The boundaries of the West African craton

ERRATUM Special Publication 297 DABO , M., GUEYE , M., NGOM , P. M. & DIAGNE , M. 2008. Orogen-parallel tectonic transp...

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ERRATUM Special Publication 297 DABO , M., GUEYE , M., NGOM , P. M. & DIAGNE , M. 2008. Orogen-parallel tectonic transport: transpression and strain partitioning in the Mauritanides of NE Senegal. In: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 483– 497. DOI: 10.1144/SP297.23 The figure on page 485 is incorrect. The correct figure is given below. The caption is correct in the printed/ online version.

The Boundaries of the West African Craton

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) PHIL LEAT (UK) NICK ROBINS (UK) JONATHAN TURNER (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY )

RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) RANDELL STEPHENSON (NETHERLANDS )

IUGS/GSL publishing agreement This volume is published under an agreement between the International Union of Geological Sciences and the Geological Society of London and arises from IGCP485. GSL is the publisher of choice for books related to IUGS activities, and the IUGS receives a royalty for all books published under this agreement. Books published under this agreement are subject to the Society’s standard rigorous proposal and manuscript review procedures.

It is recommended that reference to all or part of this book should be made in one of the following ways: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) 2008. The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297. POUCLET , A., OUAZZANI , H. & FEKKAK , A. 2008. The Cambrian volcano-sedimentary formations of the westernmost High Atlas (Morocco): their place in the geodynamic evolution of the West African PalaeoGondwana northern margin. In: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 303– 327.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 297

The Boundaries of the West African Craton

EDITED BY

NASSER ENNIH El Jadida University, Morocco and

JEAN-PAUL LIE´GEOIS Royal Museum for Central Africa, Tervuren, Belgium

2008 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ44 (0)1225 445046, Fax þ44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2008. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of the Copyright Licensing Agency, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/08/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-251-9 Typeset by Techset Composition Ltd., Salisbury, UK Printed by Cromwell Press Ltd, Trowbridge, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel þ1 800-972-9892 Fax þ1 802-864-7626 E-mail [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel þ1 918-584-2555 Fax þ1 918-560-2652 E-mail [email protected] Website http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel þ91 11 2327-9113/2326-4180 Fax þ91 11 2326-0538 E-mail [email protected]

Preface This special publication was generated by the UNESCO International Geological Correlation Program (IGCP) 485 named ‘Cratons, metacratons and mobile belts; keys from the West African craton boundaries: Eburnian versus Pan-African signature, magmatic, tectonic and metallogenic implications’, in short, ‘The boundaries of the West African craton’, the title of this book. A main issue of this IGCP485 was to bring African geologists from different countries together with European and American geologists in the same African region and to promote geological research across political boundaries. The fact that this ambitious goal was attained in Saharan regions is due to the energy and enthusiasm of geologists, colleagues and friends, from the countries where these field trips were organized, who gave a great deal of their time to organize these huge events. We are extremely grateful to Ezzoura Errami, Abdelilah Fekkak from El Jadida University and Hassan Admou from Marrakech University (AntiAtlas field trip, Morocco 2003), Khalidou Loˆ, now Director of the Office mauritanien de recherche geólogique (Reguibat field trip, Mauritania 2004), Dramane Dembe´le´, now Director of the Direction Nationale de la Geólogie et des Mines and Samba Sacko, who sadly died a few weeks after the conference (Gourma field trip, Mali 2005), Khadidja Ouzegane and Abla Azzouni-Sekkal, professors at the Universite´ des Sciences et Techniques Houari Boumedienne a` Alger (Hoggar field trip, Algeria 2006), Hassan Ouanaimani from the Ecole Normale Supe´rieure, Marrakech and again Ezzoura Errami from El Jadida University (Anti-Atlas and High Atlas field trip, Morocco, 2007). These remote field trips provided an extraordinary crucible for African geologists and African geology. The fifth and final field trip in late 2007 completed a geological cross-section from the Anti-Atlas through the High Atlas to reach the Meseta in Morocco. The IGCP497 ‘The Rheic Ocean’ has also been associated with this project. Indeed a main conclusion of this study is that the boundaries of cratons are susceptible to reactivation or even to leave the craton as occurred when the peri-Gondwanan terranes left the West African craton, generating the Rheic Ocean. This Special Publication would not have been possible without the participation of the contributors to this volume. We are grateful for their submissions and their willingness to enhance their papers as much as possible. The support of many colleagues who acted as reviewers is also greatly appreciated. Their conscientious work was necessary to ensure

the high quality of this volume. We are indebted to them and we warmly thank them. The panel of reviewers for this volume were: J. Abati (Madrid, Spain) H. Admou (Marrakech, Morocco) A. Aghzer (El Jadida, Morocco) K. Attoh (Ithaca, USA) A. Azor (Granada, Spain) T. Berza (Bucharest, Romania) P. G. Betts (Melbourne, Australia) B. Bonin (Orsay, France) N. Bournas (Algiers, Algeria) F. Bussy (Lausanne, Switzerland) A. Calvert (Burnaby, Canada) C. Carrigan (Olivet Nazarene, USA) P. Cawood (Perth, Australia) R. J. Chapman (Leeds, UK) A. Cheilletz (Nancy, France) J. R. Cottin (St Etienne, France) D. Demaiffe (Brussels, Belgium) T. De Putter (Tervuren, Belgium) M. de Wit (Cape Town, South Africa) M. Doucoure´ (De Beers, South Africa) D. Dyck (Billiton, Canada) J. F. A. Diener (Stellenbosh, South Africa) M. El Houicha (El Jadida, Morocco) E. Errami (El Jadida, Morocco) D. Gasquet (Nancy, France) P. Goncalves (Besancon, France) K. Hefferan (Stevens Point, USA) J. Hibbard (North Carolina, USA) M. Higgins (Chicoutimi, Canada) N. Ilbeyli (Hatey, Turkey) R. Kerrich (Saskatchewan, Canada) P. Koenigshof (Mainz, Germany) A. Korsakov (Novosibirsk, Russia) M. Lehtonen (GTK, Finland) U. Linneman (Dresden, Germany) B. Litvinovsky (Beer-Sheva, Israe¨l) S. Lubeseder (Manchester, UK) A. C. Maalof (Cambridge, UK) C. Marignac (Nancy, France) R. Mapeo (Gaborone, Botswana) P. Marquer (Besancon, France) J. Martignole (Quebec, Canada) S. Master (Johannesburg, South Africa) S. Masur (Wroclav, Poland) R. P. Meńot (St Etienne, France) A. Michard (Paris, France) E. Moores (California, USA) P. Morzadec (Rennes, France) H. Ouanaimi (Marrakech, Morocco) Y. Osanai (Fukuoka, Japan)

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K. Ouzegane (Algiers, Algeria) C. Passchier (Mainz, Germany) J. J. Peucat (Rennes, France) A. Pouclet (Orleáns, France) A. Preat (Brussels, Belgium) G. Rebay (Pavia, Italy) R. Resmini (Fairfax, USA) K. Sato (Tokyo, Japan)

PREFACE

A. Schmidt Mumm (Adelaide, Australia) J. Schumacher (Bristol, UK) J. Ugidos (Salamanca, Spain) A. Wilde (Monash, USA) F. Y. Wu (Changchun, China) NASSER ENNIH JEAN -PAUL LIE´ GEOIS

