. (Ed.), Frontiers in Sedimentary Geology. Springer-Verlag, New York, pp. 27–35. Sisson, V.B., Hollister, L.S., 1988. Low-pressure facies metamorphism in an accretionary sedimentary prism, southern Alaska. Geology 16, 358–361. Sisson, V.B., Pavlis, T.L., 1993. Geologic consequences of plate reorganization: An example from the Eocene southern Alaska fore arc. Geology 21, 913–916. Sisson, V.B., Hollister, L.S., Onstott, T.C., 1989. Petrologic and age constraints on the origin of a low-pressure/high temperature metamorphic complex, southern Alaska. Journal of Geophysical Research 94, 4392–4410. Sisson, V.B., Pavlis, T.L., Prior, D.J., 1994. Penrose Conference Report: Effects of triple junction interactions at convergent margins. GSA Today 4 (10), 248–249. Sisson, V.B., Pavlis, T.L., Roeske, S.M., Thorkelson, D.J., 2003. Introduction: An overview of ridgetrench interactions in modern and ancient settings. In: Sisson, V.B., Roeske, S.M., Pavlis, T.L. (Eds.), Geology of a Transpressional Orogen Development during Ridge-Trench Interaction along the North Pacific Margin. GSA Special Paper 371, 1–18. Stock, J., Molnar, P., 1988. Uncertainties and implications of the late Cretaceous and Tertiary positions of North America relative to the Farallon, Kula, and Pacific plates. Tectonics 7, 1339–1384. Stone, D.B., Panuska, B., Packer, D.R., 1982. Paleolatitude versus time for southern Alaska. Journal of Geophysical Research 87, 3697–3707. Stott, G.M., 1997. The Superior Province, Canada. In: de Wit, M., Ashwal, L.D. (Eds.), Greenstone Belts. In: Oxford Monographs on Geology and Geophysics, vol. 35, pp. 480–507. Sun, S.-s., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes. In: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the Ocean Basins. Geol. Soc. Spec. Publ. 42, 313–345. Taira, A., Katto, J., Tashiro, M., Okamura, M., Kodama, K., 1988. Cretaceous to Miocene accretionary prism. Mod. Geol. 12, 5–46. Tarney, J., Jones, C.E., 1994. Trace element geochemistry of orogenic igneous rocks and crustal growth models. Journal of the Geological Society of London 151, 855–868. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, Cambridge, p. 312. Taylor, S.R., McLennan, S.M., 1988. The significance of the rare earths in geochemistry and cosmochemistry. In: Gschneidner Jr., K.A., Eyring, L. (Eds.), Handbook on the Physics and Chemistry of Rare Earths, vol. 2. Elsevier, New York, pp. 485–578. Taylor, S.R., McLennan, S.M., 1995. The geochemical evolution of the continental crust. Review of Geophysics 33, 241–265. Tysdal, R.G., Case, J.E., 1979. Geologic map of the Seward and Blying Sound Quadrangles, Alaska. United Stated Geological Survey Misc. Inv. Ser. Map I-1150, scale 1:250 000. Tysdal, R.G., Case, J.E., Winkler, G.R., Clark, S.H.B., 1977. Sheeted dikes, gabbro, and pillow basalt in flysch of coastal southern Alaska. Geology 5, 377–383. von Huene, R., Scholl, D.W., 1991. Observations at convergent margins concerning sediment subduction, subduction erosion, and the growth of continental crust. Reviews of Geophysics 29, 279–316.

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Walker, R.G., 1979. Turbidites and associated coarse clastic deposits. In: Walker, R.G. (Ed.), Facies Models. Ainsworth, Mitchner, ON, pp. 91–104. Walker, R.G., 1992. Turbidites and submarine fans. In: Walker, R.G., James, N.P. (Eds.), Facies Models: Response to Sea Level Change. Geol. Assoc. Canada, Ontario, pp. 239–263. Weaver, J.S., Langmuir, C.H., 1990. Calculation of phase equilibrium in mineral-melt systems. Comput. Geosci. 16, 1–19. Wiltse, M.A., 1973. Fe-Cu-Zn massive sulfide deposits in an ancient outer arc ridge-trench slope environment. Geol. Soc. Am. Abstr. Programs 5, 122–123. Winkler, G.R., 1976. Deep Sea fan deposition of the Lower Tertiary Orca Group, Eastern Prince William Sound, Alaska. In: Miller, T.P. (Ed.), Recent and Ancient Sedimentary Environments in Alaska, Proceedings of the Alaska Geological Society, Anchorage, 1975. Alaska Geological Society, pp. 189–204. Winkler, G.R., 1992. Geologic map and summary geochronology of the Anchorage 1◦ × 3◦ Quadrangle, southern Alaska. U.S. Geol. Surv., Misc. Investigation Series, Map I-2283, scale 1:250 000. Winkler, G.R., Miller, M.L., Hoekzema, R.B., Dumoulin, J.A., 1984. Guide to the Bedrock Geology of a Traverse of the Chugach Mountains from Anchorage to Cape Resurrection. Alaska Geological Society, p. 40. Wyman, D., Kerrich, R., Polat, A., 2002. Assembly of Archean cratonic mantle Lithosphere and crust: Plume-arc interaction in the Abitibi-Wawa subduction-accretion complex. Precambrian Research 115, 37–62. Young, C.P., 1997. The Resurrection Peninsula ophiolite: Emplacement in the southern Alaska forearc and controls on sedimentation during ophiolite transport. M.A. thesis. Boston University, Boston, MA.

Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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PHANEROZOIC ANALOGUES OF ARCHAEAN OCEANIC BASEMENT FRAGMENTS: ALTAID OPHIOLITES AND OPHIRAGS 1 AND B.A. NATAL’IN A.M.C. SENGÖR ¸

˙ Maden Fakültesi, Jeoloji Bölümü, Ayazaˇga 34469, Istanbul, ˙ ITÜ Turkey

To the memory of Academician Alexandr Valdemarovich Peive for his advocacy of ophiolites as records of past oceans and his recognition that the Altaids was a Phanerozoic factory for the continental crust

Ophiolites have long been regarded as rock associations typical of orogenic belts. In the framework of the theory of plate tectonics, they have been interpreted as remnants of the igneous basement of former oceans, the closure of which generated the orogenic belts. Such remnants occur in orogens either in the form of a complete suite displaying a sequence from pillow lavas and pillow lavas, sills and dykes, through a sheeted dyke complex to gabbros, peridotites and dunites to tectonised harzburgites, or they occur as dismembered pieces of such a suite. Following the 1972 Penrose definition, we restrict the term ophiolite to the more or less complete suite and call its various fragments formed by processes incorporating them into the continental crust ophirags. In the minds of most geologists, orogenic belts are linear/arcuate, long and narrow zones of intense deformation. That is why, irregularly shaped areas of widespread ‘orogenic deformation’ interspersed with abundant fragments of the members of the ophiolite association in the Precambrian, but especially in the Archaean, have been thought of as products of processes no longer operative. However, the geology of the Altaid orogenic system in Asia greatly resembles in its overall map aspects, lithological content, structural characteristics, and in the distribution and types of fragments of floors of former oceans to the Archaean granite-greenstone terrains. In the Altaids, ophiolites are now encountered in three main settings: (1) Ophiolites that occur as basement of ensimatic arcs, (2) Ophiolites and ophirags that occur in former forearcs now entrapped within transform sutures: (a) Ophiolites as backstop to accretionary wedges, (b) Ophirags within accretionary wedges, and (3) Ophiolites and ophirags in collisional suture zones that have usually evolved from members of the second category. Ophiolites and ophirags have a widespread distribution within the orogenic edifice. This distribution was brought about by processes that shaped the Altaid edifice, namely, generation of supra1 Also at ˙ITÜ Avrasya Yerbilimleri Enstitüsü, Ayazaˇga 34469, ˙Istanbul, Turkey.

DOI: 10.1016/S0166-2635(04)13021-1

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subduction zone forearc basements created by pre-arc spreading, back-arc basin opening, subduction-accretion, trench-linked strike-slip faulting including arc slicing and arc shaving faults and associated ocean floor spreading processes, and collision of buoyant pieces along suture zones. Without appreciating the nature and sequence of these processes and their superimposition, it is impossible to understand the rules that govern the distribution of oceanic basement fragments in the Altaids. These processes have led to a tremendous degree of structural shuffling of previously distant environments and a large degree of dismembering of formerly more complete ocean floor fragments. The preservation is highly selective and favours upstanding and buoyant segments of ocean floors. Such pieces are embedded most commonly in metapelitic/metapsammitic or, more rarely, in serpentinitic matrices in mélange/wildflysch complexes. We contend that the same rules apply to the greenstone belts of the Precambrian and greatly hinder their deciphering in the absence of biostratigraphic control.

1. INTRODUCTION The purpose of this paper is to present an overview of the distribution, current tectonic position and mode of origin of the remnants of oceanic basement rocks in the Altaids as likely analogues of similar rocks in many Precambrian, but mainly Archaean, orogenic belts. We contend, following Burke et al. (1976), Sengör ¸ et al. (1993), Sengör ¸ and Natal’in (1996a, 1996b), and Kusky and Polat (1999), that the Archaean greenstone belts represent mostly subduction-accretion complexes that became arc basements by migration of magmatic arc fronts as a consequence of subduction back-stepping. The processes that created the Archaean greenstone belts were no different, therefore, from those now creating magmatic arcs and associated subduction-accretion complexes, with the exception of those processes related to the higher heat output of the planet in the Archaean. Before we describe the geology of the Altaids and of the oceanic basement remnants in them, we review, from a historical perspective, the reasons why the recognition of the ArchaeanAltaid similarity has long been hindered. These reasons are in part associated with the role ophiolites were believed to play in orogenic processes.

2. OPHIOLITES AND OROGENY: A HISTORICAL REVIEW AND CRITIQUE 2.1. Ophiolites and Plate Tectonics The concept of ophiolite gained great popularity after the advent of plate tectonics, because ophiolites were thought to represent pieces of crust and upper mantle of the floors of now vanished oceans (see, for example, Coleman, 1977, parts I and II; also see the contributions in Dilek and Newcomb, 2003). Ophiolites were hailed as guides to the solution of two main problems: (1) identification of places where former oceans had vanished (i.e., suture zones)

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and (2) understanding of the nature of the oceanic crust and upper mantle and all the ensimatic constructs such as island arcs, guyots, oceanic plateaux and fracture zone crust and mantle that may get preserved as pieces of ‘oceanic crust and mantle’. An associated problem concerned the mechanism by which ophiolites were brought into juxtaposition with continental crust. All three of these problems were initially approached within the context of the large-scale geometric properties of plate tectonics and its immediate predecessor, the two-dimensional spreading-subduction hypothesis of Hess (1962) and Dietz (1961). 2.2. Ophiolites as Ocean Floor Remnants It quickly became accepted that the basic ophiolitic pseudostratigraphy correlated well with the geophysical layering of the oceanic crust. Initially, such correlations (e.g., Dietz, 1963) were based on a crude (and largely incorrect) oceanic model of mainly serpentinitic layer 3 underlying a basaltic + sedimentary layer 2, which then underlay a largely unconsolidated sediment cover forming the layer 1 (Hill, 1957; Hess, 1962). However, Bishopp (1952) already had suggested that the Troodos Massif in Cyprus, displaying a clear transition downwards from pillow lavas and pillow lavas and dykes, through a sheeted dyke complex to gabbros, peridotites and dunites to tectonised harzburgites, was most likely a piece of oceanic crust. Accumulating observations on other ophiolites in the world revealed a similar sequence (e.g., Davies, 1968, in Papua New Guinea; Reinhardt, 1969, in Oman; Bailey et al., 1970, in California; for a general assessment, see Peive, 1969), which rapidly led to a corrected and more detailed correlation (see the summaries in Coleman, 1971, Fig. 1, and Dewey and Bird, 1971, Figs. 1 and 2, and references therein). In the seventies a number of models were proposed for the detailed geometry and kinematics of spreading centres on the basis of ophiolite data (e.g., Greenbaum, 1972, 1977; Gass and Smewing, 1973; Sleep, 1975; Coleman, 1977; Dewey and Kidd, 1977). Dewey (1974, 1976) suggested, on the basis of the short time interval between generation and obduction, and of the reconstructed tectonic setting, of the Caledonian/Appalachian ophiolites, that most major ophiolite nappes may have originated in back-arc basins rather than in major oceans. While much ink was being spilled in correlating ophiolites and the oceanic crust (in major oceans and in marginal basins), Miyashiro (1973) made the seminal suggestion, on the basis of geochemistry, that the Troodos ophiolite may have been born in an island arc setting. It was later pointed out that the Troodos could not have been a ‘normal’ island arc, because the sheeted dykes clearly indicate an environment of very considerable extension, similar to mid-oceanic spreading centres, not habitually encountered in active island arcs (e.g., Saunders et al., 1980). This dilemma was resolved, when Taylor and Karner (1983) showed that ‘back-arc basin spreading’ had occurred in the Marianas and the South Scotia arcs before an arc was established. Pearce et al. (1984) called this process ‘pre-arc spreading’ and showed that most Neo-Tethyan ophiolite nappes south and east of the Carpathians had indeed formed in such environments of ‘pre-arc spreading’. Sengör ¸ (1990) pointed out that most large ophiolite nappes in the world probably formed in such a setting and that this may be a unique mechanism of forming giant ophiolite nappes.

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As giant ophiolite nappes were further autopsied, however, it turned out that not only supra-subduction zone pre-arc environments, but also fracture zone environments were represented in them. Saleeby (1977) first drew attention to this fact by interpreting the Kings-Kaweah ophiolitic mélange belt in California as a fossil fracture zone (also see Saleeby et al., 1978; Saleeby, 1981). Shortly afterwards, Karson and Dewey (1978), Simonian and Gass (1978), and Lemoine (1980) documented segments of fracture zones preserved in ophiolite nappes. This strengthened a speculation that Dewey had first emphasised in 1975, namely that the vicissitudes of finite plate evolution inevitably lead to former transform faults turning into subduction zones. This had been implicit in the classical Dewey and Bird (1971) paper (see their Fig. 4), but the realisation that ophiolites preserve sections of fracture zones greatly encouraged the interpretation that such subduction zones as the Marianas and the Aleutians, which are nearly orthogonal to former magnetic sea-floor spreading anomalies in the upper plate, most probably nucleated on former fracture zones (Casey and Dewey, 1984). One such case may now be just in the process of origination in the Gorringe Bank on the Azores Fracture Zone (Auzende et al., 1984). Thus obduction began to be thought also to be related to fracture zone locations in the oceans. It was also thought clear that along the trend of mountain ranges long and narrow belts of ophiolitic material with dominantly steep fabric, popularised by Gansser’s (1964, 1966) ‘Indus suture’, marked where two continents had been apposed after the intervening oceans had entirely disappeared by subduction. Both in the Himalaya and the Middle East, large, coherent ophiolite nappes were shown to root into such suture zones (Gansser, 1964; Ricou, 1971). Hamilton (1969) (also Hamilton, 1970) and later Ernst (1970) and Hsü (1971) followed an earlier speculation by Dietz (1963) by arguing that ophiolitic mélange was a chaotic mixture of ophiolite pieces, sedimentary rocks of diverse origin and metamorphic rocks of HP/LT type, mixed and further churned along a subduction zone (see Sengör, ¸ 2003, for the evolution of the mélange concept). They argued that present-day subduction-accretion complexes at the snout of subduction zones were most likely sites of mélange origin and accumulation. Although Dietz (1963) had approached the problem of the mixing of oceanic basement rocks with what he called ‘eugeosynclinal prisms’ from the viewpoint of a world-wide review, both Hsü and Hamilton considered the mélange problem from a Pacific Ocean perspective, where both active and fossil subduction-accretion complexes still largely face the ocean. However, earlier, Gansser (1955) had shown that ophiolitic mélange (called ‘coloured mélange’ by him: A. Gansser, personal communication, 7th May 2003) also occurred along narrow belts of intense deformation in the Alpide belts of the Middle East. When plate tectonics was applied to the interpretation of these orogenic belts, it was thought that these narrow mélange belts were simply sutures (Dewey and Bird, 1970; Dewey et al., 1973). This line of thought was hardly unexpected in view of a widespread conviction that ‘Alpine-type serpentinites occur in linear swarms of subparallel or en échelon masses along the axes of old and new belts of folded mountains, . . .’ (Dietz, 1963, p. 947). In fact, the main Zagros crush belt was shown to have along it both mélange stringers and coherent ophiolite nappes (Stöcklin, 1968). Mélange belts also turned up un-

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der many major ophiolite nappes (e.g., Gansser, 1964; Reinhardt, 1969) corroborating the view that mélanges marked sites of sutures. 2.3. Ophiolites as Suture Markers: A Historical Perspective It is important to underline the fact that all authors, including Dewey (2003), have emphasised the role of the ophiolites in the classical Alpine- and Himalayan-type (sensu Sengör, ¸ 1990, 1992) of long and narrow orogenic belts that result from continental collision. Even those authors writing from a non-collisional, Pacific, viewpoint have depicted the ophiolitic mélange belts as long and narrow units attached to the prows of arcs (for a recent example, see Moores, 1998). The consideration of ophiolites along narrow, well-defined zones is hardly surprising in a retrospective interpretation of the history of the ophiolite concept in the early 20th century. In 1904, Eduard Suess published a two-and-a-half page note, with no figures, in the proceedings of the French Academy of Sciences (Suess, 1904). In it, he noted the discovery of overthrusts in the Alps by French and Swiss geologists. He then compared the frontal regions of overthrusts with the internal ogives at one of the termini of the ice cap in Greenland between Kangersuk and the Nasasuak nunatak and suggested that such ogives were homologous to the arcs of overthrusts (arc de charriage). Suess emphasised that deep moraines are brought to the surface along such listric thrust faults. Suess argued that to understand the nature of this ‘grand phenomenon’ of overthrusting one must look at the ‘roots’ of the nappes. He noted that the roots of the grand Alpine nappes exhibit much mafic and ultramafic rock and he (incorrectly) correlated the ultramafics of the Ivrea Zone with some of the Penninic ophiolites (because he had remained ignorant of the associated submarine lavas in the ophiolites as opposed to the entirely plutonic/metamorphic subcontinental Ivrea body: see Bailey’s point in Bailey (1944, p. 755)). Suess’ next point concerned the exotic blocks of the Himalaya and the associated mafic and ultramafic rocks. He supported the view that they belong to an allochthonous unit expelled from the north and northeast. The implication that a phenomenon similar to the one seen in the Alps is clear. In a short paragraph, he compared the island festoons of the western Pacific with thrust fronts. At the end of his note, Suess mentioned Auguste Daubrée’s studies (Daubrée, 1879, 1888; see Sengör, ¸ 2003) showing that ultramafic rocks represent the deeper parts of the earth and emphasised their frequent association with fronts of nappes, and the observation that they never appear in forelands of mountain belts. Suess left it there, but his implication is clear: orogeny somehow brings rocks that usually form deep in the earth to the surface by means of very large overthrusts. He noted the association of jumbled rock masses, which we now call mélange, with this phenomenon (Sengör, ¸ 2003). This is the point that Gustav Steinmann followed up and showed that in the extremely deformed and jumbled rock associations of the so-called Aufbruchszone (= zone of piercing) in Graubünden in eastern Switzerland, the bottom of the immense Austroalpine nappe system was exposed and exhibited a thin zone of extremely deformed and jumbled rock