Contents Preface ENNIH , N. & LIE´ GEOIS , J.-P. The boundaries of the West African craton, with special reference to the basement of the Moroccan metacratonic Anti-Atlas belt The Palaeoproterozoic terranes from the West African craton and their reworking SOUMAILA , A., HENRY , P., GARBA , Z. & ROSSI , M. REE patterns, Nd–Sm and U –Pb ages of the metamorphic rocks of the Diagorou– Darbani greenstone belt (Liptako, SW Niger): implication for Birimian (Palaeoproterozoic) crustal genesis KEY , R. M., LOUGHLIN , S. C., GILLESPIE , M., DEL RIO , M., HORSTWOOD , M. S. A., CROWLEY , Q. G., DARBYSHIRE , D. P. F., PITFIELD , P. E. J. & HENNEY , P. J. Two Mesoarchaean terranes in the Reguibat shield of NW Mauritania KOLB , J., MEYER , F. M., VENNEMANN , T., HOFFBAUER , R., GERDES , A. & SAKELLARIS , G. A. Geological setting of the Guelb Moghrein Fe oxide–Cu–Au– Co mineralization, Akjoujt area, Mauritania KAHOUI , M., MAHDJOUB , Y. & KAMINSKY , F. V. Possible primary sources of diamond in the North African diamondiferous province BENDAOUD , A., OUZEGANE , K., GODARD , G., LIE´ GEOIS , J.-P., KIENAST , J.-R., BRUGUIER , O. & DRARENI , A. Geochronology and metamorphic P–T –X evolution of the Eburnean granulitefacies metapelites of Tidjenouine (Central Hoggar, Algeria): witness of the LATEA metacratonic evolution ADJERID , Z., OUZEGANE , K., GODARD , G. & KIENAST , J. R. First report of ultrahightemperature sapphirine þ spinel þ quartz and orthopyroxene þ spinel þ quartz parageneses discovered in Al –Mg granulites from the Khanfous area (In Ouzzal metacraton, Hoggar, Algeria) The Pan-African orogeny along the boundaries of the West African craton VILLENEUVE , M. Review of the orogenic belts on the western side of the West African craton: the Bassarides, Rokelides and Mauritanides CABY , R., BUSCAIL , F., DEMBE´ LE´ , D., DIAKITE´ , S., SACKO , S. & BAL , M. Neoproterozoic garnet-glaucophanites and eclogites: new insights for subduction metamorphism of the Gourma fold and thrust belt (eastern Mali) ATTOH , K. & NUDE , P. M. Tectonic significance of carbonatite and ultrahigh-pressure rocks in the Pan-African Dahomeyide suture zone, southeastern Ghana BOUSQUET , R., EL MAMOUN , R., SADDIQI , O., GOFFE´ , B., MO¨ LLER , A. & MADI , A. Me´langes and ophiolites during the Pan-African orogeny: the case of the Bou-Azzer ophiolite suite (Morocco) BELKABIR , A., JE´ BRAK , M., MAACHA , L., AZIZI SAMIR , M. R. & MADI , A. Gold mineralization in the Proterozoic Bleida ophiolite, Anti-Atlas, Morocco TOUIL , A., HAFID , A., MOUTTE , J. & EL BOUKHARI , A. Petrology and geochemistry of the Neoproterozoic Siroua granitoids (central Anti-Atlas, Morocco): evolution from subductionrelated to within-plate magmatism The late Neoproterozoic –early Palaeozoic extension along the West African craton and the Peri-Gondwanan terranes ´ LVARO , J. J., MACOUIN , M., EZZOUHAIRI , H., CHARIF , A., AIT AYAD , N., RIBEIRO , M. L. & A ADER , M. Late Neoproterozoic carbonate productivity in a rifting context: the Adoudou Formation and its associated bimodal volcanism onlapping the western Saghro inlier, Morocco POUCLET , A., OUAZZANI , H. & FEKKAK , A. The Cambrian volcano-sedimentary formations of the westernmost High Atlas (Morocco): their place in the geodynamic evolution of the West African Palaeo-Gondwana northern margin

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77 111

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169 203

217 233

249 265

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CONTENTS

EZZOUHAIRI , H., RIBEIRO , M. L., AIT AYAD , N., MOREIRA , M. E., CHARIF , A., RAMOS , J. M. F., DE OLIVEIRA , D. P. S. & COKE , C. The magmatic evolution at the Moroccan outboard of the West African craton between the Late Neoproterozoic and the Early Palaeozoic NANCE , R. D., MURPHY , J. B., STRACHAN , R. A., KEPPIE , J. D., GUTIE´ RREZ -ALONSO , G., FERNA´ NDEZ -SUA´ REZ , J., QUESADA , C., LINNEMANN , U., D’LEMOS , R. & PISAREVSKY , S. A. Neoproterozoic–early Palaeozoic tectonostratigraphy and palaeogeography of the peri-Gondwanan terranes: Amazonian v. West African connections PEREIRA , M. F., CHICHORRO , M., WILLIAMS , I. S. & SILVA , J. B. Zircon U– Pb geochronology of paragneisses and biotite granites from the SW Iberian Massif (Portugal): evidence for a palaeogeographical link between the Ossa –Morena Ediacaran basins and the West African craton GU¨ RSU , S. & GONCUOGLU , M. C. Petrogenesis and geodynamic evolution of the Late Neoproterozoic post-collisional felsic magmatism in NE Afyon area, western central Turkey The Variscan orogeny along the West African craton SOULAIMANI , A. & BURKHARD , M. The Anti-Atlas chain (Morocco): the southern margin of the Variscan belt along the edge of the West African craton BAIDDER , L., RADDI , Y., TAHIRI , M. & MICHARD , A. Devonian extension of the Pan-African crust north of the West African craton, and its bearing on the Variscan foreland deformation: evidence from eastern Anti-Atlas (Morocco) OUANAIMI , H. & LAZREQ , N. The ‘Rich’ group of the Draâ Basin (Lower Devonian, Anti-Atlas, Morocco): an integrated sedimentary and tectonic approach DABO , M., GUEYE , M., NGOM , P. M. & DIAGNE , M. Orogen-parallel tectonic transport: transpression and strain partitioning in the Mauritanides of NE Senegal The Cenozoic situation along the boundaries of West African craton ATTOH , K. & BROWN , L. Deep structure of the southeastern margin of the West African craton from seismic reflection data, offshore Ghana BERGER , J., ENNIH , N., LIE´ GEOIS , J.-P., NKONO , C., MERCIER , J.-C. C. & DEMAIFFE , D. A complex multi-chamber magmatic system beneath a late Cenozoic volcanic field: evidence from CSDs and thermobarometry of clinopyroxene from a single nephelinite flow (Djbel Saghro, Morocco) Index

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The boundaries of the West African craton, with special reference to the basement of the Moroccan metacratonic Anti-Atlas belt NASSER ENNIH1 & JEAN-PAUL LIE´GEOIS2 1

Geodynamic Laboratory, El Jadida University, BP. 20, 24000, El Jadida, Morocco (e-mail: [email protected]) 2

Isotope Geology, Royal Museum for Central Africa, B-3080 Tervuren, Belgium (e-mail: [email protected]) Abstract: The West African craton (WAC) was constructed during the Archaean and the c. 2 Ga Palaeoproterozoic Eburnian orogeny. Mesoproterozoic quiescence at c. 1.7–1.0 Ga allowed cratonization. In the absence of Mesoproterozoic activity, there are no known WAC palaeogeographical positions for that time. At the beginning of the Neoproterozoic, the WAC was affected by several extensional events suggesting that it was subjected to continental breakup. The most important event is the formation of the Gourma aulacogen in Mali, and the Taoudeni cratonic subcircular basin and deposition of platform sediments in the Anti-Atlas. At the end of the Neoproterozoic, the WAC was subjected to convergence on all its boundaries, from the north in the Anti-Atlas, to the east along the Trans-Saharan belt, to the south along the Rockelides and the Bassarides and to the east along the Mauritanides. This led to a partial remobilization of its cratonic boundaries giving rise to a metacratonic evolution. The WAC boundaries experienced Pan-African Neoproterozoic to Early Cambrian transpression and transtension, intrusion of granitoids and extrusion of huge volcanic sequences in such as in the Anti-Atlas (Ouarzazate Supergroup). Pan-African tectonism generated large sediment influxes around the WAC within the Peri-Gondwanan terranes whose sedimentary sequences are marked by distinctive zircon ages of 1.8– 2.2 Ga and 0.55 –0.75 Ga. WAC rocks experienced Pan-African low grade metamorphism and large movements of mineralizing fluids. In the Anti-Atlas, this Pan-African metacratonic evolution led to remobilization of REE in the Eburnian granitoids due to the activity of F-rich fluids linked to extrusion of the Ouarzazate Supergroup. During the Phanerozoic, the western WAC boundary was subjected to the Variscan orogeny, for which it constituted the foreland and was, therefore moderately affected, showing typical thick-skin tectonics in the basement and thin-skin tectonics in the cover. During the Mesozoic, the eastern and southern boundaries of the WAC were subjected to the Atlantic opening including Jurassic dolerite intrusion and capture of its extreme southern tip by South America. The Jurassic is also marked by the development of rifts on its eastern and northern sides (future Atlas belt). Finally, the Cenozoic period was marked by the convergence of the African and European continents, generating the High Atlas range and Cenozoic volcanism encircling the northern part of the WAC. The northern metacratonic boundary of the WAC is currently uplifted, forming the Anti-Atlas Mountains. The boundaries of the WAC, metacratonized during the Pan-African orogeny have been periodically rejuvenated. This is a defining characteristic of the metacratonic areas: rigid, stable cratonic regions that can be periodically cut by faults and affected by magmatism and hydrothermal alteration – making these areas important for mineralization.

A craton is a stable part of the continental lithosphere which has not been deformed for a long time (Bates & Jackson 1980). Although cratons are not tectonically active, they can be located near active margins, such as the Brazilian craton at the rear of the Andean active margin. Cratons proximal to collision zones act as a shield, as their thick lithosphere protects them from most of the collisional effects (Black & Lie´geois 1993). However, cratons can be partly subducted or affected by transpressive tectonics. Such partial reactivation of a rigid body or of the boundaries

of a rigid body gives rise to geological characteristics different from the cratonic quiescence as well as from the intense activity occurring in a mobile belt. This has been called metacratonic evolution (Abdelsalam et al. 2002). This special volume has been generated by the UNESCO IGCP485 (International Geological Correlation Programme, now International Geoscience Programme) called Cratons, metacratons and mobile belts: keys from the West African craton boundaries; Eburnian versus Pan-African signature, magmatic, tectonic and metallogenic

From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 1– 17. DOI: 10.1144/SP297.1 0305-8719/08/$15.00 # The Geological Society of London 2008.

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N. ENNIH & J.-P. LIEGEOIS

implications. The aim of this project, and of this book, was to encompass the whole evolution of the boundaries of the West African craton, from the Archaean/Palaeoproterozoic towards Recent times. The IGCP485 organized field conferences in remote areas such as the Reguibat Rise in Mauritania, the Gourma region in Mali, the Hoggar shield in Algeria and twice in the Anti-Atlas belt in Morocco. This book contains twenty-four papers concerning these regions and other boundaries of the West African craton.

The West African craton The West African craton (WAC) is composed of three Archaean and Palaeoproterozoic metamorphic and magmatic shields separated by two cratonic sedimentary basins (Fig. 1). The WAC components include: to the south the Man shield, to which the smaller Kayes and Kenieba inliers can be associated; to the north the Reguibat shield; and to the extreme north, the Anti-Atlas belt. In between are, in the centre, the huge Taoudeni basin and to the north the Tindouf basin. The Man and Reguibat shields comprise Archaean nuclei to the west

(Feybesse & Mile´si 1994; Potrel et al. 1998; Key et al.). In the Man shield, a large part of the WAC consists of the Palaeoproterozoic Birimian continent (Abouchami et al. 1990; Lie´geois et al. 1991; Boher et al. 1992; Soumaila et al.). The Reguibat shield contains Palaeoproterozoic assemblages in the eastern part as well as Archaean components that include kimberlites (Kahoui et al.). The AntiAtlas belt comprises only a Palaeoproterozoic basement (Thomas et al. 2004). The WAC terranes were affected by the Eburnian orogeny, at around 2 Ga. During the Mesoproterozoic, no event or rocks are known in the WAC. This extremely quiet period between 1.7 Ga and 1 Ga allowed this large area to become a craton by acquiring a thick lithosphere (Black & Lie´geois 1993). The palaeoposition of the WAC during the Mesoproterozoic is not known. By contrast to the Mesoproterozoic, the Neoproterozoic was an important period of evolution for the West African craton. At the beginning of the Neoproterozoic, the WAC was affected by several extensional events associated with continental breakup. The most important event was the formation of the Gourma aulacogen in Mali (MoussinePouchkine & Bertrand-Sarfati 1978) but also the deposition of passive margin sediments on its

Fig. 1. Main geological units in West Africa, from Fabre (2005) and Lie´geois et al. (2005). The positions and shapes of the Peri-Gondwanan terranes are based on Nance et al.

THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT

northern boundary in the Anti-Atlas (Bouougri & Saquaque 2004) and to the SE in the Voltas basin (Ako & Wellman 1985; Ne´de´lec et al. 2007). The formation of the huge Taoudeni cratonic subcircular basin also began during that period (Bronner et al. 1980). At the end of the Neoproterozoic, the WAC was first subjected to island arc accretion during 760 –660 Ma on its northern and eastern sides, in the Moroccan Anti-Atlas (Thomas et al. 2002; Samson et al. 2004; D’Lemos et al. 2006; Bousquet et al.), in the Malian Tilemsi (Caby et al. 1989) and in the Gourma area (De la Boisse 1979). During the main Pan-African orogenic phase, the WAC was subjected to convergence on all its boundaries, from the north in the Anti-Atlas (Hefferan et al. 2000; Ennih & Lie´geois 2001), to the east along the Trans-Saharan belt (Black et al. 1979, 1994; Affaton et al. 1991; Attoh & Nude), to the south with the Rockelides and the Bassarides belts and to the east with the Mauritanides belt (Villeneuve), with several thrust sheets preserved on the craton itself, such as in Mali (Caby et al.). These collisions partly remobilized other cratonic regions to the east of the WAC in the Tuareg shield (Lie´geois et al. 2003; Bendaoud et al.; Adjerid et al.) and also the peri-Gondwanan terranes (Nance et al.; Pereira et al.; Gu¨rsu & Gonçuog˘lu).

The boundaries of the West African craton The Pan-African orogeny induced a partial remobilization of the WAC boundaries, inducing a metacratonic evolution. At the end of the Neoproterozoic, the West African craton belonged to the subducting plate, implying that it was never an active margin above a subduction oceanic plate (Black et al. 1979; Hefferan et al. 2000; Ennih & Lie´geois 2001; Villeneuve), with craton-ward directed fold-and-thrust structures (Jahn et al. 2001; Caby et al.) also forming outboard of the craton (Attoh & Brown). The West African craton acted as a rigid indenter during the Pan-African orogeny in a similar way as India is currently indenting the Asian continent (Black et al. 1979). The thick cratonic continental lithosphere was partly affected by transpression and transtension tectonic episodes, intrusion of granitoids and extrusion of volcanic sequences, such as in the Anti-Atlas (Ouarzazate Supergroup). High temperature/low pressure grade metamorphism occurred with large movements of fluids causing mineralization in several areas (Inglis et al. 2004; Kolb et al.; Belkabir et al.). Although partially buried by younger deposits, rock exposures contain excellent preservation of the early

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Neoproterozoic events such as the passive margin sediments (Bouougri & Saquaque 2004), the early thrust oceanic terranes (Samson et al. 2004; Bousquet et al.) and Pan-African magmatism (Touil et al.; Ezzouhairi et al. ). The Pan-African orogeny generated large sediment influx, for example outside the WAC, within the PeriGondwanan terranes whose sedimentary sequences are marked by the WAC-typical signature, a bimodal set of zircon ages at c. 1.8–2.2 Ga and at 0.55 –0.75 Ga (Nance et al.). The metacratonic evolution affected some parts of the WAC and periGondwanan terranes such as the c. 2 Ga Icartian gneisses in Brittany (Calvez & Vidal 1978; Samson & D’Lemos 1998). A better knowledge of the boundaries of the West African craton is, therefore, of importance for the study of periGondwanan terranes in Europe and in America. During the latest Neoproterozoic and early Palaeozoic, the western WAC boundaries were first subjected to a major extensional event producing sedimentary and volcanic sequences (Alvaro et al.; Pouclet et al.) and the drifting of some of the peri-Gondwanan terranes (Nance et al.). Thick Phanerozoic sedimentary sequences were deposited afterwards up to Devonian times (Baidder et al.; Ouanaimani & Lazreq). The Late Palaeozoic Variscan orogeny moderately affected the WAC by generating thick-skin tectonics in the basement and thin-skin tectonics in the sedimentary cover rocks (Caritg et al. 2004; Burkhard et al. 2006; Baidder et al.; Dabo et al.; Soulaimani & Burkhard). In the Mesozoic, the western and southern WAC boundaries were subjected to the Atlantic rifting and Jurassic dolerite intrusions and massive Central Atlantic magmatic province (CAMP) basalt flows (Marzoli et al. 1999; Deckart et al. 2005; Verati et al. 2005). The Jurassic was also marked by the development of rifts on the eastern side (e.g. Gao rift) and on the northern side (e.g. Atlas belt), contemporaneously with the development of the Central Atlantic Ocean and the Western Mediterranean Sea (Laville et al. 2004; Guiraud et al. 2005). Finally, the Cenozoic era was marked by the convergence of the African and European continents, generating the High Atlas range, the uplift of the Anti-Atlas (Malusa et al. 2007) and the Cenozoic volcanism in West Africa (Berger et al.) and the Hoggar (Lie´geois et al. 2005). The northern metacratonic boundary of the WAC is currently uplifted, forming the AntiAtlas Mountains. The boundaries of the WAC, metacratonized during the Pan-African orogeny, have been rejuvenated periodically. The rigid WAC metacraton has been affected by the reactivation of lithospheric faults that facilitate hydrothermal mineralization (Pelleter et al. 2007). This is the main characteristic