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association (composite thickness about 700 m, but in many places reduced to mere slivers: Trümpy, 1975). Before the advent of the nappe theory, the Aufbruchszone was assumed to represent a submarine ridge, separating the Western Alps from the Eastern Alps and characterised by the ‘eruption’ of gabbros and peridotites. Steinmann had selected this area for a detailed study by himself and his students. Schardt’s classic 1893 publication reinterpreting the entire Préalpes as a nappe prompted him to test the theory of nappes on Bündner examples. He found that although Schardt’s and later Lugeon’s parallelisation of the nappe structures in western Switzerland with those of eastern Switzerland was incorrect in detail, the broad outlines of the nappe theory not only stood the tests, but made many of Steinmann’s own observations intelligible. Among Steinmann’s observations was the intimate association of radiolarites, deep-sea muds and radiolarite-bearing pelagic limestones with gabbros and peridotites. Steinmann noted that this association, previously vaguely ascribed to eruptive processes along the Aufbruchszone, in reality represented an extremely deformed association that stemmed from the deepest part of the Alpine ‘geosyncline’ (he estimated the maximum water depth of the ‘geosyncline’, on the basis of the present-day analogues of the radiolarites and the associated deep-sea sedimentary rocks, to be more than 4000 m: Steinmann (1905, p. 55); see Trümpy (2003) for an excellent historical assessment). At this depth, gabbros, peridotites and the deep sea sediments somehow had come together and were then uplifted by the medial Cretaceous by some 5 km! Steinmann ascribed this uplift to an early phase of Alpine orogeny that had folded up the middle part of the Alpine geosyncline (he regarded the Helvetic and the Austroalpine shelf areas as the northern and the southern parts, respectively, of the same geosyncline). Continued shortening finally generated the two largest nappes of the Alps from the deepest parts of the geosyncline: the ophiolite (by which Steinmann meant only the ultramafic rocks and the gabbros in 1905) and deep-sea sediment-filled northern part formed the Rhaetic Nappe, whereas the relatively shallower southern part formed the Austroalpine Nappe. The Austroalpine Nappe overrode the Rhaetic Nappe and literally squashed it (‘ausgequetscht’, Steinmann, 1905, p. 58). For that reason, Termier (1903, p. 762) had called the Austroalpine Nappe the traîneau écraseur (crushing sledge). This crushing sledge had such an effect on its underlying nappe, according to Steinmann, that ‘the structure of the middle or the “piercing zone” [‘Aufbruchszone’] became dominated by numerous smaller overthrusts and squashing, so much so, that a complete agreement with the confusing structure of the northern Swiss klippes originated and that, even while mapping at a scale of 1:25,000, only “squashing zones” [Quetschzonen] can be discerned, which almost earn the designation of a friction breccia [Reibungsbrekzie] at large’ (Steinmann, 1905, p. 10). This confusing, mega-breccia structure (which the Swiss geologist Joos Cadisch called the Aroser Schuppenzone [= the Zone of Imbrication of Arosa] in Cadisch et al. (1919, p. 362)) and the stratigraphic similarities led Steinmann to correlate the Rhaetian Nappe with the northern Swiss klippes and the Préalpes. As yet, the father of the Steinmann trinity was not aware of the importance of the third member of the trinity, namely the pillow lavas. In his classic 1927 paper he completed the trinity (but without naming it as such and without admitting the extrusive nature of the ‘diabases’! see Bailey (1936, 1944) and Trümpy (2003)) and pointed out the settings in which it was

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encountered (see Bailey, 1936, pp. 1718–1722). They were what we today recognise as the suture zones of the Mediterranean and the mélange belts associated with them. Thus, Suess’ prophetic point that ultramafic rocks seemed to be associated with major thrusts culminated, in the light of the continuing Alpine discoveries, in Steinmann’s point that the trinity he noticed pointed to the presence of oceanic depths, now represented by the extremely deformed zones of confused structure along thrust belts. Already in his 1905 paper Steinmann pointed out that the ophiolites (then thought of as only the ultramafics and the gabbros) and the associated deep-sea deposits (i.e., the ophiolitic mélange, a term Steinmann did not use) not always formed ‘middle nappes’ as in the Alps. In some mountain belts (as in the Himalaya) they constituted the highest nappes. Steinmann (1905, p. 64) pointed out that ophiolites issued most commonly from the ‘internal’ parts of geosynclines. Steinmann’s point that ophiolites were associated with deep-sea sedimentary rocks and thus must have issued from the deepest part of the Alpine geosyncline found a sympathetic reader in Suess, although Suess was no believer in geosynclines (Suess, 1888, pp. 263–264, 1909, p. 722 and note 52; cf. Sengör, ¸ 1982a, 1982b, 1998). He had satisfied himself that the Alps formed from the destruction of the Tethys, which he had regarded as an ocean, no different in structure from the present-day oceans. But he agreed with Steinmann that ophiolites, representing former deep-sea environments, were confined to a narrow zone in a mountain belt and repeated that they never occurred on the foreland. However, Suess also made the important observation, which was largely ignored after him, that ophiolites never occupy an axial position in mountain belts (Suess, 1909, p. 644). He also cautioned against an overenthusiastic acceptance of the idea that all ophiolites had issued from the deep sea by pointing out that the ophiolites in the western Pyrenees were not associated with deep sea sedimentary rocks, although, there also, significant dislocation accompanied them (Suess, 1909, p. 646). Suess regarded ophiolitic igneous rocks as sills, largely confined to surfaces of tectonic movement, but also underlined that in many cases the suspicion that they had been transported in a solid state seemed unavoidable: ‘This suspicion disappears [i.e. turns into a certainty] concerning the large surface of movement on which the Tibetan slabs were transported onto parts of the Himalaya’ (Suess, 1909, p. 646). In 1911, in a short note, Émile Argand demonstrated that in the Pennine zone of the Swiss and Italian Alps, the higher one looked in the nappe edifice, the more frequently one encountered ophiolites. This interpretation, when viewed in terms of Lugeon’s rule of undoing the Alpine deformation, namely, the higher a nappe in the edifice, the more southerly must be its original palaeogeographic provenance, led Argand (1916) to combine the ideas of Suess (1904, 1909) and Steinmann (1905) in the framework of his embryonic tectonics to interpret Alpine ophiolites as submarine effusions issuing along thrust faults into the active flank of an asymmetrically shortening geosyncline (Fig. 1). Like Argand, Leopold Kober, one of the most influential tectonicians of the first half of the 20th century, also followed the views of Suess and Steinmann in his very popular textbook Der Bau der Erde (1921, pp. 34–36). By 1928, when the second, enlarged edition of his book appeared, his enthusiasm for the close association of thrusts and the ophiolites seems to have somewhat waned, but, following Steinmann (1905, 1927), he still held fast onto their deep-sea origin.

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By 1922, however, Argand had changed his mind. He no longer viewed geosynclines as products of shortening, because he had become a follower of Wegener’s continental drift theory (Sengör, ¸ 1982a, 1982b, 1982c, 1998). In this new view, geosynclines marked sites of crustal stretching. When stretching went far enough the continental crust ruptured and the area of extension became underlain by the sima (Fig. 2). Argand saw the provenance of ophiolites in this simatic basement. He pointed out that the frequent association of the ‘green rocks’ with abyssal and bathyal sedimentary rocks could thus be explained. He also noted that sedimentary rocks of shallower provenance also appeared commonly mixed with these. Argand ascribed such mixing to submarine slumps (Argand, 1924, p. 299). Staub’s (1924, 1928) scheme (Fig. 3) was a direct derivative of Argand’s earlier interpretation (1916), but Staub, coming from an eastern Switzerland experience, was not a partisan of the eccentric position of ophiolite effusion channels with respect to the geosynclinal axis. Abandoning Suess and Argand, he liberally distributed them throughout the axial part of the Tethys. It had thus become common knowledge by the teens of the 20th century that ophiolites had something to do with orogeny. Kober (1921, p. 35) quoted Suess with emphasis that they always occurred within mountain belts and never on forelands. But again, not every orogenic belt had them: Steinmann (1905, p. 63) pointed out that in the Central Andes, for example, he had been able to find no ophiolites despite intensive search. He thus distinguished mountains of Alpine-type from mountains of Cordilleran-type. Only the former, which Steinmann thought encompassed an entire geosyncline, had ophiolites. The Cordilleran-type mountain belts, Steinmann speculated, may have been only the margin of a geosyncline or they my have grown out of shallow geosynclines. In a widely cited paper, Hess (1939) pointed out that the dominant member of the Steinmann trinity was serpentinites. He believed that they were intrusive, mainly disposed of in two parallel belts 60 miles distant from the axis of any orogen in which they are located, and that their time of intrusion always coincided with the first major deformation of the orogenic belt in which they occurred. He challenged the view that they originated in geosynclines by showing that they also occurred in island arcs, which Hess believed had no geosynclinal precursors. According to Hess, geosynclines formed only after the first major deformation in a given mountain belt and were essentially foredeeps. He was thus in agreement with Franz Eduard Suess (1937), Eduard Suess’ son. Another extremely infulencial tectonician of the first half of the twentieth century, Hans Stille, offered, in 1939, the first systematic account of ophiolite genesis and significance within the framework of his geotectonic cycle. According to him, ophiolites occurred during the preparatory phase of mountain-building, when the maternal geosyncline began subsiding. He called this ‘initial magmatism’ (he had thought of calling it ‘ophiolitic magmatism’, but had decided against it, because not all initial magmatism in geosynclines was exclusively ophiolitic: see Stille, 1939). He believed that they issued along the axes of subsiding geosynclines and there mixed with deep-sea sedimentary rocks (Fig. 4). The essence of Stille’s scheme was that ophiolites were confined to geosynclines. He thus simply followed his countryman Steinmann.

2. Ophiolites and Orogeny: A Historical Review and Critique

Fig. 1. Argand’s (1916, Pate II, Fig. 1) famous block diagram showing the conditions in the Alpine ‘geosyncline’ at the time of its pre-paroxysmal stage. Note the ophiolite ‘effusions’ à la Suess and Steinmann (v and v ). We have copied this translated version from Collet (1935, Plate I).

Fig. 2. Origin of geosynclines and oceans illustrated on the example of the Ionian Sea by Argand (1924, Figs. 19 and 19 bis). Note the surfacing of the simatic rocks at the toe of the continental apron and their subsequent incorporation into an orogenic belt during shortening. 683

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Fig. 3. Staub’s (1924, Fig. 61, 1928, Fig. 1) view of the origin of ophiolites in the Alpine Tethys.

In 1936, Stille had divided the North American Cordilleran geosyncline into a western and an eastern trough. The western trough was thought to contain thicker piles of dominantly clastic/silicic sedimentary rocks closely and abundantly associated with mafic and intermediate ‘initial’ volcanics (Fig. 5). The eastern trough was believed to hold a thinner pile of more carbonate rich sedimentaries and had no initial volcanics. Stille called the western trough ‘pliomagmatic’ and the eastern trough ‘miomagmatic’. In 1940, he called the pliomagmatic trough a eugeosyncline and the miomagmatic trough a miogeosyncline. Ophiolites occurred only in eugeosynclines. As the eu- and miogeosynclinal couples formed the great mother troughs of mountain belts, the orthogeosynclines, Stille thus ultimately conceded the eccentric position of ophiolite feeding channels within a large orthogeosynclinal system. By the end of the fifties, it had thus become clear that ophiolites (1) represented a mafic/ultramafic association commonly closely coupled with deep-sea sedimentary rocks, (2) they were confined to orogenic belts, and (3) they generally occurred as tabular bodies that came into being early during the life cycle of a mountain belt in a volcanically active part of the maternal geosyncline, but they were also present as smaller bodies in enormous piles of eugeosynclinal sedimentary rocks. Only Hess made the, in retrospect, critical observation that on continents although ‘alpine-type peridotites occur only in alpine-type mountain structures’. ‘They appear to occur ubiquitously throughout the earliest Precambrian rocks’ (Hess, 1955, p. 394). All of these observations, except the last one, were easily converted to plate tectonic interpretations, once it was recognised that tabular ophiolite bodies were in fact nappes and the vast thicknesses of eugeosynclinal sedimentary rocks with associated ophiolitic magmatics were tectonically stacked sequences in accretionary complexes. Papers by Gass and Masson-Smith (1963) and Dietz (1963), respectively, already had the essences of these ideas. 2.4. Ophiolites as Distributed Fragments in Continental Structure: Historical Perspective Argand (1924) already surmised that sialic rafts drifting in sima would accumulate at their prows oceanic and continental sedimentary and igneous material in part mixed by sedimentary gliding processes and in part by tectonism. He pointed out that the closure of former oceans north of Tibet must have given rise to complicated accretionary ‘cushions’ between colliding continents (‘bourrelet complexe’: see Fig. 6 herein). He thought that one such complexly folded cushion, combining the continental slope accumulation of the colliding

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Fig. 4. Sequential cross-sections to illustrate schematically Stille’s view of the ‘geotectonic cycle’ and accompanying phenomena, such as the ‘magmatological cycle’ (see Stille, 1940, esp. p. 21) drawn after his descriptions and sketches in a number of his publications to show the place of ophiolitic magmatism in his scheme (from Sengör, ¸ 1982a, Fig. 1.9). (a) Consolidated crust before a regeneration that will form new orthogeosynclines between cratons. (b) Preparatory phase of orogeny: during this ‘anorogenic’ time epeirogenic movements (caused by epeirogenic stresses, Se , which are allegedly much weaker than the orogenic stresses So ) give rise to geosynclines and geanticlines. In geosynclines, the base of the crust is pushed into progressively deeper regions in the earth. As a result of the accompanying increase in temperature, its base weakens and its effective thickness to carry stresses decreases. At this stage initial magmatism (simatic, mostly ophiolitic) loads the geosyncline with heavy magma. As a result of subsidence, the area of the geosyncline becomes more ‘mobile’ than the bordering geanticlines (cratons). Stille never explained where the simatic magmas come from and why. It is conceivable that he might have thought of the ophiolitic ‘feeders’ as dyke systems parallel with the shortening direction. In 1939, he pointed out that not all initial magmatism was mafic. He gave examples of intermediate even felsic ‘initial magmatic rocks’. His scheme would indeed allow such magma mixing in the geosynclinal subsidence stage. (c) During an orogenic phase stresses throughout the globe allegedly increase. The crust fails at its weakest point which is the geosyncline. The geosyncline becomes compressed between the two stable blocks (cratons). This, as a rule, gives rise to a symmetric orogen. As a consequence of shortening, the crust in an orogen thickens, its base melts and this gives rise to synorogenic plutonism, which is sialic, in contrast to the simatic initial magmatism.

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Fig. 5. Hans Stille’s view of the origin of ophiolites in the North American Cordilleran geosyncline (from Stille, 1936, Fig. 2, 1940, Fig. 60).

continents, now forms the material of the Kuen-Lun range (Fig. 6). Although he did not explicitly say so, from his description of the nature of continental slopes being regions of the mixture of simatic rocks, abyssal, bathyal and shallower sedimentary rocks, it is not unreasonable to assume that in the structure of the Kuen-Lun he expected to see shreds of the ‘sima’ embedded in his ‘complexly folded cushion’. How wide the ophiolite-bearing cushions may be depended on how big the ocean was, from which the cushion had been swept off. Argand did not further elaborate on this idea. In 1955, Hess made the seminal observation that alpine-type peridotites ‘appear to occur ubiquitously throughout the earliest Precambrian rocks’ (Hess, 1955, p. 399), as we have mentioned above. He thought that the processes which created that picture were fundamentally different from those which created the Phanerozoic orogen-bound peridotite belts. He thought that the Precambrian situation arose because entire areas of ocean had somehow been converted into continents. Dietz (1963) reinterpreted Hess’ observation by assuming that successive orogenies had incorporated ocean-floor igneous rocks into what he called the eugeosynclinal prisms along continental margins. Like Hess, he related this to J. Tuzo Wilson’s (1949) idea of peripheral growth of the continents. While Hess (1955) thought that Wilson’s mechanism probably contributed to the vertical growth of the continents, Dietz (1963) pointed out that tectonic patterns of Precambrian terrains clearly indicated lateral growth also. Both Hess and Dietz, thought, however, that the processes that generated Precambrian terrains were no longer operative either in kind (Hess) or in extent (Dietz). They thought so, because they were

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Fig. 6. Argand’s cross-section across the Tibetan high plateau showing part of the Tien-Shan, Serindia (the Tarim Basin), Kuen-Lun and the Himalaya, in addition to the high plateau itself. The bourrelet complexe (shown by dense stippling added to Argand’s figure by us) represents, according to Argand, an earlier accretionary complex that had formed long the southern margin of the ‘Serindian Block’ than the mainly Mesozoic one shown by Argand himself with sparse stippling (‘The tectonic products arisen from the axial zone of the Tethys’: Argand, 1924, p. 348). Modified from Argand (1924, Fig. 13).

not familiar with regions that resembled the Precambrian terrains they were considering. The world geological opinion has largely followed either Hess or Dietz until the beginning of the 1990s. The main purpose of this paper is to describe one major region in which ophiolites occur much as they do in Precambrian, specifically Archaean, terrains and to argue that the processes responsible for creating the tectonic patterns of these regions were substantially the same. 2.5. The Term Ophiolite: Can Its Original Usage Be Taken as a Guide to Its Present Usage? The term ophiolite was introduced by the great French geologist Alexandre Brongniart in 1813 (for a detailed history, see Amstutz, 1980). When first introduced, Brongniart used it essentially for rocks made up dominantly of serpentinite enveloping different species of minerals and he gave both compositional and textural criteria for their recognition. In 1827, Brongniart gave a much more extended discussion of the term ophiolite and there too he applied the term only to a variety of serpentinites associated with diverse types of minerals. In an 1821 paper on the northern Apennines, Brongniart described the ophiolites of what is now known as the Liguride units. There he showed the close association of serpentinites of diverse types, gabbros, mafic volcanics and cherts. Brongniart’s usage of the term ophiolite was strictly mineralogical/petrographical. It can be extended to cover the present usage only in retrospect, but not by considering Brongniart’s publications alone. The present usage (Penrose field conference participants, 1972) covers an association that goes from ultramafics to mafic volcanics. Brongniart’s term applies strictly to the ultramafic part. It is clear that the original usage as intended by Brongniart (1813, 1821, 1827) cannot be taken as a guide to the present usage. But confining the term ophiolite only to the Penrose definition (Penrose field conference participants, 1972) would restrict it to only a few cases where the Penrose sequence is pre-

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served intact. What would we then call the other remnants of ocean floors now embedded in the continents? We could naturally describe each according to its petrography, but descriptive terms are not efficient communicators of interpretations. Under what term would we unite separate slivers of ultramafics, gabbros, basalts, etc., if we believe they are were plucked off now vanished oceanic crust and upper mantle? Thus, by the end of the eighties, two major classes of oceanic basement fragments had been identified: one was the giant ophiolite nappes, which were clearly generated atop subduction zones in environments of pre-arc spreading and many had been obducted across former fracture zones. The other class was far more heterogeneous: it consisted mainly of a diverse array of oceanic basement scraps in mélanges representing a full spectrum of the ‘Penrose ophiolite definition’ (Penrose field conference participants, 1972), but, in many mélange belts, being dominated by the mafic end of the spectrum. The numerous ‘diabase-phyllitoid’ and ‘slate-diabase’ associations of the Tethysides are examples of such ‘mafic dominance’ (e.g., Yilmaz and Sengör, ¸ 1985). Dewey (2003) recently called the first class ophiolites sensu stricto and the second, ophiolites sensu lato. He further argued that ‘Ophiolites sensu lato can be thought of as a simple random record of the Wilson Cycle of opening, widening, narrowing, and closing oceans but ophiolite complexes sensu stricto are events related to particular and special tectonic configurations’ (Dewey, 2003). We think that Dewey’s (2003) suggestion of separating ophiolites sensu lato, including all occurrences of oceanic basement rocks, from ophiolites sensu stricto, i.e., those ophiolites showing the complete Penrose sequence is very apposite. However, it would be handier to find single-word designations for these two classes. We here suggest to keep ophiolite for those bodies preserving the complete Penrose sequence as has become customary since 1972. For separated fragments of this complete sequence, also including sea-mounts, oceanic plateaux, and fracture zone scarps, and dominated by basalts and gabbros, and rarer ultramafic rocks (Dewey, 2003) including both lherzolitic and harzburgitic types, we suggest the term ophirag, from the Greek oφιs (= serpent, snake) and ραγos (= tatter, shred, sliver). Thus we shall henceforth call ophiolites sensu stricto in Dewey’s sense ophiolites and his ophiolites sensu lato ophirags. Ophiolites and ophirags together make up remnants of oceanic basement rocks incorporated into continents. As we shall see below, by far the largest number of the Altaid oceanic basement remnants are ophirags; the same is true for most Archaean occurrences of oceanic basement remnants.

3. OUTLINE GEOLOGY OF THE ALTAIDS 3.1. Historical Perspective: The Problem of the Altaids The Altaids (Fig. 7) constitute one of the world’s largest and most complicated orogenic systems. Their name derives from the Altay Mountains in Russia, China and Mongolia. It was coined a century ago, in the first part of the third volume of Das Antlitz der Erde (Face of the Earth) by Eduard Suess (1901, pp. 246–250) to designate the mountain chains that formed around what he called the ‘ancient vertex of Asia’ (Fig. 8). He noted that the

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Altaids constituted a series of ever widening, north-concave arcs around the ancient vertex and were delimited to the south by the chains that arose from the mid- to late Mesozoic and early Cainozoic elimination of a broad Mesozoic ocean, the Tethys (cf. Sengör ¸ and Natal’in, 1996a; Sengör, ¸ 1998). Suess characterised both the structural coherence of the Altaids (disposition in more or less concentric arcs, lack of a foreland) and their compositional peculiarity (paucity of large gneiss terrains, dominance of slates, schists, cherts with some basalts and serpentinites) and likened them to the ‘waves in the open sea’ as opposed to ‘waves breaking on a beach’ that resembled more the mountain ranges with pronounced forelands such as the Alps and the Himalaya (Sengör, ¸ 1998, p. 248). In the 20th century, the greatest problem the Altaids posed has been their unusual shape and internal composition. They resembled none of the more familiar great mountain ranges of the globe such as those along the Alpine-Himalayan ranges or the Cordilleras in the Americas. That parts of them were like parts of the North American Cordillera or Japan was not considered significant. That they greatly resembled the Archaean and some Proterozoic regions was not at all recognised. On the one hand, such Archaean and Proterozoic terrains were thought not to have Phanerozoic counterparts and, on the other, the Altaids, which in reality were similar to those Precambrian areas, were thought unique. Complex models of geosynclines were proposed for the interpretation of the Altaids, but few could go beyond being ad hoc solutions (in this, the difficulties of interpretation were remarkably similar to those in the Precambrian regions; but again, this similarity was not recognised). The early plate tectonic models for the Altaids were also ad hoc (with the very remarkable exception of the classic Uralides paper by Hamilton, 1970), in that they invoked head-on collisions for any length of suture that could be identified with no obvious genetic connexion among the various colliding entities (e.g., Kropotkin, 1972; Zonenshain, 1973; Burrett, 1974). The neutral names used for Suess’ Altaids, such as the Ural-Amurian or the Ural-Mongolian foldbelt or the Central Asian foldbelt in the literature betray this interpretative sterility. Suess’ Altaids disappeared from the literature, along with the recognition of their structural and evolutionary coherence. 3.2. Present Interpretation of the Altaids The initial step in the rehabilitation of Suess’ Altaid model was the recognition that much of Central Asia, particularly the mountain ranges around the Tarim Basin and those extending into Kazakhstan consisted dominantly of subduction-accretion complexes (Sengör ¸ and Okuro˘gulları, 1991; Sengör, ¸ 1992) and that those north of the Tarim basin constituted a single, unified orogenic system (Hamilton, 1970; Sengör ¸ et al., 1993). There are very few and spatially restricted neat, long and narrow ophiolitic sutures and no forelands (except at the outermost periphery: see Fig. 7) against which parts of the orogen abut across a clear structural front except in the extreme south (Fig. 7). The shape of the orogen is irregular and does not have the familiar linear/ arcuate, long and narrow shapes of most Phanerozoic orogenic belts. The trend of the Altaid orogen within this confusing region could most easily be followed by tracing out the magmatic arc fronts (for methodology, see Sengör ¸ and Okuro˘gulları, 1991;

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Fig. 7. Generalised tectonic map of the Altaids and related surrounding tectonic units (from Sengör ¸ and Natal’in, 1996a, Fig. 21.18). The Uralide and Baykalide orogenic systems have not been further subdivided. Dotted line shows the limit of the post-Altaid cover. Key to numbers: 1. Valerianov-Chatkal (pre-Altaid continental basement, Altaid accretionary complex and magmatic arc), 2. Turgay (pre-Altaid continental basement, Altaid accretionary complex and magmatic arc, buried under 2-km-thick Mesozoic-Cainozoic sedimentary cover), 3. Baykonur-Talas (pre-Altaid continental basement, early Palaeozoic Altaid accretionary complex and magmatic arc), 4.1. Djezkazgan-Kirgiz (pre-Altaid continental basement, Altaid Palaeozoic magmatic arc and accretionary complex), 4.2. Jalair-Nayman (pre-Altaid continental crust, early Paleozoic marginal-sea complex, early Palaeozoic magmatic arc and accretionary complex), 4.3 (or 16) Borotala (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex), 5. Sarysu (Altaid accretionary complex and magmatic arc), 6. Atasu-Mointy (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex), 7. Tengiz (pre-Altaid continental basement, Vendian-early Palaeozoic magmatic arc and accretionary complex), 8. Kalmyk Köl-Kökchetav (Vendian-early Palaeozoic magmatic arc and accretionary complex), 9. Ishim-Stepnyak (pre-Altaid continental basement, Vendian-early Palaeozoic magmatic arc and accretionary complex), 10. Ishkeolmes (early Palaeozoic magmatic arc and accretionary complex), 11. Selety (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex), 12. Akdym (Vendian-early Palaeozoic magmatic arc and accretionary complex 13. Boshchekul-Tarbagatay (early Palaeozoic magmatic arc and accretionary complex), 14. Tekturmas (Ordovician-medial Palaeozoic accretionary complex, medial Devonian-early Carboniferous magmatic arc), 15. Junggar-Balkhash (early-late Palaeozoic magmatic arc, medial and late Palaeozoic accretionary complex), 16. see unit 4.3, 17. Tar-Muromtsev (early Palaeozoic magmatic arc and accretionary complex), 18. Zharma-Saur (Palaeozoic magmatic arc, early Palaeozoic accretionary complex), 19. Ob-Zaisan-Surgut (late Devonian-early Carboniferous accretionary complex, strike-slip fault-bounded fragments of the lte Devonian-early Carboniferous magmatic arc, late Palaeozoic volcanic arc), 20. Kolyvan-Rudny Altay (early Palaeozoic accretionary wedge, early and medial-late Palaeozoic magmatic arc), 21. Gorny Altay (early Palaeozoic accretionary complex and magmatic arc superimposed by medial Palaeozoic magmatic arc; in the South Altay sector, medial palaeozoic accretionary complex with fore-arc basin), 22. Charysh-Chuya-Barnaul (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex, medial Palaeozoic fore-arc basin and magmatic arc), 23. Salair-Kuzbas (pre-Altaid continental basement, Vendian-early Palaeozoic magmatic arc and accretionary complex, Ordovician-Silurian fore-arc basin, Devonian pull-apart basin, late Palaeozoic foredeep basin), 24. Anuy-Chuya (early Palaeozoic magmatic arc and accretionary complex), 25. Eastern Altay (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex with large seamount fragments), 26. Kozhykov (early Palaeozoic magmatic arc and accretionary complex), 27. Kuznetskii Alatau (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex), 28. Belyk (Vendian-medial Cambrian magmatic arc and accretionary complex), 29. Kizir-Kazyr (Vendian-medial Cambrian magmatic arc and accretionary complex), 30. North Sayan (Vendian-early Palaeozoic magmatic arc and accretionary complex), 31. Utkhum-Oka (pre-Altaid continental basement, early Palaeozoic magmatic arc and accretionary complex), 32. Ulugoi (Vendian-early Cambrian magmatic arc and accretionary complex), 33. Gargan (pre-Altaid continental basement, early Palaeozoic magmatic arc and Vendian-early Palaeozoic accretionary complex), 34. Kitoy (early Palaeozoic magmatic arc), 35. Dzhida (early Palaeozoic magmatic arc and accretionary complex), 36. Darkhat (pre-Baykalide continental basement, Riphean magmatic arc and accretionary complex), 37. Sangilen (Baykalide microcontinent that collided with Darkhat unit in the Riphean and experienced strike-slip displacement during the early Palaeozoic Altaid evolution), 38. Eastern Tannuola (early Palaeozoic magmatic arc and accretionary complex), 39. Western Sayan (early Palaeozoic magmatic arc and accretionary complex), 40. Kobdin (pre-early and medial Palaeozoic magmatic arc and accretionary complex), 41. Ozernaya (Vendian-early Cambrian magmatic arc and accretionary complex), 42. Han-Taishir (pre-Altaid continental basement, Vendian-early Cambrian magmatic arc and accretionary complex), 43. Tuva-Mongol (equivalent to the central and eastern parts of Suess’ ‘ancient vertex of Asia’; see Fig. 8 herein): 43.1. Tuva-Mongol arc massif (pre-Altaid continental crust and Vendian to Permian magmatic arc), 43.2. Khangay-Khantey (Vendian-Triassic accretionary complex and magmatic arc), 43.3. South Mongolia (Ordovician-early Carboniferous accretionary complex), 44. South Gobi (pre-Altaid continental basement, Palaeozoic magmatic arc and accretionary complex).

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Sengör, ¸ 1993; for detailed application to the Altaids, especially Sengör ¸ and Natal’in, 1996a, Figs. 21.19–21.25). It seemed that magmatic fronts jumped forward within growing subduction-accretion complexes. This growth also enlarged the continent. When this model was applied to the Altaids as a whole, it appeared that they consisted of the fragments of at most two large magmatic arcs (Fig. 7) stacked together dominantly by arc-subparallel strike-slip faults (both arc-slicing and arc-shaving varieties (Natal’in and Sengör, ¸ 1996): Fig. 9) and not by head-on collisions as previously assumed (Sengör ¸ et al., 1993; Sengör ¸ and Natal’in, 1996a). These strike-slip faults were seen to have enormous offsets along them, amounting to hundreds, in many cases more than 1000 km. One pair of them, in particular, the Irtysh and the Gornostaev, appear as broad zones of intense shear with offsets around 2000 km or more (Fig. 7). These two are veritable keirogens dividing the Altaid orogenic system into two main domains: The Kazakhstan-Tien Shan domain in the west and the Altay-Mongol domain in the east. The Altay-Mongol domain itself is naturally divided into two by the continental Tuva-Mongol axis (unit 43.1 in Fig. 7) that encloses the Khangai and Khantey mountains (forming unit 43.2 in Fig. 7). The accretionary complexes lying west and northwest of the Tuva-Mongol continental axis form the Altay-Sayan sector and the Tuva-Mongol fragment itself, together with the Palaeozoicearly Mesozoic accretionary complexes forming the Khangai-Khantey mountains, southern Mongolia (unit 43.3. in Fig. 7) and the Gobi (unit 44 in Fig. 7), constitute the MongolOkhotsk sector (Fig. 7). Although the Tuva-Mongol continental fragment is the only major Precambrian core in the Altay-Mongol domain of the Altaids, there are at least 9 substantial pre-Altaid continental slivers in the Kazakhstan-Tien Shan domain (Fig. 7). These follow the disposition of the magmatic arc fronts here and they are separated from one another by subduction-accretion complexes with age ranges from Vendian to Carboniferous. These

Fig. 8. Tectonic map of a part of Central Asia showing Suess’ ‘ancient vertex of Asia’ and associated units after Obrutschew (1926, Plate 11). Suess never presented a map of the ancient vertex. Obrutschew was one of his most loyal followers in the definition of it. For a description of the ancient vertex of Asia, see Suess (1901, ch. 3) and Obrutschew (1926, pp. 15–17). Also see Stille (1958, pp. 61–66). For Obrutschew’s loyalty to Suess’ interpretation of the ancient vertex as a Precambrian unit, see Stille (1958 p. 65). Stille wrote (1958, p. 64) with emphasis, and ‘against the divergent opinions cited in the literature’, that Suess himself had viewed the ancient vertex as a younger structure than the Angaran Shield. It is unclear why Stille thought so (he indicates no specific writing by Suess to support his statement), because Suess expressly says that the ancient vertex is a high-lying remnant of the Angara Shield: ‘The larger part of the western half of the vertex was faulted down at an early period, and thus gave rise to the amphitheatre of Irkutsk’ (Suess, 1901, p. 98). Suess correctly recognized that what is now called the Tuva-Mongol continent (Sengör ¸ et al., 1993; Sengör ¸ and Natal’in, 1996a; see Fig. 7 herein, unit 43) was no younger than the basement of Angara Shield, even including Archaean rocks. What he did not know was that the Khangay-Khantey accretionary complex reached in age from the Vendian into the Mesozoic (cf. Sengör ¸ and Natal’in, 1996a).

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Fig. 9. Arc subparallel ‘trench-linked’ (Woodcock, 1986) strike-slip faults and their effects on the geometry of arc and subduction-accretion systems. (A) Arc slicing faults cut through the entire arc system and are proper plate boundaries. They may repeat or elide arcs. Their best active examples are seen in the Philippine Fault System (Rutland, 1968; Cardwell et al., 1980; Hall, 2002), in the Palu Fault that slices through the northern arm of Sulawesi (Celebes) (Cardwell et al., 1980), and in the northern Caribbean (Mann et al., 1991). In the Sangilen unit (unit No. 37 in Fig. 7; ophiolite No. 10 in Fig. 14), arc repetition created an ophiolitic zone that resembles a collisional suture, but is in reality a transform suture, exhibiting ophiolites entrapped by an arc repeating strike-slip fault. (B) Arc shaving faults affect only the fore-arc and shave bits of the fore-arc and transport them coastwise along the trench. The great Atacama fault in the Andean fore-arc is the best-known example (Sengör, ¸ 1990). Note how the forearc area is repeated and thus enlarged in front of arc-shaving faults and are elided and narrowed behind them. This is how the Altay region has been vastly enlarged, while South Mongolia was constantly being shaven along the immense Irtysh strike-slip system. In some cases hybrid faults originate. The Great Sumatra fault (Sieh and Natawidjaja, 2000), for example, mostly shaves the arc, but does so along the volcanic axis and in places wanders behind the axis to incorporate arc magmatics into the fore-arc ‘sliver plate’ (Jarrard, 1986). Some of the Altaid trench-linked strike-slip systems in the Altay-Mongol domain were of this hybrid type. If pull-apart geometries originate along any of such trench-linked, arc-subparallel strike-slip faults, new oceanic crust may be formed along short spreading centres trending at high angles to the arc, as, for example, in the Cayman Trough or in the Andaman Sea. The Han-Taishir ophiolite in the unit 42 of the Altaids, for example, may have formed in such an environment.

arc segments originally formed a single arc along the Ural/Yenisey margin of the then combined Russian and Angara cratons (Fig. 10). This arc became detached in the Vendian/Cambrian and formed what we call the Kipchak Arc (Fig. 11). In the internal parts of the Kazakhstan-Tien Shan sector, the ages of the accretionary complexes do not reach beyond the Ordovician-Silurian. It looks as if by Devonian time this part was wholly assembed by strike-slip stacking of the Kipchak Arc (Fig. 12; Sengör ¸ et al., 1993; Sengör ¸ and Natal’in, 1996a, 1996b). Similarly the Altay-Sayan sector was also internally assembed by the end of the Silurian (Sengör ¸ et al., 1993; Sengör ¸ and Natal’in, 1996a, 1996b). The internal Kazakhstan-Tien Shan sector formed a small continent by itself (with a narrow

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Fig. 10. Vendian palaeotectonic reconstruction of the Altaids (630–530 Ma). The legend shown here applies to all reconstructions in Figs. 10–16.

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Fig. 11. Early Cambrian palaeotectonic reconstruction of the Altaids (530–520 Ma). For legend, see Fig. 10.

attachment to the Russian craton via the Valerianov-Chatkal arc) whereas the Altay-Sayan sector was assembled as a tight attachment of the Angara craton and its Baykalide frame. These two sectors were initially united by an ensimatic segment of the Kipchak Arc now making up the Boshchekul-Tarbagatay and Zharma-Saur arc fragments (Sengör ¸ and Natal’in, 1996a). This ‘bridge’ eventually formed the backstop in front of which the enormous flysch wedge of the Junggar-Balkhash accretionary complex accumulated (Fig. 13). The Boshchekul-Tarbagatay/Zharma-Saur ‘bridge’ had a sliding attachment to the periAngara accretionary systems. As the bridge tightened in the form of a south-concave pincer, its Angara end moved along the previously-assembled units in a right-lateral sense along the Irtysh shear zone. At the same time the Valerianov-Chatkal unit was moving in a left-lateral sense along the eastern margin of the Urals forming the Denisov-Oktabyarsk suture. The driving motive of these movements were the rotation of the Angara and the

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Fig. 12. Early Devonian palaeotectonic reconstruction of the Altaids (420–390 Ma). For legend, see Fig. 10.

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Fig. 13. Early Carboniferous palaeotectonics of the Altaids (350–325 Ma). For legend, see Fig. 10.

Russian cratons (that had separated from each other in the Vendian) towards one another that progressively narrowed the oceanic area between them. This motion culminated in the late Carboniferous by the elimination of all marine areas from the Kazakhstan-Tien-Shan domain and the Altay-Sayan sector (with minor exceptions in the western Tien-Shan). The Tarim fragment had collided with the Valerianov-Chatkal, but mainly with the DjezkazganKirgiz unit (Sengör ¸ and Natal’in, 1996a) earlier to terminate the orogenic evolution of the Kazakhstan-Tien Shan segment (Fig. 14). In the late Carboniferous and the early Permian the Kazakhstan-Tien Shan domain (and, with it, the Russian craton) moved right-laterally for some 2000 km with respect to the Altay-Mongol domain and the Angara craton along the Irtysh keirogen. This motion created the Nurol extensional basin north of the Kolyvan Mountains (Fig. 15). In the late Permian a major reversal of this motion occurred along the Gornostaev keirogen. This left-lateral motion generated the major extensional basins of Nadym in the north and the Junggar-Alakol-Turfan system in the south (Allen et al., 1995;

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Fig. 16). The main tenets of the evolution here outlined and graphically depicted in Figs. 10–16 have been corroborated to a remarkable degree by the recent palaeomagnetic work by Rob Van der Voo and his colleagues in the Tien Shan and in the Kazakhstan orocline (Collins et al., 2001; Bazhenov et al., 2002, 2003; Van der Voo et al., 2002). The Mongol-Okhotsk sector continued its evolution until the Jurassic. The Solonker Ocean separating it from the Manchurides in northeastern China and the Russian Far East remained open until the early Permian (Fig. 15). The final closure and closure-related easterly extrusion of part of the Khangai-Khantey, including segments of the Tuva-Mongol arc massif (Bindeman et al., 2002), was mainly a result of the Cimmeride collisions in China and lasted into the Jurassic (Sengör ¸ and Natal’in, 1996a). The Altaid evolution displays an immense spatial and temporal continuity rarely seen in this clarity by the geologist in the evolution of a continent by orogeny. The Altaid orogeny was locally and intermittently arrested by massive keirogenic movements and local taphrogeny associated with it. The present state of the orogen is one of considerable disruption into innumerable little mosaic pieces, but it is the former continuity of its tectonic units that enables us to understand its structure and to reconstruct its evolution. This evolution betrays 0.4 million km3 /a continental growth during the evolution of the Altaids, which may account for 40% of the total Palaeozoic growth (Sengör ¸ and Natal’in, 1996b). This estimate has been largely corroborated by new geochemical work (Chen et al., 2000; Jahn, 2002) and it had been inferred in a different tectonic framework by Peive and his collaborators (Peive et al., 1972). Another important observation is the lack of any episodicity in the Altaid evolution as a whole. Changes in structural regimes such as switch from subduction to strike-slip or local collisions as seen in the western Sayan mountains did introduce local episodicity, but there never was any cessation of motion within the orogen as a whole accommodating the continuous approach of the Russian and the Angaran cratons. Although most of the high-pressure metamorphics and backstop ophiolites seem to have been generated in the early history of the orogenic system, nothing comparable with the Tethyan ‘episodes’ of high-pressure metamorphism or ophiolite obduction is seen in the Altaids. This is probably a consequence of the lack of continental collisions and the great size of the ocean consumed in the Altaids, as opposed to the Tethysides: the small size of the ocean in the Tethysides dictated a narrow time interval for the subduction-related metamorphism and the arrival at subduction zones of long continental margins triggered nearsynchronous obductions of ophiolites along considerable strike-length. The abundance of ophiolite preservation in the earlier history of the orogenic system is probably a consequence of flysch-choking of the subduction zones that hindered plucking off ocean-floor pieces as seen in the still-active Makran accretionary complex in Iran and Pakistan, where the preserved ophiolites and ophirags are almost entirely of Cretaceous age (McCall and Kidd, 1982). The Altaid evolution has an actualistic analogue in the present-day western Pacific and it also greatly resembles the Pan-African evolution in North Africa and Arabia (Sengör ¸ and Natal’in, 1996b). It is also similar in style to most of the Archaean granite-greenstone terrains (Sengör ¸ and Natal’in, 1996b; Burke, 1997). The only difference we see between the Archaean terrains and the Altaid evolution seems a consequence of the higher oceanic

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Fig. 14. Late Carboniferous palaeotectonics of the Altaids (325–300 Ma). For legend, see Fig. 10.