4


of the metacratonic areas (Lie´geois et al. 2003, 2005): being rigid but affected by faults of lithospheric scale, they constitute areas subjected to reactivation, including intraplate reactivations (Azzouni-Sekkal et al. 2003; Lie´geois et al. 2005), making them areas likely to be rich in mineralizations.

The case study of the remobilization of the Eburnian basement in the Anti-Atlas belt Our recent study on the Eburnian basement of the Anti-Atlas, in the Zenaga inlier (Figs 2 and 3) has revealed extensive REE mobility in Eburnian granites during the Pan-African orogeny. REE mobility is attributed to Pan-African transcurrent tectonism and associated Neoproterozoic Ouarzazate Supergroup volcanism.

The Anti-Atlas belt The Anti-Atlas belt (Fig. 2) is separated in two parts by the Anti-Atlas major fault (AAMF) – long considered as the northern limit (e.g. Hefferan et al. 2000) of the WAC because it is marked by ophiolitic remnants, including that of Bou Azzer, and because Eburnian outcrops are not known north of it in the Saghro Mountains. For various geological but also rheological and isotopic reasons, Ennih & Lie´geois (2001) proposed that the actual northern boundary of the WAC is the South Atlas fault that borders the High Atlas mountain range to the south (Fig. 2) of the AAMF. According to Ennih & Lie´geois (2001), the South Atlas Fault marks the edge of the deepening of the WAC basement under Neoproterozoic volcano-sedimentary series. Here we will focus on the Zenaga inlier, which consists of Eburnian basement rocks and is located just to the south of the AAMF, for deciphering the Pan-African effects on the WAC northern boundary basement.

The Zenaga inlier The Zenaga inlier is a depression of about 500 km2 containing mainly Palaeoproterozoic gneisses and granitoids unconformably overlain by the late Neoproterozoic Ouarzazate volcanic Supergroup or by the Cambrian Tata Group (Fig. 3). Within the inlier, Neoproterozoic rocks also consist of passive margin sediments (Taghdout Group), pre-Pan-African doleritic dykes and sills and a late Pan-African alkaline ring-complex. Along the AAMF, to the west and to the east, remnants of the Bou Azzer –Sirwa oceanic terrane are present. A summary of the geology of the area helps to

understand the metacratonic evolution of the Zenaga basement. The Zenaga Palaeoproterozoic metamorphic rocks include medium to high-grade amphibolite facies grey gneisses, biotite-rich schists, garnet + sillimanite paragneisses, calc-silicate rocks, migmatites and rare amphibolites. The gneissic layering and the migmatitic leucosomes are deformed by isoclinal ductile folds, whose axes have variable plunge. This basement represents a high-grade metamorphic supracrustal series. These schists have not been dated but inherited zircons at c. 2170 Ma within the c. 2035 Ma cross-cutting granitoids could be attributed to the Zenaga schists (Thomas et al. 2002). The Zenaga Palaeoproterozoic granitoids are represented by the Azguemerzi mesocratic granodiorite, and the Aı¨t Daoui, Assersa, Tamarouft and Tazenakht granites. The Zenaga plutons show quartzo-feldspathic layers separated by biotite and garnet layers, locally associated with gneisses and anatectic products. They contain rare metasedimentary xenoliths and no mafic microgranular enclaves (MME). The presence of aluminous minerals (biotite, garnet, muscovite), the association with migmatitic rocks, the absence of MME suggest that the Zenaga granitoids originated by the partial melting of crustal rocks. The granodiorite and the granites have been dated at 2037 + 7 Ma, 2037 + 9 Ma and 2032 + 5 Ma (U– Pb zircon ages, Thomas et al. 2002). These dates give a minimum age for the gneisses and schists. The Zenaga granitoid basement is unconformably covered by the Taghdout sedimentary Group, also known as the Tizi n-Taghatine Group (Thomas et al. 2004). The Taghdout Group displays brittle tectonic faults folds associated with a southverging thrust event. Portions of this unit have been metamorphosed to greenschist facies. The Taghdout Group contains well-preserved sedimentary features, such as ripple marks, desiccation cracks or oblique stratification (Bouougri & Saquaque 2004). The Taghdout Group is a 2 km-thick succession deposited during three stages of an extensional event (Bouougri & Saquaque 2004): (1) a shallowwater and gently dipping mixed siliciclastic – carbonate ramp facing north and attached to braided alluvial plain in the south, indicating a relatively stable margin; (2) tholeiitic sills and dykes of the Ifzwane Group that cut the sedimentary sequence and the basement of the Zenaga inlier (particularly to the NW; Fig. 3a); (3) deepening of the margin marked by turbidites. Although they are not directly in contact with the Zenaga basement, remnants of the Bou Azzer and Sirwa oceanic island arc complex occur along the AAMF (Fig. 2). This complex comprises ophiolitic sequences in which plagiogranites have been


Fig. 2. (a) Geological map of the Anti-Atlas (Morocco), from Thomas et al. (2004) and Gasquet et al. (2008). The rectangle outlines the area in Figure 1B. (b) Satellite photograph of the Anti-Atlas (Orthorectified Landsat Thematic Mapper Mosaics as compressed colour imagery in MrSIDTM file format from Lizardtech).

5

6


Fig. 3. (a) Satellite photograph of the Palaeoproterozoic Zenaga inlier (Orthorectified Landsat Thematic Mapper Mosaics as compressed colour imagery in MrSIDTM file format from Lizardtech). The inlier is noted ‘Zenaga’ on Figure 2a and 2b. (b) Sketch geological map of the Zenaga inlier (from Ennih & Lie´geois 2001). The scales of the satellite image and of the geological map are the same.

dated at 761 + 2 Ma and 762 + 2 Ma at Taswirine (Sirwa area; U –Pb zircon; Samson et al. 2004) and a tonalitic migmatite at 743 + 14 Ma at Iriri (Sirwa area; U –Pb zircon; Thomas et al. 2002). The zircon rims of the latter gave an age of 663 + 13 Ma, interpreted as the age of the metamorphism that accompanied the island arc accretion towards the craton (Thomas et al. 2002). In Bou Azzer, juvenile metagabbros (752 + 2 Ma), augen granite gneiss (753 + 2 Ma) and leucogranites (705 + 3 Ma; 701 + 2 Ma) are linked to this

750–700 Ma event but in a way still to be deciphered (D’Lemos et al. 2006). The Zenaga basement is overthrust by the Tamwirine rhyolitic unit, attributed to the Bou Salda Group, which has been dated at 605 + 9 Ma (U –Pb zircon, Thomas et al. 2002). The Tamwirine rhyolites are unconformably overlain by the Ouarzazate Group. Rhyolites and granitoids of the Ouarzazate Group have been dated between 581 + 11 Ma and 543 + 9 Ma (Gasquet et al. 2005). Within the Zenaga inlier, the Sidi El Houssein alkaline