3. Outline Geology of the Altaids

Fig. 15. Early Permian palaeotectonics of the Altaids (300–250 Ma). For legend, see Fig. 10. 701

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Fig. 16. Late Permian palaeotectonics of the Altaids (250–245 Ma). For legend, see Fig. 10.

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geothermal gradients in the Archaean. The immense flattening of supracrustal units such as pillow-lavas observed in some Archaean terrains (0
k
1) accompanied by amphibolitegrade metamorphism (Dr. John Myers, personal communication, 1999) indicates considerable extension in a hot environment, the equivalents of which have not yet been reported from the Altaids.

4. OCEANIC BASEMENT FRAGMENTS IN THE ALTAIDS: SOME REPRESENTATIVE EXAMPLES In the Altaids all known types of oceanic basement fragments, except large pre-terminal ophiolite nappes similar to those in Oman, Papua New Guinea or Newfoundland (contrary to certain claims such as those of Buchan et al., 2001: see below), occur in a wide variety of tectonic positions. Altaid ophiolites and ophirags may be classified according to the tectonic environments in which they now occur in the following manner: (1) Ophiolites that occur as arc basement of ensimatic arcs; (2) Ophiolites and ophirags that occur in former forearcs now entrapped within transform sutures: (a) Ophiolites as backstop to accretionary wedges; (b) Ophirags within accretionary wedges; (3) Ophiolites and ophirags in collisional suture zones. In the Altaid edifice, classes (1) and (2a) may in places include the same body. In very rare cases a single body may fit all three classes except (2b) because of the peculiarities of the Altaid evolution. In the descriptions below there are examples of such bodies that fit more than one category. It is impossible to give an exhaustive description or even a simple list of all occurrences. There must be vast numbers of individual bodies located in the three tectonic environments indicated above. It is impossible to walk a long profile across the Altaids without encountering some ophirag or ophiolite. Instead we have chosen a few examples from among these myriads to highlight their characteristics. In our choice we have been guided not only by the mere availability of the data, but also according to how well each locality is known by the international community. We chose those that have been visited by the largest number of geologists both from local countries and from afar. Among those we have preferred those that have been recently (re)described in internationally widely circulated journals, preferably in English, so that our preferred interpretations may be compared with existing ones in the literature by the widest possible readership. 4.1. Ophiolites that Occur as Arc Basement of Ensimatic Arcs Some ophiolites in the Altaids form basements of ensimatic arcs representing upper plate oceanic basement of former Altaid intraoceanic subduction zones. We mention below those in units 10, 13, 25, 26 and 42 (Figs. 7 and 17). This naturally does not mean that they do not occur in other ensimatic arc remnants. We do not mention them here, because some

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are now buried under younger cover and their presence is only surmised by geophysical means (especially under the West Siberian Basin: see Fig. 7) and about some others we simply have not yet come across an adequate publication. Yet others, as in the ZharmaSaur accretionary complex (unit 18 in Fig. 7), are so disrupted by later tectonism that it is difficult to tell whether they represent a once-coherent ophiolite disrupted in situ or ophirags assembled together from disparate areas in an accretionary complex mainly by arc-shaving strike-slip tectonism. In the following accounts all unreferenced geological description is from Sengör ¸ and Natal’in (1996a). Unit 10 (Ishkeolmes). This is a remarkable ensimatic arc, sitting on the Middle Cambrian Kuyanbai ophiolite complex (1 in Fig. 17). The ophiolite is topped by komatiitic basalts and boninites described along the western side of the unit (Spiridonov, 1991), i.e., near the fore-arc/arc contact. West of it is an accretionary complex including Ordovician cherts, shales and Middle to Upper Ordovician flysch. Unit 13 (Boshchekul-Tarbagatay). This is a middle Cambrian to Silurian arc system constructed atop an ophiolite unit yielding isotopic ages between 568 and 525 Ma. Clearly a 43 Ma spreading history may be implied or the unit may have brought together complete ophiolite fragments of different ages. Parts of the ophiolite are metamorphosed giving rise to garnet amphibolites. Atop this ophiolite are volcanic rocks consisting of rhyolites and dacites and, together with these arc magmatics, the ophiolite now builds the arc massif of the Boshchekul-Tarbagatay arc system (Peive and Mossakovsky, 1982; Khromykh, 1986; Borisenok et al., 1989; Yakubchuk and Degtyarev, 1991), very similar to the present-day ensimatic arc of the Marianas. The arc produced in the late Ordovician magmatic rocks as evolved as granodiorites, similar to those known from the ensimatic arc of the KingsKaweah ophiolite in California (Jason B. Saleeby, personal communication, 2002). Collins et al. (2002) corroborated the ensimatic arc interpretation by pointing out that in this unit the sediments had not been fed from any pre-existing continental source. Unit 25 (Eastern Altay). This unit has boninites as the lowest unit (Simonov et al., 1994; Berzin and Kungurtsev, 1996) and Dobretsov and Buslov (in press) believe that they sit on an ophiolite basement (2 in Fig. 14). The boninites are of Vendian-earliest Cambrian age. The arc evolved to a calc-alkalic composition in the early to medial Cambrian. Unit 26 (Kuznetskii Alatau). The Kuznetskii Alatau range has an ophiolite basement (3 in Fig. 14) on which supposedly the best-known ensimatic arc of the entire Altaid edifice was constructed (Dr. Alexander V. Vladimirov, personal communication, 2002). The basement is of late Precambrian age and the arc constructed on top of it has an age range from

Fig. 17. Some of the major ophiolite and ophirag occurrences in the Altaids. The numbers are referred to in the text. A vast number of smaller occurrences could not be shown at the chosen scale. Also not shown are those under the cover of the West Siberian Basin, which are known through drilling and geophysical means (for the extent of the post-Altaid cover, see Fig. 7). Compare this map with the maps of Precambrian greenstone belts in the end-papers in de Wit and Ashwal (1997).

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the Vendian to the Middle Cambrian, separated into two sequences by a Lower Cambrian reefal limestone unit (Volkov, 1988; Kungurtsev, 1993). Unit 42 (Han-Taishir). This unit has in its northern half a pre-Altaid piece of continental crust dated by single zircon evaporation ages on a xenocryst as 1715 Ma. The granite carrying the xenocryst was emplaced at 1127 Ma (Kröner et al., 2001). To this basement is attached the complete Han-Taishir (transliterated as Khantaishir in some publications) ophiolite (4 in Fig. 17). In fact, the ophiolite has been pieced together from various separate fault-bounded blocks and inclusions in a serpentinitic mélange (Khain et al., 2003). Several generations of dykes, sills and gabbro intrusions indicate a protracted history of the ophiolite formation that in turn indicates formation of the ‘ophiolite’ in different tectonic settings. Boninites have been reported from the sheeted dyke complex, although it is impossible to guess from the descriptions to which generation of dykes the boninites actually belong (Simonov et al., 1994). The plagiogranites from the ophiolite gave a zircon age of 568 Ma (Khain et al., 2003). Khain et al. (2003) interpret the Han-Taishir ophiolite to have resulted from forearc spreading, whereas Zonenshain and Kuzmin (1978) had originally interpreted the geochemical data from these ophiolites to indicate a back-arc basin origin. Taking into account the complex overprinting relationships in the dyke complex, its protracted history and its proximity to the continental margin, we consider the Han-Taishir ophiolites to have formed in an Andaman Sea-type situation with spreading next to the continent which shortly later nucleated on top an ensimatic arc that provided the boninites. The presence of the Neoproterozoic (?) Shargyngol dyke complex in the pre-Altaid continental piece (Badarch et al., 2002), if indeed of appropriate age, may be a part of this margin-parallel shear regime. 4.2. Ophiolites and Ophirags that Occur in Former Forearcs Now Entrapped within Transform Sutures Ophiolites as Backstop to Accretionary Wedges Unit 37 (Sangilen). This unit consists of a Baykalian continental nucleus and an Altaid accretionary complex that grew in front of it (present north). The Agardag ophiolites to the north (5 in Fig. 17), recently have been dated as 579 Ma by Sm/Nd (Pfänder et al., 2002). They occur in blocks embedded in a mélange, where previous studies have resulted in the recognition of several types of rock associations differing in mineralogical composition and major-element geochemistry among the gabbros (Izokh et al., 1988). Sedimentary rocks in the mélange are represented by schists, sandstones, cherts and carbonate rocks. Pillow basalts, dykes and microgabbro from a region (7.5 × 20 km) stretching parallel with the mélange zone reveal geochemical and isotope features similar to OIB, island arc and back-arc tectonic settings. Pfänder et al. (2002) have defined the nature of the Agardag ophiolite on the basis of the island arc and back-arc signatures. As a result of their study, the whole of the forearc region of the Sangilen unit (at most 40 km wide) has been interpreted as an independent back-arc/fore-arc Agardag terrane (Badarch et al., 2002), the tectonic connexions of which, as in most ‘terrane interpretations’, are entirely obscure. The close association of the island arc and back arc lavas now separating the accretionary com-

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plex from the continental nucleus of the Sangilen unit leads us to interpret the ophiolites, in the light of our large-scale Altaid tectonic model (Sengör ¸ and Natal’in, 1996a), as the product of a pre-arc spreading event in front of the continental nucleus of the Sangilen unit and now forming a backstop to the accretionary complex, whose age ranges from Vendian to the early Cambrian. It is thus similar to the Coast Range ophiolite in California and it is, like the Coast Range ophiolite south of San Francisco (W.R. Dickinson, personal communication, 2001, and in preparation), in places highly disrupted. Unit 30 (North Sayan). The North Sayan unit was indicated to be an ensimatic arc in Sengör ¸ and Natal’in (1996a). Indeed, the Vendian (?) through Lower Cambrian andesites, basalts, dacites, hyaloclastites, tuffs, cherts and reef limestones constitute the arc massif. In the accretionary wedge to the (present) south of this arc, in the Boruss mélange belt (6 in Fig. 17), Simonov et al. (1994) and Dobretsov et al. (1995) report ophiolites with geochemical signatures indicating island arc boninites and back-arc basin basalts. These ophiolites are bereft of a sheeted dyke complex. We interpret these sequences as parts of an ophiolitic forearc in front of which an accretionary complex had developed that include high-pressure schists, metapeiltes, metacherts, marbles and greenschists with ages around 450 to 400 Ma. The structure of this forearc complex is now much disrupted owing to considerable coastwise transport and oroclinal bending of the entire Western Sayan Mountains (including the Northern Sayan (30 in Fig. 7) and the Western Sayan (39 in Fig. 7) units. Unit 39 (Western Sayan). This unit also consists of an ensimatic arc massif and a large accretionary wedge. The arc basement is exposed to the west of the unit and consists of island arc-boninite-bearing Kurtushiba ophiolite (7 in Fig. 17). The ophiolite consists of a package of nappes thrust onto glaucophane-bearing accretionary complex rocks. The lower part of the package consists of three nappes each of which being represented by volcanic and sedimentary rocks. Their thickness individually ranges from 1.5 to 3 km. It is clear that with such thicknesses, if they are indeed stratigraphic, we cannot have a normal oceanic crustal succession. It is probably some sort of a pre-arc spreading product. Above is an ultramafic/gabbro/diabase thrust sheet that is 7 km thick. Its internal structure is said to be ‘weakly disturbed’. The nappe includes metamorphic peridotites, cumulate ultramafics and gabbros followed by isotropic gabbros and diabases and volcanic rocks. The diabase section contains sheeted dykes in the central part of the ophiolite outcrop and sills in its northeastern sector. The entire succession is terminated by Lower Cambrian volcanic rocks of island arc type forming a small area in the central and southern part of the ophiolite belt (Berzin and Kungurtsev, 1996). Farther east are coeval arc magmatic rocks including basalts, andesites, felsic volcanics, tuffs and conglomerates and reef limestones. Westwards, shallow-marine rocks of the forearc basin appear. Within the forearc area are MORB-type ophirags that probably represent offscraping from the downgoing slab. The Kurtushiba ophiolite exposed along the western margin of the arc massif thus resembles, in its tectonic position, to the ophiolite believed to underlie the Great Valley Sequence in California.

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Ophirags Within Accretionary Wedges Ophirags occurring as accretionary wedge scraps are the most common types of oceanic basement fragments preserved in the Altaid edifice. Our Fig. 17 shows those that can be shown at the chosen scale. Naturally, there are many more in all of the Altaid accretionary complexes. In the following, we discuss only those that have been recently studied in terms of their regional geology, petrology and geochemistry. Unit 4.1 (Djezkazgan-Kirgiz). In this unit, the ophirags of Kara-Archa (8 in Fig. 17), Kenkol (9 in Fig. 17), Toluk (10 in Fig. 17), Karachi-Karakty (11 in Fig. 17), Karadzhorgo (12 in Fig. 17), and Archaly (13 in Fig. 17), have been studied geochemically with a view to determining their tectonic setting of origin. Fossil finds in the associated pelagic sedimentary rocks indicate that they are of Cambrian-Middle Ordovician age. Basaltic rocks associated with these ophirags are mostly of normal MORB type, but some, such as Toluk, reveal E- and T-MORB features indicating ocean island remnants (Lomize et al., 1997). In Toluk and Kaarchi-Karakty there are basalts exhibiting arc-related features; therefore they have been compared with SSZ ophiolites. The study of spinel from the peridotite in the KarachiKarakty body also indicates mid-oceanic and supra-subduction zone tectonic settings (Demina et al., 1995), possibly indicating a pre-arc spreading event. In the Kara-Archa body, arc-related Ordovician volcanic rocks overlie the ophiolitic pseudostratigraphy. This region has been interpreted as an ensimatic arc. Lomize et al. (1997) thought that the arctrench system of the Djezkazgan-Kirgiz unit stopped operating in the Caradocian. However, there is clear evidence to the south of these ophirags that the subduction-accretion complex continued its growth until the early Carboniferous (Sengör ¸ and Natal’in, 1996a). The Southern Tien Shan Accretionary Complex in China. Farther east, in China, a piece of accretionary complex formed south of several units of the Altaids in the KazakhstanTien Shan domain after these units had been assembled by strike-slip stacking by the early Devonian (Figs. 7 and 12). This late-forming unit was left nameless by Sengör ¸ and Natal’in (1996a). The Youshugou ophirag (14 in Fig. 17) in the southern Tien Shan occurs in this nameless accretionary complex and is of pre-Middle Devonian age. Its mafic dykes display an REE pattern interpreted to indicate mixed magma sources (Allen et al., 1992). A first group of magma is characterised by depleted HFSE, characteristic of supra-subduction ophiolites. A second group has enriched HFSE content that is typical for off-spreading sites such as seamounts. However, the sampling is from a mélange and the two separate magma sources may in fact indicate two separate bodies of ophirags stemming from two contrasting tectonic setting now mixed in the mélange (Allen et al., 1992). The rest of the known southern Tien Shan ophirags in China (15 in Fig. 17) are reported in the Changawuzi, Kule, Heiyinshan and Kumishi regions. Their basalts are of MORB type and Chen et al. (1999) consider them as having originated in mid-ocean ridge environments. An Ar/Ar plateau age from a pyroxene from the gabbro section in one of these bodies gives an age of 439.4 ± 26.7 Ma (Llandoverian). Blueschists from the mélange yield 351 ± 2 Ma (Famennian). The nameless accretionary complex of the southern Tien-Shan clearly has accumulated pieces of oceanic basement from diverse environments. Normal oceanic crustal fragments plus ocean islands were clearly swept into the subduction zone in front of it. Some of the

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supra-subduction zone signatures may indicate Mariana-type outbowed migratory arcs or they may have been like the eastern Aleutians barring embayments formed by indentations in the pre-Altaid continental piece that forms the backstop of the unit. The present state of knowledge about its structural geology unfortunately does not allow us to choose among these possibilities. Unit 14 (Tekturmas). Ophirags to the southeast of the Tekturmas unit have been recently studied from the viewpoint of their regional geology, petrology and geochemistry by Wang et al. (in press). These are the Dabut (16 in Fig. 17) and the Honguleleng (17 in Fig. 17) ophirags. Darbut has a Sm-Nd age of 395 Ma (Emsian) and is closely associated with Ordovician radiolarian cherts in the subduction mélange of the Tekturmas unit. The geochemical discrimination diagrams indicate that the Dabut ophirag represents a MORB-type ophiolite later enriched in an ocean island setting probably by plume contribution. It may be that the original ocean floor was of Ordovician age onto which the later hot-spot edifice was built. It is clear that the Dabut ocean island was clipped off the downgoing plate and added as an ophirag into the subduction mélange. The Honggueleng ophirag also has MORB signatures, but also island arc signatures. If one considers the reconstruction shown in Fig. 13, the reason for this becomes readily apparent. Pieces of unit 18 (Zharma-Saur) representing an ensimatic Ordovician island arc were fed laterally into the accretionary complex of unit 14 by arc-shaving strike-slip faults to become ophirags within it. Thus, ocean floor ophirags and island arc ophirags were brought into close proximity within an accretionary complex by means of arc-shaving strike-slip faults. Unit 15 (Junggar-Balkhash). In the interior position of the Junggar-Balkhas unit, Wang et al. (in press) studied the Karameili (18 in Fig. 17) and the Mayila (19 in Fig. 17) ophirags. Of these, the Karameili ophirags give a pure MORB signature and they are clearly just plucked off pieces of the downgoing oceanic crust. The Mayila ophirags are also of MORB-type, but they are enriched by a plume source and probably are pieces of former ocean islands. Their tectonic history was probably not dissimilar to those of the Dabut and the Honggueleng complexes (compare Figs. 13 and 17). Unit 19 (Ob-Zaysan-Surgut). The Ob-Zaysan-Surgut unit is one of the largest tectonic units of the Altaids. It was transported from the present-day southeastern Mongolia to its present position along the giant Irtysh keirogen. The Aermentai ophiolites are located within this unit (20 in Fig. 17) and give a Sm-Nd isochron age of 561 ± 41 Ma (latest Precambrian or earliest Cambrian) according to Hu et al. (2000). By contrast, Wang et al. (in press) report a Sm-Nd age of 479 ± 27 Ma (Arenigian), measured by Liu and Zhang (1993) on the same ophiolite. A whole rock isochron age of the cumulate gabbro, diabase and andesite is said to give 561 ± 41 Ma (Huang et al., 1997). It seems that the ophiolite is probably Cambro-Ordovican in age and gives geochemical signatures indicating an ensimatic island arc setting. This inferred setting is consistent with the reconstruction given in Fig. 11, except that the arc was probably formed offshore of the Tuva-Mongol fragment and was later moved by coastwise transport to its present position as suggested by Sengör ¸ and Natal’in (1996a) along arc-slicing strike-slip faults. It was during this process that they