7

Table 1. Major element compositions of Palaeoproterozoic Zenaga plutons Sample

SiO2

TiO2

Azguemerzi granodiorite TZG2 63.80 0.77 ASZ16 65.30 0.54 Tim40 66.00 0.53 Asra12 61.50 0.19 TGZ7 63.10 0.55 TT58 63.30 0.73 TL52 63.70 0.73 TL49 66.00 0.59 TL47 66.50 0.24 Tiz42 67.20 0.53 Ti31 67.90 0.33 TT21 68.00 0.51 TZG6 68.10 0.46 Tiz46 69.00 0.47 Tiz43 69.20 0.42 Ti30 69.70 0.26 Ti34 71.50 0.16 Ait Daoui monzogranite AD24 67.40 0.47 AD26 67.60 0.28 AD28 70.80 0.28 AD25 70.50 0.29 Assersa monzogranite Asra9 72.20 0.04 Asra11 74.70 0.01 As112 74.40 0.07 Asra10 73.40 0.01 As111 73.00 0.04 As114 71.90 0.47 Tamarouft monzogranite AN39 74.20 0.04 TGR43 70.60 0.03 TGR48 72.40 0.02 TAM31 66.30 0.54 TAM32 72.50 0.25 TAM33 71.50 0.02 TAM35 69.50 0.22 TOU37 71.10 0.28 TOU38 72.80 0.04 AN40 73.30 0.14 AN41 71.30 0.12 AN42 68.30 0.12 TGR44 71.20 0.00 TGR45 71.60 0.03 TGR46 71.10 0.02 TGR47 72.90 0.07 TAM44 70.10 0.32 Tazenakht monzogranite TZK3 73.00 0.07 TA4 73.10 0.15 TA8 75.10 0.02 Ti38 71.70 0.14 TZi101 72.85 0.10 TZi10B 72.88 0.10 TZi300 72.26 0.10 TZi500 73.02 0.09 TZi800 68.00 0.20 TZi100 73.28 0.12 TL51 70.00 0.12 AK18 72.30 0.24

Al2O3

Fe2O3t

MgO

MnO

CaO

Na2O

K2O

P2O5

LOS

SUM

15.90 17.60 16.50 15.60 19.00 17.50 17.10 17.00 17.90 15.90 16.50 15.90 15.70 14.70 15.50 16.90 14.40

7.00 3.80 4.17 7.03 6.10 5.65 5.22 3.99 2.73 4.18 2.56 4.93 3.92 4.51 3.44 2.24 1.45

2.00 1.18 1.54 2.11 1.54 2.39 2.04 1.80 1.44 1.21 0.81 1.80 0.90 1.06 1.25 0.49 0.56

0.06 0.03 0.03 0.06 0.03 0.05 0.04 0.04 0.04 0.03 0.03 0.09 0.03 0.02 0.02 0.01 0.02

2.60 2.31 2.57 2.42 0.77 2.40 2.64 2.24 1.16 2.83 1.83 0.25 2.48 2.50 1.34 0.35 1.90

2.94 2.50 2.42 2.99 2.57 2.58 2.38 2.62 5.81 2.81 2.70 2.26 3.26 2.58 2.60 2.51 1.73

3.04 4.64 3.94 2.37 3.42 2.68 4.09 3.56 2.12 3.22 4.79 3.58 3.28 3.96 4.56 6.10 5.99

0.13 0.17 0.17 0.19 0.09 0.24 0.29 0.17 0.10 0.28 0.20 0.09 0.33 0.19 0.15 0.13 0.11

1.27 1.50 1.87 4.77 2.52 2.17 1.29 1.62 1.94 1.45 1.96 2.45 1.25 0.77 1.27 1.15 2.02

99.50 99.57 99.74 99.22 99.69 99.69 99.52 99.63 99.98 99.64 99.61 99.86 99.71 99.76 99.75 99.84 99.84

16.30 16.80 15.40 15.20

3.46 2.59 2.54 2.29

1.62 1.53 0.68 1.38

0.02 0.00 0.02 0.02

0.35 0.81 2.37 0.62

2.34 2.79 2.75 2.50

5.89 5.76 3.76 5.37

0.12 0.15 0.21 0.10

1.69 1.47 1.03 1.52

99.66 99.78 99.83 99.78

16.10 14.20 15.00 15.30 15.10 14.80

1.25 0.91 1.00 4.98 1.08 3.45

0.25 0.11 0.25 0.10 0.30 1.19

0.08 0.01 0.01 0.02 0.01 0.02

0.47 0.66 0.51 0.55 0.69 1.28

3.53 3.28 2.81 4.04 2.80 2.44

5.17 5.46 6.22 4.04 6.15 2.64

0.10 0.18 0.10 0.12 0.17 0.10

0.84 0.44 0.59 0.47 0.59 1.51

100.03 99.96 100.96 103.02 99.93 99.80

14.40 18.10 16.30 15.70 14.20 16.10 20.20 15.50 15.40 17.10 16.00 18.50 16.10 17.40 17.10 15.60 16.90

0.62 1.31 0.33 5.24 2.39 0.96 1.06 2.41 0.57 1.59 1.54 0.53 0.22 0.90 0.73 0.85 1.11

0.15 0.48 0.21 2.23 1.18 0.27 0.23 0.44 0.18 0.55 0.43 0.51 0.07 0.22 0.17 0.20 0.42

0.01 0.01 0.01 0.03 0.01 0.01 0.01 93.00 0.01 0.01 0.01 0.01 0.00 0.01 0.01 0.00 0.01

0.35 0.47 0.42 1.37 1.53 0.75 0.34 0.53 0.50 0.13 0.34 0.47 0.20 0.74 0.40 0.59 0.85

2.94 3.50 3.81 4.67 5.59 2.94 1.07 3.46 3.71 0.88 3.12 2.23 2.53 5.00 4.42 3.84 2.87

6.32 3.68 4.79 2.47 1.26 5.58 5.20 4.91 5.45 4.80 5.65 5.62 8.87 2.63 4.21 4.75 5.70

0.17 0.32 0.33 0.20 0.20 0.30 0.30 0.11 0.20 0.09 0.20 0.53 0.31 0.48 0.31 0.22 0.49

0.55 1.44 0.89 1.11 0.75 1.42 1.86 1.05 0.89 2.12 1.11 1.72 0.36 0.97 0.87 0.86 1.25

99.75 99.93 99.51 99.86 99.87 99.84 99.98 192.79 99.75 100.70 99.81 98.53 99.86 99.97 99.33 99.89 100.01

14.80 14.70 15.50 16.20 15.44 15.38 15.79 15.29 19.45 15.27 17.00 14.90

0.85 1.55 0.70 1.06 1.11 1.23 1.16 1.13 1.36 1.10 0.80 1.94

0.25 0.07 0.19 0.28 0.35 0.34 0.37 0.34 0.61 0.33 0.27 0.42

0.01 0.00 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.00 0.01

0.40 0.45 0.51 0.53 0.38 0.41 0.30 0.26 0.33 0.32 0.43 0.66

2.00 3.88 5.22 2.91 3.73 3.83 3.20 3.01 3.35 3.49 2.76 3.45

7.53 5.34 1.72 6.12 4.80 4.69 5.53 5.50 4.70 4.82 7.08 4.69

0.17 0.03 0.27 0.20 0.16 0.19 0.16 0.16 0.12 0.18 0.14 0.12

0.75 0.51 0.83 0.80 1.07 0.96 1.13 1.20 1.87 1.07 1.29 0.99

99.84 99.77 100.07 99.95 100.00 100.02 100.01 100.01 100.00 99.99 99.89 99.72

8


Table 2. Trace element compositions of Palaeoproterozoic Zenaga plutons Sample

V

Rb

Azguemerzi granodiorite TZG2 44.2 98.8 ASZ16 21.4 89.0 Tim40 25.1 92.6 Ait Daoui monzogranite AD26 26.6 153.9 AD28 10.0 130.3 AD25 13.4 87.8 Assersa monzogranite Asra11 0.76 152.3 As112 0.44 121.5 Asra10 1.07 178.3 Tamarouft monzogranite TGR43 3.30 135.6 TGR48 ,0.1 166.2 TAM31 ,0.1 131.4 Tazenakht monzogranite TA4 1.21 157.1 TA8 3.86 135.3 Ti38 0.93 67.7