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become incorporated into the large accretionary complex of the Ob-Zaysan-Surgut unit. The presence of early Devonian adakites not too far away from this ophiolite complex support their location in a setting not too dissimilar to the Aleutian arc system exhibiting considerable strike-slip. Unit 25 (Eastern Altay). Dobretsov and Buslov (in press) have claimed that more than 50% of the Eastern Altay unit is made up of stranded seamounts of considerable size, in fact, fragments of a veritable plateau. However, during the IGCP 283 international excursion in 1993 we noted that the mafic rocks and their shallow-water limestone cover appeared in extremely deformed ophirags incorporated into a vast subduction mélange terrain. Much of the mélange Dobretsov and Buslov (in press) interpret as debris flows accumulated along the aprons of seamounts. This naturally gave the impression that seamounts (or an oceanic plateau) had been preserved entire and had, in fact, choked and stopped the subduction. However, we thought that their ‘olistostromes’ were, in fact, tectonic mélanges. Associated pillow lavas were highly deformed and dismembered, indicating large amounts of internal strain. Therefore, the seamounts themselves were deformed and dismembered and incorporated into a mélange wedge together with normal MORB-type basalts, boninites, high pressure rocks and pelagic sedimentary rocks such as cherts during active subduction. What the eastern Altay unit exhibits is nothing more than diverse types of ophirags stemming from different provenances now embedded in a subduction-accretion mélange. Unit 35 (Dzhida). Small ophirag lenses (22 in Fig. 17) are scattered in the accretionary wedge of the Dzhida unit. Stratigraphic estimates limit their age to Vendian-Early Cambrian time. Volcanic rocks of these ophirags reveal an oceanic signature, and they have been interpreted as relicts of oceanic crust (Gordienko, 1987; Kuzmin et al., 1995; Parfenov et al., 1995). Some subalkalic basalt may represent fragments of seamounts. To the south of the accretionary wedge there is a body of an ensimatic magmatic arc of Vendian-early Cambrian age in which felsic magmatism was active till the end of the Cambrian. A Neoproterozoic ophiolitic basement is inferred for this arc (Badarch et al., 2002) though direct evidence for this suggestion is lacking. The accretionary wedge/magmatic arc disposition as well as the nature of ophirags clearly indicate a north facing for the arc and several recent reconstructions have so portrayed it (e.g., Parfenov et al., 1995; Sengör ¸ and Natal’in, 1996a). However in a more recent paper the Dzhida arc has been shown as a south-facing arc without any compelling reason (Badarch et al., 2002). Unit 43.2 (Khangay-Khantey). Khangay-Khantey is a vast collage of accretionary complexes of various ages including various ophiolites and ophirags. In fact, there is no intact ophiolite in this entire zone. However, in the Bayankhongor area a large ophiolitic outcrop extends for 300 km in a NW-SE orientation exhibiting a serpentinitic mélange including a complete ophiolite suite but in form of collections of ophirags (21 in Fig. 17: Buchan et al., 2001, 2002). A 569 ± 21 Ma Sm-Nd age on pyroxene and whole rock from ophiolitic gabbro (Kepezhinskas et al., 1991) indicates a late Precambrian (Vendian) age for this ‘ophiolite’. Khain et al. (2003) reported a similar age (571 ± 4 Ma). The ‘ophiolite’ is interpreted to have formed as a result of sea-floor spreading on the basis of enriched MORB signatures (Kepezhinkas et al., 1991; Badarch et al., 2002; Buchan et al., 2001, 2002). The

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shallow water sedimentary rocks covering it may indeed indicate a plume environment for its original production and may render a clue to its preservation. The presence of extensive serpentinitic mélanges in it is very similar to the Kings-Kaweah ophiolite in California and may indeed suggest a fracture zone origin that originally may have localised an ocean island similar to the Coastal Complex in Newfoundland (Karson and Dewey, 1978). Buchan et al. (2001, 2002) maintain that the Bayankhongor ophiolitic mélange marks a suture zone, because ‘It seems unlikely that an ophiolite fragment 300 km long would remain intact and unmixed with the rest of the rocks in an accretionary wedge’ (Buchan et al., 2001, p. 459). This is a surprising statement in view of (1) the extremely disrupted and mélanged aspect of the Bayankhongor occurrence (Buchan et al., 2001), (2) the presence of similar areas in zones of accretion, such as the 500 km-long Kings-Kaweah ophiolite belt in the western foothills of the Sierra Nevada (Saleeby, 1977, 1981; Saleeby et al., 1978), the entire Coastal complex in the Coast Ranges in California that is coherent to the north of San Francisco for 200 km and incoherent but continuous (like the Bayankhongor suite) for 300 km south of it (e.g., Evarts, 1977; Contenius et al., 2000), and coherent slabs of considerable size within accretionary complexes in Guatemala (Aubouin et al., 1984) and Alaska (Plafker et al., 1994), (3) the location of the Bayankhongor ophiolite within a steep zone of metapelites and metapsammites, which, by the admission of Buchan et al. (2001) mostly show an incoherent mélange or broken formation style of deformation. There is no evidence northeast of the Bayakhongor ophiolitic belt of a continental piece onto which it could have been ‘obducted’. The alleged ‘passive continental margin’ of Badarch et al. (2002, p. 94) is described by them to include ‘lenses of limestone, sandstone, chert, tuff, minor felsic volcanic material and vesicular basalt’ which have been ‘intensively deformed and telescoped into a tectonic mélange’. The entire style of deformation is one of close to isoclinal folding with much strike-slip faulting with greenschist metamorphism typical of accretion complexes. To the northeast of the Bayankhongor ophiolite, the strongly deformed DevonianCarboniferous turbidites, erroneously interpreted as passive continental margin deposits by Buchan et al. (2001) and Badarch et al. (2002), continue for more than 1500 km into Russian territory. The turbidites contain slivers/beds of mafic and intermediate volcanic rocks and red pelagic cherts as recognised by Buchan et al. (2001) in Mongolia (Marinov et al., 1973; Zonenshain, 1972). Fault-bounded blocks and sheets of basalts, cherts and sedimentary rocks metamorphosed in greenschist facies appear among them but mainly along the margins of the Devonian-Carboniferous turbidite units (see Figs. 7 and 17). The age of these rocks has been interpreted as Vendian-Cambrian or as early Paleozoic, although evidence for these ascriptions is sparse. Majority of researchers agree that the DevonianCarboniferous and Vendian-Cambrian or Lower Paleozoic rocks of the Khangai-Khantey belong to subduction-accretion complexes (Sengör ¸ et al., 1993; Gusev and Khain, 1995; Sengör ¸ and Natal’in, 1996a, 1996b; Zorin, 1999; Parfenov et al., 1999). Finds of small lenses of ophirags support their conclusions (e.g., 23 in Fig. 17). In Mongolia, the age of the ophirgs/ophiolites is poorly constrained as pre-Silurian (Parfenov et al., 1999; Tomurtogoo, 1997). Pillow lavas and sheeted dykes reveal MORB-type features.

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The negative ε Nd signatures from the Palaeozoic and Mesozoic granites intruding Badarch et al.’s (2002) ‘passive continental margin’ most likely tap the highly deformed and thickened turbidite wedge of the Khangai-Khantey subduction-accretion complex that had been fed by the Precambrian Tuva-Mongol continental fragment (Zonenshain, 1973; Sengör ¸ and Natal’in, 1996a) much like the negative ε Nd signatures known from the Lachlan subduction-accretion complexes in the Tasman orogenic belt in Australia. The structural sections in the paper by Buchan et al. (2001) clearly indicate that the Bayakhongor ophiolite is an intra-accretionary complex ophiolite slab, much disrupted during incorporation. We are not even sure that it was indeed originally a single slab. In the Russian part of the Khangai-Khantey unit, the subduction-accretion complex includes Triassic turbidites. Fault-bounded blocks and sheets of strongly deformed metabasalts, cherts, and turbidites metamorphosed in greenschist facies as well as gabbro and ultramafic rocks are mapped as the Riphean Kulinda Suite (24 in Fig. 17; Suite as used in the Russian Stratigraphic nomenclature is exactly equivalent to the formation as defined in the International Stratigraphic Guide: Salvador, 1994, p. 42; de Wever and Popova, 1997, p. 394). Its age is not constrained by fossil findings or isotopic dating but inferred on the basis of long-distance correlation. Basalts of the Kulinda Suite reveal NMORB and T-MORB signatures indicating presence of fragments of normal oceanic crust and off-spreading magmatic constructions (Gusev and Peskov, 1993, 1996). Similar types of ophiolitic rocks have been established in the Molodovsk (25 in Fig. 17) and Gorbits regions (26 in Fig. 17). Younger, Devonian?-Triassic? ophiolites of the Ust-Tura region (27 in Fig. 17) represent a back-arc basement (Gusev and Peskov, 1996). This short overview shows that along the whole length of the Khangai-Khantey unit there is a unity of rock types and type of ophiolites/ophirags that mainly occur as slivers in mélanges and belong to N-MORB and OIB types. Only along the margins of the unit arc-related settings are known. Unit 43.3 (South Mongolian). Small tectonic lenses of ophirags are abundant in the South-Mongolian unit. They are commonly described as blocks in mélanges or imbricate thrust sheets (Marinov et al., 1973; Ruzhentsev et al., 1985, 1987; Ruzhentsev and Pospelov, 1992). Ophirags in the South Mongolian unit are associated both with tholeiitic and island arc volcanic rocks, shallow-marine and pelagic sediments. We are not aware of any recent detailed studies of these ophirags. They may be of both oceanic and island arc origin as it is indicated by recent detailed studies of sedimentary rocks (e.g., Lamb and Badarch, 1997). Unit 44 (South Gobi). New petrologic and isotopic studies are available for ophiolites of this unit. Petrographic, geochemical (major elements) and REE studies of the Hegeshan ophirags (28 in Fig. 17) of the Nei Mongol region, China, have shown a similarity with rocks of mid-ocean ridges (Nozaka and Liu, 2002). These ophirags represent basement fragments of the floor of the former Solonker ocean that closed in the Permian (Fig. 15; Zhang et al., 1984, Wang and Liu, 1986, 1991; Sengör ¸ and Natal’in, 1996a; Chen et al., 2000). However, metamorphic minerals of amphibolites yield unexpectedly young K-Ar ages of 110 and 130 Ma (early Cretaceous: Nozaka and Liu, 2002). The ophirags formed much earlier as it is shown by the Sm-Nd whole rock age of 403 ± 27 Ma

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(early Devonian: Chen et al., 2000). At the same time Robinson et al. (1999) report inhomogeneity of mineral composition and the geochemistry of the Hegeshan ophirags. Major element and trace element studies allowed them to distinguish rocks that originated in arc and backarc settings and in zones of within plate basaltic magmatism. The latter can be fragments of seamounts. We must note, however, that the Hegeshan ophirags occur in a tectonic mélange and tectonic settings, inferred solely on geochemical grounds from small fragments, need the structural relationships to be established to be entirely credible (e.g., Hsü et al., 1991). 4.3. Ophiolites and Ophirags in Collisional Suture Zones Unit 13 (Boshchekul-Tarbagatay). We have already pointed out that the BoshchekulTarbagatay arc system was built on an oceanic foundation. This arc system, that had originated as a single arc in the Cambrian (Sengör ¸ and Natal’in, 1996a), represents in fact a double arc, namely the Boshchekul-Tarbagatay sensu stricto (unit 13.1) and the BayanaulAkbastau (unit 13.2). These two arcs are now separated by the Maikain-Balkybek ophiolitic suture zone that forms a 700 km-long (i.e., almost exactly as long as the present-day Okinawa back-arc or the Mariana inter-arc/back-arc basin) and very narrow, linear/arcuate steep belt of ophiolites (Yakubchuk and Degtyarev, 1991). Sengör ¸ and Natal’in (1996a) had interpreted this suture to represent a now vanished early Ordovician intra-arc/backarc basin, the ophiolites of which had been clearly generated in a supra-subduction zone environment. The marginal basin closed by the early to medial Llandoverian. Unit 4.2 (Jalair-Nayman). A similar long and narrow, linear/arcuate pre-Ordovician ophiolite belt separates the Jalair-Nayman unit (unit No. 4.2) into two moieties. The preOrdovician ophiolites contain hornblende in the cumulate gabbro section and that is why Sengör ¸ and Natal’in (1996a) had interpreted them as of supra-subduction zone origin. Cambro-Ordovician turbidites cover the ophiolite stratigraphically. The basin closed during the Ordovician-(?)Silurian interval.

5. CONCLUSIONS The bold tone of many of the interpretations we present of the original tectonic settings of the ophiolites and the ophirags in the Altaid edifice in this paper must have struck the reader. In reality, our mood in formulating our interpretations was one of much diffidence. The reason for this diffidence was lack of precise and accurate observations in an orogenic system that is immense and occupies a terrain not entirely conducive to easy work. However, even where observations are plenty and multifarious, our diffidence was not entirely alleviated. The reason for this is the extremely disrupted state of the entire orogenic edifice. The scale of disruption ranges from literally millimetric scale in shear zones to tens or even hundreds of kilometres where large tectonic units are disrupted between the large cratons of Russia and Angara. Our bold presentation aims at making our choices clear with a view to encouraging field checks.

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What makes the tectonic interpretation of the oceanic basement remnants difficult is that most of them occur as ophirags and torn from their oiriginal places of incorporation into the continental crust. Processes of incorporation of oceanic basement rocks into the continental crust are complicated enough. Therefore, when intending to reconstruct the nature of ophiolitic rock assemblages in the Altaids (and in Turkic-type orogens: Sengör ¸ and Natal’in, 1996b), their original structure and original disposition with respect to neighbouring units, one must account for not only the complicated processes of incorporation into continents, but also for further deformation, tearing of the assemblages in pieces and their following distribution in areas with length scales in excess of 1000 km. All we see in the Altaids is that there are no privileged linear/arcuate zones along which ophiolites crowd as in familiar Alpine- and Himalayan-type collisional orogenic belts. Neither do they commonly define narrow, but long belts of accretionary complexes as in the Cordilleran belts. They appear to have formed in or adjacent to such accretionary complexes only after the pre-assembly form of the Altaid collage is reconstructed (see Figs. 10–16 herein). A glance at Fig. 17 shows that they are now literally everywhere within the orogenic collage! They are literally everwhere in form of small bodies compared with the intact, large ophiolite nappes of Oman, Turkey, Newfoundland or Papua New Guinea. In that respect they greatly resemble the Archaean ultramafics and mafic rocks. Moreover, many of the Altaid ophiolites have been interpreted as ensimatic island arc remnants or products of pre-arc spreading events. That is why they have thicker and more varied mafic, in places even intermediate and felsic, volcanic components. Such components range from mafic komatiites and boninites to even rhyolites, with island arc tholeiite and andesite dominance. This variety is also one we encounter in the Archaean greenstone belts (for a few recent cases relevant for rock-types mentioned in this paper, see Polat and Kerrich, 2001; Polat et al., 2002). When we add the reported slivers of oceanic plateaux and guyots in the Altaid ophirags, the resemblance to the Archaean terrains with their thick komatiitic lava piles becomes essentially complete. Perhaps the most valuable lesson a Precambrian geologist would learn from the Altaids is how careful she or he must be in judging the present setting and tectonic history of any piece of oceanic basement. In the Altaids, oceanic basement fragments became incorporated into continental structure by a truly bewildering diversity of mechanisms ranging from the familiar, cartoon-style, head on subduction-related trapping and offscraping effects through multifarious oblique-subduction-related strike-slip events and their associated complexities including arc-slicing and arc-shaving faults, releasing and restraining bends along them, numerous kinds of secondary structures that form along the paths of emerging shear zones such as Riedel and anti-Riedel shears, P- and X-shears, rotations around vertical axes of both blocks and their bounding structures, to a variety of collisional and post-collisional processes that both help incorporate ocean floor basement remnants into continents and, once incorporated, further deform, dismember and distribute them. Alternating head-on and oblique subduction events would repeatedly superpose their effects onto the incorporated and preserved basement fragments until they become almost totally bereft of any indications as to their original provenance.

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Hamilton (1992 and 1998) has complained that mélange belts were lacking in the Archaean. We think that this complaint is not justified. Sengör ¸ has been shown such jumbled rock associations resembling Phanerozoic mélange belts both in the Superior Province by Ali Polat (see Polat and Kerrich, 1999) and in the Yilgarn craton by Nick Archibald (see Sengör ¸ and Natal’in, 1996b). However, the extreme disruption seen in the greenstone belts that interleaves and intercalates ultramafics, mafic volcanic and plutonic rocks and sedimentary rocks commonly creates geometries that greatly resemble the mélange belts of the Phanerozoic. The problem in recognising Archaean mélanges is the associated rock types (ultramafic-mafic igneous rocks and dominantly pelites and cherts with little to no neritic limestone knockers), extreme deformation usually leading to complete transposition of original lithological contacts and the commonly accompanying, multi-phase amphibolite grade metamorphism that make the recognition of original rock types and geometries extremely difficult. In Isua, for instance, John Myers showed Sengör, ¸ in a deformed pillow lava, individual pillows with thicknesses of only a few tens of cm and lengths exceeding a few tens of metres! Under such conditions only a most meticulous, large-scale mapping by a geologist experienced in mélange tectonics can possibly decipher a mélange. But the resemblance does not stop at the rock association level. The structures seen in the Altaids and in the Archaean greenstone belts are very similar pointing to the operation of similar processes and in similar sequence (see Sengör ¸ and Natal’in, 1996b and the many contributions in de Wit and Ashwal, 1997). We are unable to agree with Hamilton’s (1992, 1998) defense of interpretations of simple deformation of greenstone belts. Sengör ¸ has first-hand experience in some of the Canadian and Yilgarn greenstone belts and in the Isua region of Greenland, which he was able to visit under the guidance of local experts of structural geology who had mapped them. In all of them the tectonic deformation is intense and multiphase and not confined to the surroundings of rising igneous diapirs as maintained by Hamilton. All of them very greatly resemble the tectonic style we are familiar with from the Altaids. In the Yilgarn, even the occurrence and sequence of strike-slip faulting as established by Nick Archibald (see Sengör ¸ and Natal’in, 1996b) are very reminiscent of the Altaids. Most Altaid oceanic basement fragments are ophirags. There is not in them a single giant ophiolite nappe of the kind we know from Oman or from Papua New Guinea or Turkey or Newfoundland. This is also the case in the Archaean. The entire intact Archaean terrain preserved today on the face of the earth occupies some 6% of the entire land surface, in other words an area about 8 million square kilometres (Condie and Sloan, 1998, Fig. 8.3; we think this is an overestimate!). The Altaid orogenic system, by contrast occupies some 8.5 million square kilometres. Therefore, the entire Altaid System covers almost exactly as much area as the entire area occupied by the intact Archaean crust! Since the whole of the Altaids do not contain a single giant ophiolite nappe, is it so surprising that the Archaean does not either? If only the Altaids were to be preserved 2500 million years hence, the geologists of that remote era (assuming that they would have the same means as we do today) would not have the slightest notion of the types of mountain-building, namely Alpine, Himalayan and the Andean, on which most of our theories of orogeny have so far been built.

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When one considers that the style of the Altaids is not unique to them in the Phanerozoic, but is represented by the Mesozoic-Cainozoic tectonics of the Nipponides (Sengör ¸ and Natal’in, 1996a, 1996b), of the North American Cordillera in the Mesozoic (Burchfiel et al., 1992), of the eastern part of a part of the Cimmerides in the Palaeozoic and the Mesozoic (Sengör, ¸ 1984), and of the South American Cordillera in the Palaeozoic-early Mesozoic and also in part of the eastern part of the Gondwanides in the late Palaeozoicearly Mesozoic (Sengör ¸ et al., 2001), we realise that the Archaean tectonic style is very much well and alive today. During the Phanerozoic, only in few places strike-slip faulting has gathered such an immense chunk of subduction-accretion complexes into a continental form as in the Altaids and that is why the great similarity of the Altaid style Turkic-type orogeny to Archaean terrains is not readily obvious. But a simple glance at Proterozoic tectonics in such places is in the Pan African system in northeast Africa and Arabia or in the Mazatzal System in North America convinces us that much of the continental crust has been built by Turkic-style orogens and that they have remained operative from the earliest Archaean to the present-day (Kusky and Polat, 1999).

ACKNOWLEDGEMENTS We thank Tim Kusky for inviting this contribution and waiting for it with endless and good-humoured patience. We are grateful to Brian F. Windley for helpful discussions on the geology of the Altaids. We also thank Bor-ming Jahn for his leadership in generating and administering international projects to test various tectonic models for the Altaids, from which we have derived great benefits. Rob van der Voo kindly informed us before publication of the results of his and his colleagues’ palaeomagnetic work in Central Asia. Evgeny V. Khain has discussed with us the tectonics of the Altay-Mongolian sector of the Altaids and generously shared his unpublished data and ideas on the Eastern Sayan and Central Mongolian ophiolites. V.E. Khain and Warren B. Hamilton have been, as always, helpful with discussions and excellent advice. Sengör ¸ is indebted to Ali Polat for leading him in the Schreiber-Hemlo belt in the Superior Province in Canada, to Ali Polat and John Myers for showing him part of the Isua region in western Greenland and to Nick Archibald for showing him a cross-section across the Yilgarn craton in Australia. He is also grateful to William R. Dickinson, for, among numerous other things, showing him the Californian Coast Ranges under his own incomparable guidance.

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Wang, Q., Liu, X., 1986. Paleoplate tectonics between Cathaysia and Angaraland in Inner Mongolia. Tectonics 5, 1073–1088. Wang, Q., Liu, X., 1991. Pre-Jurassic tectonic evolution between Cathaysia and Angaraland. In: Ishii, K., Liu, X., Ichikawa, K., Huang, B. (Eds.), Pre-Jurassic Geology of Inner Mongolia, China. Report of China-Japan Cooperative Research Group, 1987–1989. Osaka City University, Osaka, pp. 127–148. Wang, Z., Sun, S., Li, J.L., Hou, Q.L., Qin, K.H., Xiao, W.J., Hao, J., in press. Paleozoic tectonic evolution of the northern Xinjiang, China: geochemical and geochronological constrains from the ophiolites. Tectonics. Wilson, J.T., 1949. The origin of continents and Precambrian history. Transactions Royal Society of Canada 43 (Ser. 3, Sect. 4), 157–184. Woodcock, N.H., 1986. The role of strike-slip fault systems at plate boundaries. Philosophical Transactions of the Royal Society of London A 317, 13–29. Yakubchuk, A.S., Degtyarev, K.E., 1991. O kharaktere sochleneniya Chingizskogo i Boshchekulskogo napravlenii v kaledonidakh severo-vostoka Tsentralnogo Kazakhstana. Doklady Akademii Nauk SSSR 317, 957–962. Yilmaz, Y., Sengör, ¸ A.M.C., 1985. Paleo-Tethyan ophiolites in northern Turkey: Petrology and tectonic setting. Ofioliti 10, 485–504. Zhang, Z.M., Liou, J.G., Coleman, R.G., 1984. An outline of the plate tectonics of China. Bulletin of the Geological Society of America 95, 295–312. Zonenshain, L.P., 1972. Uchenie o Geosinklinalyakh i Ego Prilozhenie k Tsentralno-Aziatskomu Skladchatomu Poyasu. Nedra, Moscow, p. 240. Zonenshain, L.P., 1973. The evolution of central Asiatic geosynclines through sea-floor spreading. Tectonophysics 19, 213–232. Zonenshain, L.P., Kuzmin, M.I., 1978. Khan-Taishirskii ofiolitovyi kompleks zapadnoi Mongolii i problemy ofiolitov. Geotektonika (1), 19–42. Zorin, Y.A., 1999. Geodynamics of the western part of the Mongolia-Okhotsk collisional belt, TransBaikal region (Russia) and Mongolia. Tectonophysics 306, 33–56.

Precambrian Ophiolites and Related Rocks Edited by Timothy M. Kusky Developments in Precambrian Geology, Vol. 13 (K.C. Condie, Series Editor) © 2004 Elsevier B.V. All rights reserved.

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Chapter 22

EPILOGUE: WHAT IF ANYTHING HAVE WE LEARNED ABOUT PRECAMBRIAN OPHIOLITES AND EARLY EARTH PROCESSES? TIMOTHY M. KUSKY Department of Earth and Atmospheric Sciences, St. Louis University, St. Louis, MO 63103, USA

The chapters in this book have presented clear, even unequivocal evidence that Precambrian ophiolites are preserved in many Precambrian terranes. Proterozoic examples are abundant, especially in the Arabian Nubian Shield, where ophiolites have been recognized for many years. Archean examples are more controversial but a number of excellent examples of whole, dismembered, and metamorphosed ophiolites are described in this volume. In this brief epilogue, we assess what, if anything, we have learned about the early Earth from the identification of specific sequences as ophiolitic. In addition, we present a new list of criteria to help discriminate between ophiolitic and other sequences. The recognition that many of the allochthonous mafic/ultramafic complexes in Archean and Proterozoic greenstone belts are ophiolites provides researchers with a much longer record of oceanic processes than the record from Phanerozoic ophiolites alone. From this record we are able to deduce that the classical Penrose model (Anonymous, 1972) for the structure of ophiolitic lithosphere is too simplistic to explain the great variations found in ophiolites over this greater sample of time. The Penrose model for ophiolite stratigraphy is too restrictive to explain even present day sea floor and Paleozoic ophiolites, which all show much greater diversity (related to spreading rate, temperature, magma supply, etc.). Since modern environments and young ophiolites rarely conform to this strict definition, it makes little sense for Precambrian ophiolites to be held to this standard for recognition. It is more sensible to allow the diversity of modern ophiolites to be a guide to recognizing older ophiolites and their tectonic settings, and then to try to determine, through comparison, if there are any significant secular changes in ophiolitic structure and stratigraphy with time. With this caveat in mind, the chapters in this book have identified dozens of Precambrian ophiolites that contain an ophiolitic igneous stratigraphy. This basic recognition opens the way for a myriad of other studies on the chemistry, structure, thickness, rheology, biology, and other aspects of ancient oceanic crust and lithosphere that are only beginning to be appreciated. Once this recognition becomes more widespread and accepted, even greater insight to processes on the early Earth will be obtained. DOI: 10.1016/S0166-2635(04)13022-3

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Chapter 22: Epilogue

1. KOMATIITES, BONINITES, BIF’S, AND PODIFORM CHROMITES It has long been held that komatiites are abundant in Archean greenstone belts and that the Archean oceanic crust may have been dominantly komatiitic, reflecting early higher mantle temperatures. However, komatiites are much less common than many workers originally thought, and they do not necessarily mean much hotter mantle (see Parman and Grove, 2004). There has been a disproportionate number of studies of komatiites from Archean greenstone belts compared to other rock types, because petrologists have focused on the unusual aspects of these rocks, but they typically do not form more than a few percent of any greenstone terrain. However, they do appear to be more abundant in Archean terrains than younger ophiolites (e.g., Alvarado et al., 1997). Boninites are geochemically distinct mafic rocks that have been suggested to be absent from Archean terrains. As reported in several chapters in this volume (see Polat and Kerrich, 2004; Shchipansky et al., 2004; Stern et al., 2004), boninites have now been identified in several ophiolitic Proterozoic and Archean greenstone belts extending back in time to the 3.8 Ga Isua belt, suggesting that these ophiolites formed in environments similar to their modern counterparts. Boninites of Phanerozoic age occur in ophiolites or intraoceanic island arcs, such as the Izu-Bonin-Mariana arc system. These primary liquids are interpreted as second-stage high-temperature, low-pressure melting of a depleted refractory mantle wedge fertilized by fluids and/or melts, above a subduction zone. Precambrian boninitic lavas are likely products of the same conjunction of processes, suggesting that mantle melting processes above subducting slabs was broadly similar in the Archean to that of today. Podiform chromites form very distinctive deposits in many Phanerozoic ophiolites, and have been found in a few places on the modern sea-floor. Podiform chromites form small clusters of typically orbicular and nodular textured chromite in dunite pods, enclosed within mantle harzburgite tectonite. These chromite pods are distinctive, both physically and chemically, from layered chromite of layered ultramafic intrusive complex in continents (such as the Bushveld) and arcs (see Lago et al., 1982; Nicolas and Azri, 1991; Leblanc and Nicolas, 1992; Stowe, 1994; Butcher et al., 1999; Edwards et al., 2000). Until recently, podiform chromites were not known from any Archean greenstone belts, but their documentation in the Zunhua ophiolitic mélange and Dongwanzi ophiolite of North China (Kusky et al., 2004a; Huang et al., 2004) shows clearly not only that these rocks are ophiolitic, but that mantle melting processes in the Archean were similar to those of younger times. We suggest that since podiform chromites are only known from ophiolites, that they are as distinctive for recognizing a rock sequence as an ophiolite as the presence of the entire Penrose sequence. Banded Iron Formations (BIF’s) are a major component of many Archean greenstone terranes, and are described from several of the ophiolitic sequences in this volume. While the origin of BIF’s has been controversial, and there are several different origins (e.g., Fowler et al., 2002; Coward and Ries, 1995; Simonson, 1985), Hofmann and Kusky (2004) have shown how BIF’s in low-grade greenstone terranes may mark sites of regional structural detachment, with iron and sulfide mineralization focused along early shear zones.

2. Transitional Ophiolites

729

Workers in other greenstone terranes, particularly those that are more highly deformed and metamorphosed, should note the relationships at Belingwe, and re-assess whether or not BIF’s in other greenstones and Precambrian ophiolite terranes may mark the sites of major regional detachment and displacement. Several authors (e.g., Bickle et al., 1994; Hamilton, 2003) have noted that some Precambrian greenstone belts show evidence of contamination by continental type material, and have then suggested that this means that they cannot be fragments of oceanic crust and lithosphere. These authors have failed to note that many modern and Phanerozoic ophiolites also show such contamination (e.g., Moores, 2002), invalidating those arguments. Nonetheless, apparent contamination by continental crustal material presents interesting constraints on the origin of these ophiolites. For instance, apparent crustal contamination can mean lavas were derived from unusual mantle, such as an older forearc environment, where subduction-related processes may have depleted the mantle leading to unusual, apparently contaminated geochemical signatures (see Parman and Grove, 2004). Alternatively, some ophiolites may be truly contaminated, having formed near a stretched continental margin. Some ophiolites seem to preserve magmatism near these margins, and some even have subcontinental lithospheric mantle and/or crust preserved. We coin a new term for these ophiolites, and call them transitional ophiolites.

2. TRANSITIONAL OPHIOLITES Several of the ophiolites described in this volume appear to have formed within the transition from rifted continental margins to ocean spreading centers during early stages of ocean opening, then were structurally detached and/or deformed and incorporated into convergent margins during ocean closure. These ophiolites are distinctive from classical Penrose-style ophiolites and others formed in forearc and back arc environments. During early stages of ocean formation, continental crust and lherzolite of the subcontinental mantle is extended forming graben on the surface, and ductile mylonites at depth. Sedimentary basins may form in the graben, and as the extension continues magmatism sometimes affects the rifted margin, either forming volcanic rifted margins, or migrating to a spreading center forming a oceanic spreading center. New asthenospheric mantle upwells along the new ridge, and may intrude beneath the extended continental crust. In some cases, wedges of extended mid-to-lower continental crust overlying mylonitic lherzolitic subcontinental mantle become intruded by numerous dikes and magmas from this new asthenospheric mantle. In this case, magmas may pool both above and below the stretched continental crust, forming mafic/ultramafic cumulates in igneous contact with older continental crust (see Fig. 1 in the Introduction to this volume). Dikes from these magma chambers may then feed a crustal gabbroic magma chamber closer to the surface, which in turn may feed a dike complex and basaltic pillow/massive lava section. If preserved, this unusual sequence forms what we term a “transitional ophiolite”, grading down from subaquatic sediments, to pillow lavas, dikes, sheeted dikes, layered gabbro, dunite and pyroxenite cumulates, then remarkably into stretched, typically mylonitic granitic mylonites, underlain by lherzolite.

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Chapter 22: Epilogue

The lherzolite tectonic may be underlain by harzburgite tectonite or harzburgite. Recognition of this relationship represents a major advance in understanding some of the ophiolitic complexes described in this volume, and elsewhere. Examples of this type of transitional ophiolite are found in the Proterozoic Jourma complex, and in some of the Slave Province ophiolites (see papers by Peltonen and Kontinen, 2004, and Corcoran et al., 2004). Modern analogs for such transitional ophiolites are found around the Red Sea, including at Tihama Asir, Saudi Arabia, where a 5–10 Ma old transitional ophiolite has a dike complex overlying layered gabbro, which in turn overlies continental crust. Also, on Egypt’s Zabargad Island, oceanic mantle is exposed, and it is likely that the crustal structure near this region preserves transitional ophiolites as well. The main lesson here is that ophiolites may form in many tectonic settings, from extended continental crust, to mid ocean ridges, to forearcs, arcs, back arcs, to triple junctions along convergent margins.

3. PROTEROZOIC OPHIOLITES The formation of the Gondwanan supercontinent at the end of the Precambrian and the dawn of the Phanerozoic represents one of the most fundamental problems being studied in Earth Sciences today. It links many different fields, and there are currently numerous and rapid changes in our understanding of events related to the assembly of Gondwana. One of the most fundamental and most poorly understood aspects of the formation of Gondwana is the timing and geometry of closure of the oceanic basins which separated the continental fragments that amassed to form the Late Proterozoic supercontinent. Final collision between East and West Gondwana most likely occurred during closure of the Mozambique Ocean, forming the East African Orogen including the Arabian-Nubian Shield. Neoproterozoic ophiolite fragments have been recognized as a component part of many nappe complexes associated with sutures in the Arabian-Nubian Shield. The recognition of these ophiolite-decorated sutures played a major role in understanding the formation of the Arabian-Nubian Shield as an amalgam of different arc and microcontinental terranes that collided during the closure of the Mozambique Ocean (Stern, 1994; Kusky et al., 2003), but also contributed to many scientists’ acceptance that plate tectonics extended back in time to 890 Ma, the age of the oldest Arabian ophiolite (see Stern et al., 2004; Johnson et al., 2004). The chemistry of the Arabian Shield ophiolites include both tholeiitic and calc-alkaline varieties, with minor boninites, suggesting that they largely formed in a forearc environment, with extensive partial melting of the mantle. Many of the ArabianNubian shield ophiolites formed over a critical interval of Earth history that saw many changes in the Earth’s biota and climate, yet very few studies have yet been aimed at the sedimentary sequences that overlie these ophiolites, potentially preserving a treasure drove of information about the Neoproterozoic Earth.

4. Archean Ophiolites

731

4. ARCHEAN OPHIOLITES Over the course of several decades, a number of possible partial and dismembered ophiolite sequences have been described from a number of Archean greenstone belts of different ages and locations (e.g., de Wit et al., 1987; de Wit and Ashwal, 1997; Fripp and Jones, 1997; Harper, 1985; Kusky, 1989, 1990, 1991). However, few complete Phanerozoic-like ophiolite sequences have been recognized in Archean greenstone belts, leading some workers to the conclusion that no Archean ophiolites or oceanic crustal fragments are preserved (Bickle et al., 1994; Hamilton, 1998, 2003). These ideas were challenged by the recognition of a complete but partially dismembered Archean ophiolite sequence from the North China Craton (Kusky et al., 2004a), that was later found to be associated with mantle tectonites in mélange beneath the ophiolite (Li et al., 2002). This discovery has important implications for understanding other Archean greenstone belts, many of which contain only part of the typical ophiolite assemblage. With a complete assemblage present in at least one locality, it is more likely that the other reported partial sequences are truly parts of ophiolites, and not representative of some other tectonic setting that was unique in the Precambrian. Some workers have even suggested that the mechanisms of planetary heat loss changed so much with time so that what resembles an ophiolite from the Archean record is actually equivalent to a continental rift in the younger rock record. The Penrose definition of ophiolites (Anonymous, 1972; cf. Brongniart, 1813, 1821) includes “dismembered”, “partial”, and “metamorphosed” varieties, with rock types the same as those that typify Archean greenstone belts. Ophiolite-like relationships have been described for many years from Archean greenstone belts (Hess, 1955), yet many of the examples of partial ophiolites in Archean terrains were questioned, because no complete sequences were found anywhere. If such sequences were found in younger, Phanerozoic mountain belts, the ophiolitic origin for the rock sequence would not likely be questioned. In this book, many such ophiolitic sequences are described, and the authors take the approach of using the same criteria to identify ophiolites in very old rocks as they do in younger orogenic belts. The application of different paradigms to the Archean and Phanerozoic is no longer necessary, although detailed studies are beginning to reveal some differences in the style of older and younger sea floor spreading. Better quantification of these differences and similarities will help constrain geochemical, geodynamic, and thermal modeling of what effects the Archean mantle thermal and melting regime had on the structure of oceanic lithosphere produced in those times. Archean oceanic crust was possibly thicker than Proterozoic and Phanerozoic counterparts, resulting in accretion predominantly of the upper basaltic section of oceanic crust. However, structural repetition and complexities in greenstone belts makes it very difficult to assess original thicknesses, as shown by papers in this book, and in Kusky and Vearncombe (1997). The crustal thickness of Archean oceanic crust may have resembled modern oceanic plateaux (e.g., Kusky and Kidd, 1992; Kusky and Winsky, 1995; Kusky, 1998), but if average oceanic crust was this thick, then they would not be topographically high standing plateaus, and the term plateau would be meaningless. If this were the case, the rheological stratification of the oceanic lithosphere would have been different (Hoffman

732

Chapter 22: Epilogue

and Ranalli, 1988), and complete Phanerozoic-like MORB-type ophiolite sequences would have been very unlikely to be accreted or obducted during Archean orogenies. In contrast, only the upper, pillow lava-dominated sections would likely be accreted. Future research should be directed at considering what the consequences of changes in the style and/or composition of accreted oceanic material has on the structure and composition of the continental crust. For instance, the observation that Archean greenstone belts have an abundance of accreted ophiolitic fragments compared to Phanerozoic orogens suggests that thick, relatively buoyant, young Archean oceanic lithosphere may have had a rheological structure favoring delamination of the uppermost parts during subduction and collisional events (see Hoffman and Ranalli, 1988; Kusky and Polat, 1999). Subcrustal oceanic lithosphere slabs may have been underplated beneath cratons, forming mantle roots (Kusky, 1993). Descriptions of the various Precambrian ophiolites in this volume has shown that many are contained within thrust complexes that include elements formed at different stages of ocean opening and closing, with a strong bias toward convergent margin environments such as suprasubduction zone or arc-related spreading centers, accretionary wedge material, triple-junction related magmas and arc magmas that have migrated through these orogenic collages. Many of these ophiolites are parts of orogenic complexes that have experienced complex tectonic histories, similar to those of material accreted to younger accretionary orogens (Kusky et al., 2004b; Sengör and Natal’in, 2004). As in younger ophiolites, at present we observe a huge variation in inferred crustal thickness of ophiolites, and in the units that are preserved. With present limited data, and the amount of structural complication, it is not yet possible to assess whether or not there has been a demonstrable secular change in the thickness of oceanic crust (e.g., Moores, 2002). However, obtaining better constraints on the thickness and petrological relationships in Precambrian ophiolites remains a high priority for the Earth Sciences, since these are sensitive indicators to the nature of how the Earth lost the extra heat produced during the Precambrian. Since the current range of known ophiolitic thicknesses and internal stratigraphic relationships from Precambrian ophiolites are within the range of those from the Phanerozoic to present regime, we favor the idea that Precambrian ophiolites were not drastically thicker than those of younger times, and that much of the heat from the Precambrian Earth was lost though a greater total ridge length, and faster spreading rates, rather than production of dramatically thicker melt columns. Such relationships would produce a Precambrian Earth dominated by smaller oceanic plates, more triple junction interactions, and a younger average age of subducting lithosphere. Thicker sedimentary piles of graywacke turbidities on subducting plates would lead to fewer mélanges being formed (see Kusky et al., 2004b), and to more low-angle subduction with many ridge subduction events and belts of near-trench magmas intruding the accretionary margins and ophiolites, forming the TTG suite. The smaller oceanic plate size does not necessarily mean that continents were also smaller. We know, for instance, that Precambrian quartzites such as the Mt. Narryer required long rivers on large continents to form the extensive mature sands, and that some Archean strike slip faults (Sleep, 1992; Kusky and Vearncombe, 1997) and Archean passive margin sequences (Kusky and Hudleston, 1999) both had lengths exceeding 1,000 km.