Y

Zr

Nb

Ba

La

Ce

Pr

Nd

Sm

Eu

32.8 15.0 11.8

228.4 196.8 165.8

8.3 7.1 7.2

958 1539 812

50.0 47.4 35.0

105.6 96.2 74.1

12.24 11.04 8.76

47.0 42.2 33.1

8.44 6.56 6.14

1.45 1.33 1.03

15.1 15.7 14.7

167.8 122.3 219.6

8.9 4.7 4.1

1628 894 773

8.20 1.84 41.8

20.1 5.11 87.2

2.38 1.04 10.30

9.73 6.64 39.4

2.53 3.42 6.89

0.78 0.55 0.98

6.6 12.5 10.3

37.0 26.9 45.8

1.7 0.8 4.3

51 79 65

5.13 7.00 4.07

12.74 14.38 9.46

1.60 1.78 1.22

6.31 6.90 4.70

2.19 2.31 1.62

0.07 0.06 0.13

2.0 5.6 1.9

4.0 35.2 27.7

0.7 5.9 2.4

891 89 141

2.27 1.32 1.97

4.32 2.70 4.11

0.51 0.42 0.51

1.97 1.92 1.79

0.28 0.90 0.57

0.71 0.10 0.11

26.1 39.4 4.8

42.9 198.5 16.5

1.4 9.6 2.9

841 867 27

19.16 36.0 4.98

39.2 76.5 8.68

4.76 9.00 0.95

17.7 35.2 3.24

3.97 7.13 0.81

0.75 0.74 0.26

granitic ring-complex (579 + 7 Ma; U –Pb zircon; Thomas et al. 2002) is contemporaneous with the Ouarzazate Group. Structurally, the c. 2 Ga Eburnian deformation was high-grade and north– south to NE –SW orientated (Ennih et al. 2001). The Pan-African deformation occurred under greenschist conditions and is mostly NW –SE oriented along the AAMF corridor (Ennih et al. 2001). The Anti-Atlas was later deformed during the Late Palaeozoic Variscan orogeny, which was responsible for the generation of domes and for major dećollements between the Palaeoproterozoic basement and the Neoproterozoic/Phanerozoic cover. Variscan deformation produced spectacular disharmonic folds in the Ouarzazate and Tata Groups and listric extensional faults within the basement (Faik et al. 2001; Burkhard et al. 2006), reactivating faults generated at the end of the Pan-African orogeny (Soulaimani et al. 2004). Those extensional structures were again reactivated during the Cenozoic Alpine orogeny generating the current relief of the Anti-Atlas with its Precambrian inliers and Cenozoic volcanism (Berger et al.). Within the Palaeoproterozoic basement, attributing a structure to the Pan-African, Variscan or to the Alpine events is not easy because of the strong rheological contrast between the rigid basement and the softer sedimentary cover, inducing partition of the deformation. Even the rare thrust faults in the Zenaga basement that are generally interpreted as Pan-African in age could be Variscan or even Alpine in age (Thomas et al. 2002). However, it seems that the

Variscan and Alpine events never induced thermal effects above 300 8C, based on the 580–525 Ma biotite-whole-rock mineral Rb– Sr dates obtained (Thomas et al. 2002). The deformed Eburnian Zenaga granitoids are strongly peraluminous in character.

Analytical techniques Whole-rock major and trace elements. Major elements have been measured by X-ray fluorescence (Universite´ Catholique de Louvain) and the trace elements by ICP-MS (VG PQ2þ, Royal Museum for Central Africa). For trace elements, the result of the alkaline fusion (0.3 g of sample þ 0.9 g of lithium metaborate at 1000 8C during one hour) has been dissolved in 5% HNO3. The calibrations were set using both synthetic solution (mixture of the considered elements at 2, 5 and 10 ppb) and international rock standards (BHVO-1, W1, GA, ACE). For all these elements, the precision varies from 5 to 10% (for details, see Navez 1995). Results are given in Table 1 (major elements) and Table 2 (trace elements). Sr– Nd isotopes. After acid dissolution of the sample in a beaker or in a pressure vessel if any solid is present after centrifugation and Sr and Nd separation on ion-exchange resin, Sr isotopic compositions have been measured on Ta simple filament (VG Sector 54), Nd isotopic compositions on triple Ta –Re– Ta filament (VG Sector 54) in the Section of Isotope Geology of the Africa Museum,


9

Gd

Dy

Ho

Er

Yb

Lu

Hf

Ta

W

Pb

Th

U

6.45 4.63 4.48

5.73 3.10 2.59

1.31 0.64 0.47

3.54 1.41 1.05

3.64 1.32 0.94

0.55 0.21 0.14

6.85 6.04 5.16

0.28 0.26 0.35

0.77 0.70 1.15

20.3 18.2 16.2

14.1 11.0 11.1

1.48 0.84 1.03

3.19 3.97 4.65

3.16 3.11 3.04

0.67 0.63 0.59

1.76 1.60 1.33

1.65 1.36 1.15

0.23 0.19 0.17

5.22 3.94 7.01

0.31 0.20

0.57 0.37 0.57

21.7 22.6 20.7

21.1 10.0 11.3

1.74 1.35 1.29

2.28 2.97 1.92

1.73 2.93 2.03

0.25 0.37 0.36

0.42 0.49 0.71

0.26 0.16 0.52

0.04 0.02 0.06

1.90 1.54 2.17

0.06 0.03 0.04

1.04 1.13 0.37

17.1 19.1 24.8

3.04 2.52 2.70

1.58 1.57 1.19

0.62 1.15 0.71

0.73 1.38 0.89

0.15 0.24 0.17

0.41 0.64 0.41

0.37 0.96 0.48

0.06 0.14 0.08

0.14 1.60 1.29

0.10 0.80 0.23

0.85 2.30 1.42

26.0 1.7 9.2

0.85 0.88 0.82

0.49 1.20 1.97

4.24 7.28 0.92

4.47 7.06 1.23

0.85 1.54 0.26

2.00 4.17 0.76

1.55 4.20 0.95

0.22 0.62 0.15

1.57 7.02 0.64

0.15 0.75 0.51

0.57 1.16 1.27

32.2 11.3 2.8

5.09 13.9 0.44

2.67 2.40 0.43

Tervuren. Repeated measurements of Sr and Nd standards have shown that between-run error is better than 0.000015 (2s). The NBS987 standard has given a value for 87Sr/86Sr of 0.710275 + 0.000006 (2s on the mean of 12 standards, normalized to 86Sr/88Sr ¼ 0.1194) and the Rennes Nd standard a value for 143Nd/144Nd of 0.511959 + 0.000006 (2s on the mean of 24 standards, normalized to 146Nd/144Nd ¼ 0.7219) during the course of this study. All measured ratios have been normalized to the recommended values of 0.710250 for NBS987 and 0.511963 for Nd Rennes standard (corresponding to a La Jolla value of 0.511858) based on the 4 standards measured on each turret together with 16 samples. Decay constant for 87Rb (1.42 10211 a21) was taken from Steiger & Ja¨ger (1977) and for 147Sm (6.54 10212 a21) from Lugmair & Marti (1978). Results are given in Table 3.

Geology and petrography The Azguemerzi porphyritic biotite-granite is a coarse-grained, zoned and mesocratic rock. The primary minerals are mainly plagioclase and megaalkali feldspar which are resorbed and largely sericitized. The mega-alkali feldspar occurs as coarse perthitic orthoclase megacrysts and medium grained microcline. Myrmekitic intergrowths are uncommon, but when present, pervasive and large orthoclase megacrysts are occasionally surrounded by a composite mantle of plagioclase and quartz. The quartz is typically

interstitial and forms late stage inclusions. The main mafic magmatic silicates are biotite, garnet and the epidote, which is included in or associated with biotite (most of the epidote is secondary). Accessory minerals include anhedral, rounded zoned zircon, euhedral apatite, titanite, and ilmenite. Secondary minerals are sericite, chlorite and most of the epidote. The Azguemerzi pluton shows quartzo-feldspathic layers separated by biotite layers, locally associated with gneisses and anatectites in the Assersa and in Tizi-n-Taguergoust valleys. The Azguemerzi pluton displays a magmatic fabric which evolves locally to a true foliation, which probably explains why it was sometimes regarded as porphyritic gneiss. It contains xenolithic micaschists and gneisses of metasedimentary nature which could represent the source of these rocks but are most probably xenoliths from the country-rocks. The Assersa, Aı¨t Daoui and Tamarouft are granodiorite and monzogranite plutons that show the same mineralogical assemblage as the Azguemerzi pluton without biotite, conferring their leucocratic character. The Tazenakht granite in the northern part of Zenaga is a heterogeneous coarse-grained rock. It consists of abundant euhedral alkali feldspar phenocrysts, xenomorphic crystals of quartz, subhedral sericitized polycrystalline plagioclase, twisted biotite and sometimes twisted muscovite. Decimetre-size pegmatitic pockets are associated with acidic pegmatitic and aplopegmatitic dykes. Accessory minerals are mainly oxides; rare corundum has been observed. This granite was generally

10


Table 3. Sr and Nd isotopes of Palaeoproterozoic Zenaga plutons Pluton

Sample

Rb

Sr

Azguemerzi Azguemerzi Azguemerzi Assersa Assersa Assersa Aı¨t Daoui Aı¨t Daoui Aı¨t Daoui Tamarouft Tamarouft Tamarouft Tazenakht Tazenakht Tazenakht

TZG2 ASZ16 Tim40 Asra9 Asra11 As112 AD 24 AD 26 AD 28 An39 TGR43 TGR48 TZK3 TA4 TA8

110 92 95 165 140 178 162 150 90 143 193 155 159 135 68

172 185 140 16 17 21 116 156 138 170 14 29 135 34 63

87

Rb/86Sr

87

Sr/86Sr

1.858 1.444 1.975 32.291 24.876 26.443 4.074 2.798 1.895 2.449 42.804 15.974 3.437 11.576 3.155

0.749223 0.741828 0.766482 1.539725 1.152029 1.499209 0.790233 0.764744 0.749736 0.768874 1.447583 1.039992 0.792343 0.783725 0.809272

2s

Sri 600 Ma

Sri 2035 Ma

Sm

0.000012 0.000009 0.000014 0.000015 0.000015 0.000015 0.000008 0.000009 0.000007 0.000008 0.000017 0.000012 0.000011 0.000011 0.000008

0.733323 0.729473 0.749583 1.263430 0.939183 1.272951 0.755372 0.740802 0.733521 0.747920 1.081334 0.903314 0.762937 0.684674 0.782279

0.694741 0.699494 0.708577 0.592997 0.422712 0.723935 0.670782 0.682706 0.694175 0.697076 0.192629 0.571665 0.691584 0.444326 0.716779

8.44 6.56 6.14 2.19 2.31 1.62 2.53 3.42 6.89 0.28 0.90 0.57 3.97 7.13 0.81

TDM model ages calculated following Nelson & DePaolo (1985). For altered samples, the magmatic eNd and the TDM model ages have been calculated, using the 147Sm/144Nd ratio from 600 to 2035 Ma, of an unaltered sample, from the same pluton (column ‘147Sm/144Nd used’).

deformed in a solid state; it has a planar structure formed with mega-alkali feldspar sometimes fractured, twisted and kinked muscovite and biotite with heterogeneous levels corresponding to mylonitic rocks. Towards the south, the mega-alkali feldspars are deformed within a very intense foliation near the contact of the Azguemerzi granite. These characters reflect an intense and heterogeneous deformation, locally transforming the Tazenakht granite into orthogneiss, porphyroblastic mylonites and phyllonitic layers. A main characteristic of the Zenaga granitoids is the abundant presence of peraluminous minerals with abundant muscovite and almandine-rich garnet (Alm71-89 Pyr3-14 Sps2-12), except the Tazenakht granite which does not bear garnet but is particularly rich in muscovite and locally contains corundum.

melts compared with the ASI (Patinõ-Douce 1992) shows that the Zenaga granitoids are chemically comparable to other garnet bearing granitoids (Fig. 4c) pointing to a heterogeneous peraluminous source rather than to alkali mobility. This high peraluminosity indicates the presence in the source of a strongly peraluminous component, i.e. a pelitic continental crust, with, considering the quite high Ca, Na, Mg and Fe concentrations in these granitoids, the addition of a basaltic component. The garnet can crystallize from the melt itself (Dahlquist et al. 2007) or from the incongruent melting of a muscovite þ biotite þ quartz assemblage in the pelitic source that gives melt and garnet at c. 650 8C. In all cases, a muscovite-rich source is needed, the only mineral that undergoes substantial dehydration at temperature ,800 8C (Miller et al. 2003).

Major elements

REE and Nd isotopes as markers of the Pan-African metacratonic evolution of the Zenaga Eburnian basement

The studied Palaeoproterozoic granitoids (location in Fig. 3b) are all mainly felsic (SiO2 . 68%) except the Azguemerzi granodiorite, which is intermediate in composition (most samples are in the range 61 –70% SiO2). They have variable compositions in alkalis and straddle the boundaries defined for the alkalic, alkali-calcic, calc-alkalic and calcic series (Fig. 4a) suggesting a heterogeneous source or some alkali mobility. They are always strongly peraluminous (Fig. 4b) and yield an alumina saturation index (ASI) decreasing with silica, which points to a peraluminous melt crystallizing aluminous minerals. The Al2O3 activity in the

The rare earth element (REE) patterns of the five plutons studied display very varied shapes considering that all rocks are within the 64–75% SiO2 range. The Azguemerzi granodiorites display normal REE patterns for granodiorite, except the large variability in HREE abundance that can be attributed to variable garnet control (Fig. 5a). The Tazenakht granite (Fig. 5b) has one sample (TA8) very low in REE and with no Eu negative anomaly, which is very different from the two


Nd 47.00 42.22 33.14 6.31 6.90 4.70 9.73 6.64 39.40 1.97 1.92 1.79 17.70 35.24 3.24

147

Sm/144Nd 0.1085 0.0939 0.1121 0.2104 0.2026 0.2084 0.1573 0.3117 0.1058 0.0843 0.2833 0.1911 0.1357 0.1223 0.1520

143

Nd/144Nd

0.511472 0.511356 0.511546 0.512862 0.512699 0.512539 0.511747 0.512344 0.511510 0.511967 0.512959 0.512524 0.511978 0.511972 0.511963

2s

1Nd 600 Ma

1Nd 2035 Ma

0.000009 0.000006 0.000007 0.000008 0.000007 0.000008 0.000008 0.000007 0.000011 0.000032 0.000009 0.000011 0.000014 0.000007 0.000010

216.01 217.15 214.84 3.32 0.74 22.84 214.38 214.58 215.05 24.47 20.38 21.80 28.21 27.30 29.76

0.30 1.86 0.81 0.79 20.36 25.03 27.12 235.99 1.77 16.37 216.45 20.77 3.07 6.49 21.50

other samples; the Eu abundance of sample TA4 is similar to that of sample TZK3 but its other REE are much higher, suggesting a possible beginning of tetrad effect (enrichment in REE due to F-rich fluids; Bau 1996; Veksler et al. 2005). The three samples from the Tamarouft pluton (Fig. 5c) display low abundance of REE, low REE fractionation and variable Eu anomaly. Sample TGR48 displays an REE pattern that could be magmatic; the two other samples show unusual spectra: sample TGR43 is richer in normalized HREE than in LREE and sample AN39 has a strong positive Eu anomaly and also higher normalized HREE than LREE. The three samples from the Ait Daoui pluton (Fig. 5d) have similar HREE but very dissimilar LREE. Sample AD28 has a normal magmatic pattern whereas sample AD26 suffered a strong loss in LREE and sample AD24 has nearly no Eu anomaly and a weak LREE/HREE fractionation. The three samples from the Assersa granite (Fig. 5e) display from La to Eu spectra similar to seagull patterns but a strong depletion in HREE. This suggests the influence of F-rich fluids (for the tetrad effect) coupled to the destabilization of a HREE-rich mineral such as garnet, present in this granite. All these REE patterns indicate that some Zenaga granitoids possess normal REE patterns whereas others do not; the latter suggest an important role for fluids, probably enriched in elements such as F (Veksler et al. 2005). Such fluids are not typical in strongly peraluminous magmas, except for extremely differentiated melts, which is not the case in Zenaga. They are much more similar to the characteristics of alkaline magmas,

147

Magmatic 1Nd (see text) Sm/144Nd used 2035 Ma

measured measured measured measured measured 0.18 (estimated) 0.1058 ¼ AD28 0.1058 ¼ AD28 0.1058 (measured) 0.1911 ¼ TGR48 0.1911 ¼ TGR48 0.1911 (measured) 0.1357 (measured) 0.1357 ¼ TZK3 0.1357 ¼ TZK3