4. Archean Ophiolites

733

Table 1. Criteria for recognition of a rock sequence as an ophiolite Indicator

Importance

Full Penrose sequence diagnostic in order

Status Status in Phanerozoic in Dongwanzi Ophiolites rare, about 10% suggested, needs documentation and verification

Conclusion

not conclusive

Podiform chromites w/nodular textures

diagnostic

about 15%

present

diagnostic

Full sequence dismembered

convincing

about 30–50%

dismembered units present

convincing

3 or 4 of 7 main units present

typical for accepting Phan. Ophiolite

about 80%

6 of 7 units known; dikes still uncertain (age)

convincing

Sheeted dikes

distinctive, nearly diagnostic

about 20–30%

suggested, age needs verification

not conclusive

Mantle tectonites

distinctive

about 20–30%

present

distinctive

Cumulates

present, not distinctive

about 70%

present

supportive

Layered gabbro

typical

about 70%

present

supportive

Pillow lavas

typical, not distinctive

about 85%

present

supportive

Chert, deep water seds

typical

about 85%

present

supportive

Co-magmatic dikes and gabbro

necessary, rare to about 15% observe

present

distinctive

High-T silicate defm. rare, but as inclus. in melt pods distinctive

about 10%

present

distinctive

Basal thrust fault

necessary (except in rare cases), not diag.

about 60%

present

supportive

Dynamothermal aureole

distinctive, almost diagnostic

about 15%

not determined

inconclusive

Sea floor metamor

distinctive

all

present

supportive

Hydrothermal vents black smoker type

distinctive

rare

present

strongly supports (continued on next page)

734

Chapter 22: Epilogue

Table 1. (Continued) Indicator

Importance

Status Status Conclusion in Phanerozoic in Dongwanzi Ophiolites Ophiolites are defined on the basis of field relationships and the overall rock sequence. Many workers have added chemical criteria to the ways to recognize and distinguish between different types of ophiolites. Some of the more common traits are: MORB chem.

distinctive

about 65%

present

distinctive

CA chemistry

common

about 30%

present in some units

inconclusive

Flat REE

distinctive

about 65%

present

distinctive

Boninites

distinctive of SSZ

rare but increasingly recognized

not known

inconclusive

Terranes in which Precambrian ophiolites are found resemble younger accretionary orogens, such as Alaska, the Altaids, or the Philippines, where each of these has long history and many magmatic events (Kusky et al., 2004b; Sengör and Natal’in, 2004; Encarnacion, 2004). Comparative studies between accretionary orogens and Precambrian cratons, orogens, and ophiolites are likely to continue to yield useful insights about how the early Earth operated. 5. IS IT AN OPHIOLITE? Several authors have presented various schemes to purportedly discriminate between ophiolitic and other sequences (e.g., Pearce, 1987; Wood et al., 1979), although most of these are either arbitrary, or based on models of what the authors believe Precambrian ophiolites should have looked like (e.g., Bickle et al., 1994). Here, we present a shamelessly uniformitarian list of criteria that can be used to determine the likelihood of whether or not a partial, dismembered, or complete sequence is ophiolitic, though comparison with betterunderstood Phanerozoic sequences. For comparison, the Dongwanzi ophiolite is compared to Phanerozoic ophiolites, and it stands up well to such comparison, and would clearly be called an ophiolite if it were preserved in a Phanerozoic orogen. Table 1 can be used for other questionable sequences, by replacing the column for the Dongwanzi ophiolite with the sequence in question. REFERENCES Alvarado, G.E., Denyer, P., Sinton, C.W., 1997. The 89 Ma Tortugal komatiitic suite, Costa Rica: implications for a common geological origin of the Caribbean and eastern Pacific region from a mantle plume. Geology 25, 439–442.

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Johnson, P.R., Kattan, F.H., Al-Saleh, A.M., 2004. Neoproterozoic ophiolites in the Arabian Shield: Field relations and structure. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 129–162. Kusky, T.M., 1989. Accretion of the Archean Slave Province. Geology 17, 63–67. Kusky, T.M., 1990. Evidence for Archean ocean opening and closing in the southern Slave Province. Tectonics 9, 1533–1563. Kusky, T.M., 1991. Structural development of an Archean orogen, western Point Lake, Northwest Territories. Tectonics 10, 820–841. Kusky, T.M., 1993. Collapse of Archean orogens and the generation of late- to post-kinematic granitoids. Geology 21, 925–928. Kusky, T.M., 1998. Tectonic setting and terrane accretion of the Archean Zimbabwe craton. Geology 26, 163–166. Kusky, T.M., Hudleston, P.J., 1999. Growth and Demise of an Archean carbonate platform, Steep Rock Lake, Ontario, Canada. Canadian Journal of Earth Sciences 36, 1–20. Kusky, T.M., Kidd, W.S.F., 1992. Remnants of an Archean oceanic plateau, Belingwe greenstone belt, Zimbabwe. Geology 20, 43–46. Kusky, T.M., Polat, A., 1999. Growth of Granite-Greenstone Terranes at Convergent Margins and Stabilization of Archean Cratons. In: Marshak, S., van der Pluijm, B. (Eds.), Special Issue on Tectonics of Continental Interiors. Tectonophysics 305, 43–73. Kusky, T.M., Vearncombe, J., 1997. Structure of Archean Greenstone Belts. In: de Wit, M.J., Ashwal, L.D. (Eds.), Tectonic Evolution of Greenstone Belts, Oxford Monograph on Geology and Geophysics, pp. 95–128. Kusky, T.M., Winsky, P.A., 1995. Structural relationships along a greenstone/shallow water shelf contact, Belingwe greenstone belt, Zimbabwe. Tectonics 14, 448–471. Kusky, T.M., Abdelsalam, M., Tucker, R., Stern, R., 2003. Evolution of the East African and Related Orogens, and the Assembly of Gondwana. Special Issue of Precambrian Research 123, 81–344. Kusky, T.M., Li, J.H., Glass, A., Huang, X.N., 2004a. Origin and emplacement of Archean ophiolites of the Central Orogenic belt, North China craton. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 223–274. Kusky, T.M., Ganley, R., Lytwyn, J., Polat, A., 2004b. The Resurrection Peninsula ophiolite, mélange, and accreted flysch belts of southern Alaska as an analog for trench-forearc systems in Precambrian orogens. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 627–674. Lago, B.L, Rabinowicz, M., Nicolas, A., 1982. Podiform chromitite ore bodies: a genetic model. J. Petrol. 23, 103–125. Leblanc, M., Nicolas, A., 1992. Ophiolitic chromitites. International Geological Reviews 34, 653– 686. Li, J.H., Kusky, T.M., Huang, X., 2002. Neoarchean podiform chromitites and harzburgite tectonite in ophiolitic melange, North China Craton: Remnants of Archean oceanic mantle. GSA Today 12 (7), 4–11. Moores, E.M., 2002. Pre-1 Ga (Pre-Rodinian) ophiolites: Their tectonic and environmental implications. Geological Society of America Bulletin 114, 80–95. Nicolas, A., Azri, H.A., 1991. Chromite-rich and chromite-poor ophiolites: the Oman case. In: Peters, Tj., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer Academic, Boston, pp. 261–274.

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Parman, S.W., Grove, T.L., 2004. Petrology and geochemistry of Barberton komatiites and basaltic komatiites: evidence of Archean fore-arc magmatism. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 539–565. Pearce, J., 1987. An expert system for the tectonic characterization of ancient volcanic rocks. J. Volcanol. Geotherm. Res. 32, 51–65. Peltonen, P., Kontinen, A., 2004. The Jormua ophiolite: A mafic-ultramafic complex from an ancient ocean-continent transition zone. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 35–71. Polat, A., Kerrich, R., 2004. Precambrian arc associations: Boninites, adakites, magnesian andesites, and Nb-enriched basalts. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 567–597. Sengör, A.M.C., Natal’in, B.A., 2004. Phanerozoic analogues of Archean oceanic basement fragments: Altaid ophiolites and ophiorags. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 675– 726. Shchipansky, A.A., Samsonov, A.V., Bibikova, E.V., Babarina, I.I., Konilov, A.N., Krylov, K.A., Slabunov, A.I., Bogina, M.M., 2004. 2.8 Ga Boninite-hosting partial suprasubduction zone ophiolite sequences from the North Karelian greenstone belt, NE Baltic Shield, Russia. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 425–486. Simonson, B.M., 1985. Sedimentologic constraints on the origins of Precambrian iron-formations. Geological Society of America Bulletin 96, 244–252. Sleep, N.H., 1992. Archean plate tectonics: what can be learned from continental geology?. Canadian Journal of Earth Sciences 29, 2066–2071. Stern, R.J., 1994. Arc Assembly and continental collision in the Neoproterozoic East African Orogen: Implications for the assembly of Gondwanaland. Annual Reviews of Earth and Planetary Sciences 22, 319–351. Stern, R.J., Johnson, P.R., Kröner, A., Yibas, B., 2004. Neoproterozoic ophiolites of the ArabianNubian Shield. In: Kusky, T.M. (Ed.), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology, vol. 13. Elsevier, Amsterdam, pp. 95–128. Stowe, C.W., 1994. Compositions and tectonic settings of chromite deposits through time. Economic Geology 89, 528–546. Wood, D.A., Joron, J.L., Treuil, M., 1979. A re-appraisal of the use of trace elements to classify between magma series erupted in different tectonic settings. Earth Planet. Sci. Lett. 45, 326–336.

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739

SUBJECT INDEX

A Abitibi, 10, 567, 584 Abitibi greenstone belt, 577, 580 Abitibi terrane, 23 Abitibi-Wawa subprovince, Superior Province, 663 accreted oceanic plateau, 592 accretion, 662, 732 accretionary complex, 714 accretionary lapilli, 520 accretionary orogen, 429, 623, 689, 732, 734 accretionary prism, 619, 628, 647, 656, 661 accretionary wedge, 14, 207, 628, 703, 707 accretionary wedge material, 732 adakite, 23, 448, 477, 568, 569, 581, 592, 663 adakite-high magnesian andesite, 577 adakitic melt, 478 Agardagh Tes-Chem ophiolite, 17, 207–219, 706 Al Amar suture, 147 Al-depleted komatiite, 554 Al-depleted type of komatiitic rocks, 413 Al-undepleted komatiite, 554 Al-undepleted type komatiite, 413 Alaska, 10, 25, 628, 734 Aldan Shield, 21, 405, 406 Aleutian Islands convergent margin, 584 Aleutians, 678 Alfred Wegener, 682 alkaline, 67 Allaqi-Sol Hamid suture, 132 Alpine-type peridotite, 200, 608 Altaid evolution, 699 Altaids, 26, 208, 676, 688–716, 734 Altay Mountains, 688 Altay-Mongol domain, 692 Altay-Sayan sector, 698 amphibole gabbro, 111, 183, 184 amphibolite, 146, 572 amphibolitization, 150 Andes, 584 andesite, 113, 155, 432, 620 andesite-dacite-rhyolite suite, 477

Angara craton, 694 anhydrite, 353 anhydrous melting, 541 anorthosite, 240 antinodular chromite, 249, 250 antinodular texture, 250, 325 Anton complex, 366 Anton terrane, 397 apatite, 185 Arabian-Nubian Shield, 11, 15, 16, 95–123, 129–159, 163–202, 582, 727, 730 arc pluton, 620 arc tholeiite, 577 arc-like chemistry, 647 arc-related spreading center, 732 arc-shaving fault, 714 arc-slicing, 714 arc-trench assemblage, 448 Archaean granite-greenstone terrain, 699 Archaean greenstone belt, 714 Archaean mélange, 715 Archaean tectonic style, 716 Archaean ultramafic and mafic rocks, 714 Archean, 264 Archean greenstone belt, 599, 731 Archean mantle, 23 Archean mantle temperature, 560 Archean plate tectonics, 481, 560, 568, 591, 610, 662, 732 Archean spreading rate, 610 Arizona, 75 assimilation, 215, 647 assimilation of sediments, 666 Atacama fault, 694 Aufbruchszone, 679 augite, 545, 548 Austroalpine nappe system, 679 Azores fracture zone, 678 B back-arc basin, 202, 619, 677 back-arc magma, 217 back-arc tholeiite, 551

740

backstop, 707 backstop to accretionary wedge, 703 Baltic Shield, 22, 425, 426, 567, 579 banded iron formation, 80, 117, 224, 236, 506, 571, 572, 663, 728 Banting Group, 368 Barberton greenstone belt, 8, 23, 539–561, 606 Barrovian-type metamorphism, 442 basalt, 113, 141, 148, 155, 210, 581, 638 basaltic andesite, 638 basaltic komatiite, 541, 544, 551 basaltic sill, 80 basement complex, 84 basement contacts, 363 basement of ensimatic arc, 703 Batomga, 406 Beaulieu River volcanic belt, 379–397 Belingwe belt, 22 Belomorian mobile belt, 428 Beniah Lake fault zone, 364 Besshi-type deposit, 339, 358 Besshi-type sulfide deposit, 355 Besshi-type VMS deposit, 356 Betts Cove ophiolite, 583 Bi’r Tuluhah ophiolite, 143–147 Bi’r Umq ophiolite, 141–143 Bi’r Umq suture, 137 Bi’r Umq-Nakasib suture, 99, 158 BIF, 224, 225, 343, 345, 608, 663 birbirite, 118 black shale, 39 black smoker, 7 black smoker chimney, 20, 339, 349, 357 black smoker model, 357 Bogoin-Boali greenstone belt, 582 boninite, 16, 22, 23, 99, 114, 420, 425, 426, 438, 541, 551, 554, 567–592, 606, 619, 663, 704, 706, 728, 730 boninite series, 429, 434, 442, 457 Boruss mélange belt, 707 Boshchekul-Tarbagatay, 704, 713 brine-pool model, 357 Brongniart, 687 Bulawayan Group, 490, 527 C Calaguas ophiolite, 619 calc-alkaline, 113, 298, 644, 647, 730 calc-alkaline arc association, 580 calc-alkaline association, 577 calc-alkaline basalt, 645

Subject Index

calc-alkaline ophiolite, 647 calc-alkaline series, 195 Cameron and Beaulieu River belts, 366 Cameron and Beaulieu River greenstone belts, 9 Cameron and Beaulieu River volcanic belts, 20 Cameron River, 605 Cameron River belt, 379–397 carbonate alteration, 122, 152 carbonate-sulfide chimney, 351 carbonated, 141 carbonatite, 64 carbonatization, 44, 118, 171, 173, 179 carbonatized serpentinite, 172 Casiguran ophiolite, 619 Central African Republic, 582 Central America, 584 Central Asian mobile belt, 17, 207 Central Orogenic belt, 226, 228, 229, 265, 275, 283, 322, 340, 600 chemical stratigraphy, 656 chert, 4, 80, 116, 141, 142, 210, 225, 571, 572, 608, 637 chert horizon, 488 chert tectonite, 501, 506, 510 chert-BIF, 349 Cheshire Formation, 498, 504–506, 525 Chile Rise, 657 Chilimanzi suite, 517 China, 226 Chingezi gneiss complex, 490, 493 Chingezi suite, 493 chrome number, 58, 106, 200, 216, 252, 257, 258, 412 chromite, 53, 122, 150, 173, 180, 417 chromite vein, 326 chromitite, 248 Chugach accretionary wedge, 661 Chugach terrane, 13 Chugach-Prince William terrane, 25, 627, 628 climate, 662 clinopyroxenite, 41, 146 coherent flysch terrain, 658 collisional suture zone, 703, 713 Coloumb wedge, 661 coloured mélange, 678 conglomerate, 571 contact relationships, 370 continental flood basalt, 520 continental growth, 8, 11, 543, 568, 591, 658, 663, 665, 686, 699 continental rifting, 36, 64

Subject Index

Contwoyto terrane, 397 convergent margin, 12, 664 convergent margin environment, 732 convergent plate boundary, 591 Cordilleran-type, 165 Cordilleran-type ophiolite, 201, 368 critical taper, 661 crustal contamination, 521–523, 528, 558, 561, 729 crustal growth, 14, 543, 569 crustal thickness, 8, 731 crustal thickness of ophiolites, 732 cumulate, 2, 146, 171, 178, 180, 212, 216, 293 cumulate layer, 42, 173 cumulate rocks, 134 cumulate ultramafics, 110 cumulate zone, 410 Cyprus-type sulfide body, 356 D D1 ophiolitic thrusting, 450 dacite, 83, 432 dacitic, 113, 620 dacitic dike, 78 dacitic dyke, 433 dacitic flow, 80 deep water sediments, 7 deep-sea environment, 681 deep-sea sediments, 4 deformation, 714 depleted mantle, 241 Destor-Porcupine-Manneville fault zone, 577 Dibut Bay ophiolite, 619 differentiated flow, 409 differentiation process, 610 dike complex, 394 diorite, 176, 211, 477 diorite-tonalite, 615 diorite-tonalite intrusion, 622 discrimination diagram, 734 dismembered, 7, 714, 731 dismembered ophiolite, 102 disseminated chromite, 53, 55, 252 Djezkazgan-Kirgiz, 708 Dongwanzi ophiolite, 18, 19, 223–267, 275, 276, 283–318, 600 ductile shear zone, 231 Duncan Lake Group, 368 dunite, 110, 134, 146, 172, 179, 210, 216, 240, 244, 248, 289, 293, 321, 412, 606, 608, 632 dunite-peridotite, 409

741

dunitic, 171, 605 dunitic tectonite, 606 dyke, 706 dynamic magmatism, 263 Dzhida, 710 E earliest Precambrian rocks, 684 earliest thrust faulting, 433 early life, 7, 339 early nappe-style deformational event, 450 early shear zone, 728 Earth’s early biosphere, 225 East African Orogen, 11, 96, 99, 164, 730 Eastern Altay, 704, 710 Eastern Block, 227, 228, 275, 283, 340 eclogite, 10, 265, 609, 664 eclogitic residue, 664 Eduard Suess, 679 Émile Argand, 681 emplacement, 265 enriched mid-ocean ridge basalts (EMORB), 47 ensimatic arc, 707 epidosite, 85, 339, 354 epidote-clinozoisite, 238 epidotized, 60 Ethiopia, 99 exhalite, 345 F felsic tuff, 544 Fennoscandian Shield, 38 filter pressing, 55 flexural loading, 519 flower structure, 158 flysch, 12, 26, 267, 630, 634, 637, 648, 657 fold-thrust belt, 229, 542 forearc, 560, 583, 703, 706, 729 forearc evolution, 657, 665 forearc extension, 479 forearc ophiolite, 120 forearc spreading, 706 foreland basin, 229, 267, 519, 525, 526 foreland fold-thrust belt, 228, 232 foreland-thrust belt, 20 Fortescue Group, 521 Fox Island shear zone, 627, 650, 652, 653 fractionation history, 455 fracture zone environment, 678 frontal accretion, 661 Frotet-Evans greenstone belt, 579

742

G gabbro, 2, 39, 47, 50, 77, 111, 134, 141, 146, 148, 152, 155, 173, 174, 183, 238, 240, 246, 294, 295, 444, 610, 635, 706, 707 gabbro sill, 409, 433 gabbrodiorite, 210, 211 gabbroic, 171, 210 gabbroic dike, 58 gabbroic plutonic sequence, 605 gabbroic rocks, 619 gabbroic unit, 443 gabbronorite, 183, 210, 211 garnet amphibolite, 704 geochemical discrimination, 647 geothermal gradient, 7, 24, 568, 583, 591, 703 giant ophiolite nappe, 678 glaciation, 118 gold, 122, 173, 634, 658, 663 Gondwana, 165 Gondwanaland, 96 Gondwanan supercontinent, 730 Gorringe Bank, 678 granite, 84 granodiorite, 477, 633, 664 granodiorite suite, 477 granodioritic magma, 224 granulite, 228, 231, 265 graywacke, 39 Great Dyke, 490, 517 Great Lakes tectonic zone, 580 Greenland, 570 greenstone belt, 731 growth faulting, 357 H Hackett River arc, 397 Halaban ophiolite, 147 Halaban suture, 159 Hamisana shear zone, 165 Han-Taishir, 706 Hans Stille, 682 harzburgite, 2, 39, 104–110, 133, 137, 146, 179, 210, 216, 224, 240, 241, 243, 246–248, 262, 289, 290, 321, 326, 606, 619, 632 harzburgite tectonite, 104, 110, 240, 247, 275, 285, 728 harzburgitic ultramafics, 99 hazelwoodite nickel, 182 heat flow, 426, 592 heat loss, 1, 6, 732 heat production, 6

Subject Index

Hess, 682 Hess’s Precambrian oceans, 686 high degree of mantle melting, 420 high H2 O contents, 550 high P-T mineral, 448 high-Mg andesite, 432 high-pressure granulite, 228, 264 high-pressure granulite belt, 229, 267 high-temperature deformation, 104, 137, 241, 327, 329, 334 high-temperature ductile fabric, 178 high-temperature fabric, 180 high-temperature foliation, 247 high-temperature isoclinal folding, 178 high-temperature metamorphism, 433 high-temperature plastic flow, 19, 321 high-temperature shear strain, 183 high-temperature shearing, 331 high-Ti-ophiolite, 195 history of ophiolite studies, 676 Hoogenoeg Formation, 544, 606 Hoogenoeg magma, 560 hornblende gabbro, 172, 210, 211, 216 hornblende-rich gabbroic, 211 hornblende-rich phase, 146 hornblendite, 41, 247, 290, 293 hornblendite-garnetite cumulate, 35 hot spot, 4 hot subduction, 588, 592 hydrated mantle wedge, 478 hydrothermal alteration, 84, 179 hydrothermal circulation, 339 hydrothermal system, 85 hydrothermal vent, 122, 349 hypabyssal sill, 606 I IAT-MORB, 201 Iceland, 4 imbricate fold-thrust wedge, 433 imbricate stack, 446 immature sediments, 615 Indus suture, 678 inherited zircon, 523 Inner Mongolia, 226 intra-oceanic arc, 577 Iringora ophiolite, 443 Iringora SSZ ophiolite, 451 Iringora Structure, 426, 435, 442 iron formation, 38 ironstone, 494, 499, 501 Ishkeolmes, 704

Subject Index

island arc, 677 island arc magma, 216 island-arc tholeiite, 645 isotropic gabbro, 2, 42, 174, 184 Isua, 567, 584 Isua greenstone belt, 570, 573 Isua terrane, 23 Itogon ophiolite, 620 Ivrea Zone, 679 Izu-Bonin-Mariana arc, 24, 567 Izu-Bonin-Mariana fore-arc, 583 Izu-Bonin-Mariana subduction zone, 568 J Jabal al Uwayjah ophiolite, 155–157 Jabal Ess ophiolite, 130–136 Jabal Tays ophiolite, 152–155 Jalair-Nayman, 713 Jamestown ophiolite, 560 Jamestown ophiolite complex, 606 Japan, 10 Jormua ophiolite, 14 Jormua ophiolite complex, 36 Junggar-Alakol-Turfan system, 698 Junggar-Balkhash, 709 K Kaapvaal craton, 8, 23, 539 Kaapvalley allochthon, 606 Kainuu Schist Belt, 38 Kalgoorlie, 524 Kalgoorlie terrane, 9, 605 Kam Group, 368 Kamuikotan complex, 661 Kapuskasing structural zone, 579 Karelian, 38 Karelian Craton, 15, 35, 36, 38, 66, 428 Karelian province, 428 Kazakhstan-Tien Shan, 692 Kazakhstan-Tien Shan domain, 698 keirogen, 692 keirogenic movement, 699 Kenya, 99 Khangai-Khantey mountains, 692 Khangai-Khantey subduction-accretion complex, 712 Khangay-Khantey, 710 Khantaishir, 706 Khizovaara boninite series, 434, 464 Khizovaara greenstone structure, 429 Khizovaara Structure, 426, 430

743

kimberlite, 67, 609 Kings-Kaweah ophiolite, 711 Kipchak arc, 694, 696 Knight Island ophiolite, 630, 637 Komati Formation, 544, 606 Komatii river, 608 komatiite, 23, 306, 410, 494, 501, 539–561, 577, 728 komatiite alteration, 546 komatiite series, 551 komatiite water contents, 550 komatiitic basalt, 409, 704 Kromberg allochthon, 608, 609 Kukasozero belt, 429 Kula-Farallon ridge, 634, 654 Kurtushiba ophiolite, 707 Kuznetskii Alatau, 704 L Lagonoy ophiolite, 619 Lake Irinozero, 443 lamprophyre, 41, 61 Lapland-Kola province, 426 lava, 35 lava unit, 443 layered gabbro, 2, 50, 110, 143, 171, 174 layered mafic rocks, 444 Leopold Kober, 681 leucogabbro, 184 lherzolite, 5, 110, 137, 146, 293 lherzolitic subcontinental mantle, 729 Liaoxi ophiolitic mélange, 322 Limpopo belt, 490 Limpopo mobile belt, 516 listwaenite, 118, 136, 141, 148, 157 listwanitization, 173 lode gold, 663 low-angle shear zone, 506 low-Ti tholeiite, 551 Luzon volcanic arc, 617 M mafic and ultramafic sills, 420 mafic metavolcanics, 409 mafic mylonite, 363, 371, 379 magma chamber, 6, 610 magma flow structure, 332 magmatic arc front, 689 magmatic arc system, 218 magmatic flow structure, 327, 331 magmatic front, 692 magmatic layering, 77, 325

744

magnesian andesite, 568, 663 magnesite, 179 magnesium number, 54, 113, 199, 216, 456, 458, 546, 548, 549, 569, 581, 583 Maikain-Balkybek ophiolitic suture, 713 Makran, 662 Makran accretionary complex, 699 Manjeri Formation, 498–501, 527, 528 mantle cooling, 561 mantle deformation, 263 mantle diapir, 36 mantle flow, 134, 250 mantle lithosphere, 609 mantle melting process, 728 mantle peridotite, 42, 53, 600, 619 mantle plume, 65, 540 mantle root, 732 mantle tectonite, 39, 43, 53, 146, 356, 731 mantle temperature, 264, 539, 541, 728, 731 mantle wedge, 468, 478, 569, 583, 584, 588, 591 mantle xenolith, 67 Marianas, 678 martitization, 182 Mashaba ultramafic suite, 517 Mashaba-Chibi dyke, 524 Mashaba-Chibi dyke swarm, 524 massive chromitite, 55, 252 massive sulfide, 85, 238 massive sulfide deposit, 343 massive tholeiite, 608 Mazatzal, 90 Mazatzal block, 87, 89 Mazatzal crustal block, 75, 88 Mazatzal orogeny, 86 Mberengwa allochthon, 499, 510, 512, 528, 529 McHugh complex, 628, 658 mélange, 9, 12, 17, 26, 131, 136, 141, 142, 150, 152, 172, 207, 210, 212, 224, 231, 234, 235, 242, 262, 276, 339–341, 443, 446, 447, 450, 516, 517, 600, 605, 628, 657, 658, 662, 706, 708–710, 712, 715, 731, 732 melt channels, 248, 263 melt-rock reaction, 263 melting process, 539 metagabbro, 147, 605 metamorphic sole, 119, 443, 447, 448, 450 metamorphic tectonite, 172, 244 metamorphosed, 731 metazoa, 118 microgabbro, 176, 210, 706 mid-ocean ridge basalt, 645

Subject Index

Middle Marker, 544 migration of magmatic arc fronts, 676 Moho, 2 molasse, 265 Mongol-Okhotsk sector, 692, 699 Mongolia, 207 Mongolia-Okhotsk intracontinental collision, 661 Montalban ophiolite, 619 MORB, 236, 368, 442, 448, 645, 647, 708 MORB-like crust, 619 MORB-like pillow lava, 218 MORB-type ophiolite, 709 Mozambique belt, 164 Mozambique Ocean, 96, 159, 730 Mtshingwe Group, 490, 494–498, 527 mylonite, 506, 510 mylonitized, 146 N N-MORB, 47, 217, 377 N-MORB basalt, 433 Najd fault system, 99 Nankai accretionary wedge, 662 nappe, 65, 95 narrow ocean basin, 215 Nb-enriched basalt, 568 near-trench magma, 732 near-trench magmatism, 13, 634 near-trench pluton, 12, 657 near-trench plutonic rocks, 657 Neoproterozoic, 130 Ngezi Group, 22, 490, 498–530 niobium-enriched basalt, 581 nodular, 55 nodular chromite, 249, 321 nodular chromitite, 248, 250 North China, 275, 340 North China craton, 18, 226, 228, 283, 322, 340, 600, 731 North Karelia schist belt, 36 North Karelian greenstone belt, 22, 425–481, 579 North Karelian terrane, 567 North Korea, 226 North Sayan, 707 O Ob-Zaysan-Surgut, 709 obduction, 7, 99, 265, 732 ocean floor basement remnant, 714 ocean floor remnant, 677 ocean floor tholeiite, 645

Subject Index

ocean floor tholeiitic basalt-komatiite association, 580 ocean-continent transition, 35 oceanic basement remnant, 714 oceanic crust, 677 oceanic mantle, 250, 264, 321 oceanic plate stratigraphy, 10 oceanic plateau, 4, 5, 121, 528, 529, 577, 580, 731 ocelli, 610 OIB, 47 Olekma gneiss-greenstone terrain, 406 olivine, 545 olivine composition, 104 olivine gabbro, 210, 211, 240, 246, 293 olivine spinifex komatiite, 544 Olondo greenstone belt, 21, 405, 407–423 one-way chilling, 4 Onib ophiolite, 169 Onib-Sol Hamed ophiolite-decorated suture, 200 Onib-Sol Hamed suture, 163 Onverwacht allochthon, 606, 608 Opatica, 579 Opatica subprovince, 579 ophiolite, 560, 657 ophiolite conundrum, 21, 364 ophiolite nappe, 703 ophiolite sequence, 3 ophiolite thickness, 609 ophiolites and orogeny, 676 ophiolitic mélange, 242, 265, 340, 678 ophiolitic nappe, 443, 447 ophiolitic suite, 170 ophirag, 688, 708, 710, 712, 714, 715 orbicular chromite, 250, 321 orbicular chromitite, 55, 249 Orca Group, 630 Ordos block, 227 orogenic belt, 676 orogenic complex, 732 orogenic lherzolite, 42, 43, 62 orthopyroxenite, 134 osmium, 276 P Palaeo-Asian Ocean, 207 Pan-African evolution, 699 Pannotia, 96 partial, 731 partial melting of oceanic slab, 478 passive margin, 35, 65 passive margin ophiolite, 36

745

Patterson Lake structural complex, 379 Payson ophiolite, 15 pelagic rocks, 137 pelagic sediments, 116, 177 Penninic ophiolite, 679 Penrose-style, 170 Penrose-type ophiolite, 2, 99 peridotite, 39, 133, 141, 143, 146, 148, 155, 171, 172, 240, 244, 289, 412, 429, 605, 635, 707 peridotite massif, 248 peridotite xenolith, 609 Philippine arc, 89 Philippine fault system, 694 Philippine Sea plate, 616 Philippines, 25, 584, 615–623, 734 picrite, 663 picritic rocks, 433 Pilbara craton, 10 pillow, 35, 80 pillow basalt, 4, 112, 134, 140, 146, 296, 497, 504, 634, 636, 637, 706 pillow lava, 7, 39, 47, 140, 172, 176, 177, 186, 210, 236, 243, 248, 371, 379, 409, 444, 501, 502, 571, 600, 606, 619, 638 pillow structure, 410 pillowed tholeiitic basalt, 429 plagiogranite, 52, 120, 172, 184, 210, 636, 706 plate boundary length, 662 plate tectonics, 225 plateau, 710 platinoid, 182 plume, 561 plume model, 540 plume-arc interaction, 583 plume-arc scenario, 542 podiform chromite, 19, 110, 182, 263, 275, 285, 321, 324, 728 podiform chromitite, 39, 55, 200, 224, 235, 243, 262, 334, 600 Point Lake, 20 Point Lake belt, 366 Point Lake greenstone belt, 8, 605 Point Lake volcanic belt, 371 polyphase deformation, 225 porcelainite, 84 pore fluid pressure, 661 post-orogenic intrusive suite, 623 pre-arc spreading, 677 primitive island arc, 619 Prince William terrane, 630 Purtuniq ophiolite, 14

746

pyroxene spinifex komatiite, 545 pyroxenite, 110, 155, 172, 181, 210, 216, 240, 247, 289, 290, 293, 606, 608, 635 Q Qinling-Dabie Shan orogen, 226 R radiolarite, 680 Raquette Lake Formation, 385 Re-Os, 417 Re-Os model age, 275, 280 recumbent folding, 450 Red Sea, 96, 730 Red Sea Hills, 17, 163, 201 Reliance Formation, 498, 501–504 residual mantle, 216 Resurrection Peninsula ophiolite, 26, 630 rhenium, 276 rhyolite, 345, 478, 663 rhyolitic dyke, 433 ribbon chert, 172, 177 ridge length, 732 ridge subduction, 13, 14, 657, 658, 664, 665 rifted continental margin, 729 rodingite, 44, 48, 60, 119, 147, 150 Rodinia, 96, 159 Russian craton, 694 S Sanak-Baranof belt, 633 Sangilen, 706 sanukitoid, 580, 592, 664 Saudi Arabia, 98, 730 sea-floor alteration, 194 sea-floor hydrothermal alteration, 187, 452 sea-floor hydrothermal vents, 225 sea-floor metamorphism, 44, 150 seamounts, 710 Sebakwe protocraton, 493 Sebakwian Group, 488, 526 secular change, 732 secular change in the melting conditions in subduction zone, 561 sediment contamination, 647 sediment subduction, 207, 215, 561, 588 sedimentary cap, 177 serpentinite, 118, 147, 152, 155 serpentinitic mélange, 706, 711 serpentinization, 179 serpentinized peridotite, 172

Subject Index

Shabani gneiss complex, 490, 493 Shabani ultramafic complex, 517, 524 shallow subduction, 567 Shamvaian Group, 527 shear zone, 8, 16, 130, 136, 140, 143, 173, 433, 489, 513, 516, 527, 650 shearing, 158 sheeted dike, 79, 111, 140, 185, 243, 479, 605, 619, 635, 637 sheeted dike complex, 4, 6, 9, 39, 42, 48, 78, 134, 176, 211, 236, 248, 296, 368, 389, 443, 444, 605, 706 sheeted intrusion, 608 sheeted mafic dyke complex, 172 sheeted mafic-ultramafic intrusion, 606 shoshonite, 577 Siberia, 10 Siberian craton, 406 siliceous unit, 156 siliceous zone, 83 silicification, 44, 173, 339, 343, 506 silicified, 141 sill, 605, 706 sinter vent complex, 346 slab fluid, 561 slab window, 658 Slave Craton, 20, 363–399 Slave Province, 8 Sleepy Dragon complex, 379–397 Sleepy Dragon terrane, 397 slow-spreading type ophiolite, 43 Snowball Earth, 118 Solonker Ocean, 699, 712 source composition, 548 South Gobi, 712 South Mongolian, 712 South Pass, 9 spinifex texture, 550 spreading rate, 610, 732 SSZ ophiolite, 178 staurolite, 435 staurolite-amphibolite, 438 staurolite-bearing amphibolite, 434, 435 staurolite-garnet geothermometer, 442 Steinmann trinity, 680 Stille’s geotectonic cycle, 685 stockwork feeder vein, 351 structural detachment, 728 structural dislocation, 265 structural repetition, 731 subcontinental lithospheric mantle, 35, 43, 67, 664

Subject Index

subducted continental sediments, 218 subducted sediments, 216 subduction back-stepping, 676 subduction erosion, 662 subduction mélange, 628, 662 subduction process, 560 subduction-accretion complex, 12, 580, 676, 692 Subgan granite-greenstone complex, 407 submarine hydrothermal alteration, 452 Sudan, 99, 163, 201 sulfide, 85 sulfide deposit, 339 sulfide mound, 349 sulfidic shear zone, 519 Sunda accretionary prism, 662 superimposed ophiolite, 615 Superior Province, 10, 23, 567, 569 364, 579 suprasubduction belt, 321 suprasubduction complex, 608 suprasubduction zone, 96, 120, 163, 200, 201, 264, 281, 283, 285, 364, 418, 425, 523, 528, 579, 615, 619, 732 suprasubduction zone ophiolite, 298, 317, 419, 426, 479 suprasubduction zone ophiolitic belt, 356 Susunai complex, 662 suture, 98, 143, 265 suture zone, 95, 136, 158, 676, 678 Svecofennian collision, 67 Svecofennian Ocean, 36 Svecofennian orogeny, 428 T Taito ophiolite, 657 tectonic ironstone, 501, 506, 514–516, 518 tectonic settings, 5 tectosphere, 227 Tekturmas, 709 term ophiolite, 687 Tethyan ophiolite terrane, 399 Tethyan-type, 165 Tethyan-type ophiolite, 20, 363, 366 Tethysides, 699 Tharwah ophiolite, 136–141 thermal plume, 540 thin-skinned thrusting, 506 tholeiite, 448 tholeiitic, 85, 113, 195, 377, 394, 644, 730 tholeiitic basalt, 368, 379, 409, 410, 433, 577 tholeiitic igneous rocks, 67 tholeiitic pillow basalt, 608

747

tholeiitic pillow lava, 605 thrust, 450 thrust complex, 732 Tien Shan accretionary complex, 708 Tihama Asir, 730 Tihama Asir, Saudi Arabia, 5 Tokwe gneiss, 493 Tokwe gneiss complex, 488 Tokwe segment, 525 Tokwe terrane, 493 tonalite, 83, 84, 471, 633, 664 tonalite-trondhjemite gneiss (TTG), 231, 406 tonalite-trondhjemite-granodiorite, 13, 25, 658 tonalite-trondhjemite-granodiorite batholith, 569 tonalite-trondhjemite-granodiorite pluton, 580, 627 tonalitic magma, 224 trace element, 298, 299 Trans-Hudson orogen, 567, 582 transform suture, 703 transition zone, 2, 10, 78, 110, 111, 171, 172, 240, 248, 263 transitional continental crust, 397 transitional IAT/MORB, 195 transitional ophiolite, 5, 729 transposition, 265 transpression, 158 trench axis, 656 triple junction, 647 triple-junction interaction, 732 triple-junction related magma, 732 troctolite, 201, 262 trondhjemite, 52, 184, 471, 636, 650, 664 Troodos, 464, 583 Troodos Massif in Cyprus, 677 Troodos ophiolitic lava, 469 TTG, 13, 228, 229, 569, 580, 592, 663 TTG magma, 664 TTG suite, 732 tuff, 649, 651 turbidite, 265, 448, 449, 571, 630, 649 Tuva, 207 Tuva-Mongol arc, 699 Tuva-Mongol continental fragment, 692, 712 Tuva-Mongolian Massif, 208 U ultramafic, 290, 293, 707 ultramafic cumulate, 240, 619 ultramafic layered complex, 606 ultramafic mylonite, 8, 510, 513, 514 ultramafic plutonic rocks, 409

748

ultramafic rocks, 289, 572, 635 ultramafic tectonite, 134, 241 ultramylonite, 147 Umtali line, 530 underplating, 661 uniformitarianism, 731 Ural/Yenisey margin, 694 uralitic amphibole, 184, 186 V Valdez Group, 630, 658 vanished ocean, 676 vent chimney, 348 Ventersdorp, 521 Ventersdorp Group, 521 Vermilion district, 579 viscosity, 6 VMS, 663 VMS deposit, 664 volcaniclastics, 620 volcanogenic massive sulfide, 340, 663 W Wabigoon, 10 Wadi Onib, 17 Wadi Onib mafic-ultramafic complex, 163, 165 Wadi Onib ophiolite, 17, 163–202 Wawa greenstone belt, 580 Wawa subprovince, 569 websterite, 146, 182, 293 wehrlite, 110, 134, 146, 181, 210, 211, 216, 240, 293, 606, 608 wehrlitic transition, 171, 181 West Greenland, 23, 567 Western Block, 226, 228, 275, 283, 322, 340 Western Sayan, 707 wet melting, 541

Subject Index

Wind River Range, Wyoming, 9 Wrangellia composite terrane, 628 Wutai mélange belt, 357 Wutai Mountains, 20, 340 Wutai ophiolitic mélange, 322 Wutai VMS, 354 X xenocryst, 62 xenocrystic zircon, 62, 141, 523 Y Yakutat terrane, 662 Yanbu suture, 132 Yavapai orogeny, 86, 87, 89, 90 Yavapai-Mazatzal orogenic belt, 73, 87 Yellow Sea, 226 Yellowknife, 20 Yellowknife belt, 366 Yellowknife greenstone belt, 605 Yellowknife volcanic belt, 368–379 Yilgarn craton, 9, 524 Yinshan-Yanshan orogen, 226 Z Zabargad Island, 5, 730 Zagros crush belt, 678 Zambales ophiolite, 619 Zambales range, 617 Zeederbergs Formation, 498, 502–504 Zimbabwe, 22 Zimbabwe craton, 488, 516, 526, 527 Zunhua ophiolitic mélange, 275, 281, 322 Zunhua podiform chromite, 19 Zunhua structural belt, 229, 231, 232, 234, 242, 275, 284, 322 Zunhua tectonic zone, 600