0.38 1.93 0.89 0.94 20.22 0.38 2.52 2.32 1.84 23.32 0.79 20.63 3.17 4.09 1.62

11

TDM (Ma) 2284 2151 2253

2120 2135 2170

2086 1992 2245

frequently displaying seagull REE spectra (Bau et al. 1996). This leads to the question of the age of the fluid influence: these could be late-magmatic fluids or much younger fluids derived by reactivation of the northern boundary of the WAC. The use of the Nd isotopes can constrain this topic. Initial 143Nd/144Nd values relative to bulk Earth (1Nd) are highly variable in the studied plutons: they vary from – 36 to þ16. This variation is observed within all the plutons studied, although the most extreme values belong to the Tamarouft and Ait Daoui plutons (Table 3). When looking closely at the evolution of the 1Nd through time for the different plutons (Fig. 6), several observations emerge. The Azguemerzi granodiorite (Fig. 6a) displays the expected evolution for a magmatic rock: similar slope (proportional to the 147Sm/144Nd ratio), grouped 1Nd at 2035 Ma (U– Pb zircon crystallization age), and a progressively larger variability of 1Nd while time is elapsing (common 1Nd at 2035 Ma, slight difference in 147Sm/144Nd ratios between samples inducing progressive difference in the produced radiogenic 143Nd). The Tazenakht granite (Fig. 6b) shows the opposite behaviour: 1Nd are distinct at 2035 Ma and become more and more similar with time. The Tamarouft and Ait Daoui plutons (Fig. 6c, 6d) display crossed patterns: 1Nd are very different at 0 Ma and 2035 Ma, having a common value during the Neoproterozoic. Finally, the Assersa pluton (Fig. 6e) shows parallel evolution, the difference in 1Nd of the three samples remaining nearly constant through time. The spectra of the Tamarouft and Ait Daoui plutons are particularly enlightening: their 1Nd

12

N. ENNIH & J.-P. LIEGEOIS 143

(a) 12

Na2O+K2O-CaO Azguemerzi Tazenakht Tamarouft Ait Daoui Assersa

10

8

lic

6

alka

lcic

li-ca

4

alka

alic

-alk

calc

2

Ca

lcic

SiO2

0 60

62

64

66

68

70

72

74

76

78

(b) 2.8 2.6

Alumina saturation index (ASI)

2.4 2.2 2.0 1.8 1.6 1.4

Strongly peraluminous 1.2

Slightly peraluminous 1.0

%SiO2

Metaluminous 0.8 60

62

64

66

68

70

72

74

76

(c) A*

13

garnet + Al-silicates granitoids

11

9

7

opx + garnet & opx + crd granitoids

5

cpx + opx granitoids

ASI 3 0.8

1.0

1.2

1.4

1.6

1.8

2.0

2.2

2.4

2.6

Fig. 4. (a) SiO2 vs. Na2O þ K2O-CaO (MALI index; Frost et al. 2001). (b) Alumina saturation index (ASI) ¼ Al2O3/(Na2O þ K2O þ CaO) in molar proportions vs. SiO2 showing the strong peraluminous character of the Palaeoproterozoic Zenaga granitoids. (c) A* (¼ASI *(Na2O þ K2O)) vs. ASI (Patinõ-Douce 1992); boxes for the three kind of peraluminous granitoids are based to the analyses compiled by Patinõ-Douce (1992).

values are very close to ages corresponding to the Pan-African orogeny. Three-point isochrons can even be calculated: the Ait Daoui gives an age of 612 + 300 Ma (initial 143Nd/144Nd ¼ 0.51110 +0.00040; MSWD ¼ 8.5) and Tamarouft an age of 761 + 300 Ma (initial 143Nd/144Nd ¼ 0.51156 +0.00041; MSWD ¼ 2.4). These ages are imprecise but they strongly suggest that, during the Pan-African, the REE of some studied samples were remobilized, as indicated by both

Nd/144Nd isotopic ratios and REE abundances. Such a remobilization requires aggressive F-rich fluid percolations. The hydrothermal event can be linked to the Ouarzazate Supergroup that crosscuts and covers the Zenaga inlier: this is a huge volcanic episode, alkali-calcic in nature and associated with fluorine and beryl that has formerly been mined. This hypothesis can be tested by using an evolutionary model in two stages, with the measured 147Sm/144Nd ratios of the sample from now to 600 Ma (Pan-African orogeny) and with the magmatic 147Sm/144Nd ratio that can be estimated from unaltered samples from 600 to 2035 Ma (U –Pb on zircon crystallization age). The Variscan and Alpine events are not considered here for two reasons: (1) the above mentioned convergence of 1Nd occurred during the Pan-African and (2) the two Phanerozoic events happened at temperature , 3008C in the Zenaga inlier (Thomas et al. 2002). The unaltered Azguemerzi samples (Fig. 5a) can be taken as reference for the magmatic signature of the Eburnian Zenaga plutons: their 1Nd at 2035 Ma vary between þ0.3 and þ1.9. In the Ait Daoui pluton, the AD28 sample has a magmatic REE pattern (Fig. 5d) and gives a 1Nd of þ1.8, within the Azguemerzi range. This sample can, therefore, be considered as having a REE magmatic signature. Its 147Sm/144Nd ratio can be used from 600 to 2035 Ma for the two other samples of the Ait Daoui pluton: with this two stage evolution, sample AD24 get a 1Nd at 2035 Ma of þ2.5 and sample AD26 a 1Nd of þ2.3 (Fig. 6d), very close to the Azguemerzi range. In a similar way, sample TGR48 from the Tamarouft pluton can be considered as having a magmatic REE signature (1Nd at 2035 Ma ¼ 20.8) and when using its 147 Sm/144Nd ratio from 600 to 2035 Ma (Fig. 6e), one of the other samples get an Azguemerzilike 1Nd values (sample TGR43, 1Nd at 2035 Ma ¼ þ0.8) and the other sample (AN39) get a lower 1Nd of 23.3 but however much more magmatic compatible that its single stage 1Nd of þ16.4 (Fig. 6d). The Tazenakht TZK23 sample, which shows a classical magmatic REE pattern (Fig. 5b), has a 1Nd at 2035 Ma of þ3.1. With the TZK23 147 Sm/144Nd ratio, the two other samples provide 1Nd at 2035 Ma of þ1.6 and þ4.1 (Fig. 6b). Finally, if the samples from the Assersa pluton have lost a part of their HREE content (Fig. 5e), a feature likely to be linked to the fact that the garnet in this pluton is altered to chlorite and epidote, their LREE values appear to be less affected. Their 1Nd at 2035 Ma are þ0.8 (Asra 9), –0.36 (Asra11) and –5.03 (As112). The first two are within the range of the Azguemerzi pluton. Sample As112 has a lower 1Nd but this sample is


(a) 1000

1000

(b)

1000

Chondrite normalized


Azguemerzi

Tamarouft

73–75% SiO2

64–66% SiO2

71–74% SiO2

TZK3 TA4 TA8

100

100

100

10

10

10

1

1

1

0.1

La Ce Pr Nd

Sm Eu Gd

Dy Ho Er

Yb Lu

0.1 La Ce Pr Nd

Sm Eu Gd

(d) 1000

Dy Ho Er

AN39 TGR43 TGR48

Yb Lu

0.1 La Ce Pr Nd

Sm Eu Gd

Dy Ho Er

Yb Lu

(e) 1000 Chondrite normalized


Assersa

Ait Daoui

72–75% SiO2

67–71% SiO2 AD24 AD26 AD28

100

100 Asra9 Asra11 As112

10

10

1

1

0.1

(c) Chondrite normalized

Tazenakht

TZG2 ASZ16 Tim40

13

La Ce Pr Nd

Sm Eu Gd

Dy Ho Er

Yb Lu

0.1 La Ce Pr Nd

Sm Eu Gd

Dy Ho Er

Yb Lu

Fig. 5. Rare Earth Elements normalized to chondrites (Boynton 1984) for the studied Zenaga granitoids, showing the F-rich fluids influence on some of the samples. (a) Azguemerzi granodiorite; (b) Tazenakht granite; (c) Tamarouft granite; (d) Ait Daoui granite; (e) Assersa granite.

REE-poor and with a Sm concentration of 1.4 ppm (against the measured value of 1.6 ppm), this sample would have a 1Nd at 2035 Ma of þ0.38 (Fig. 6e). These results show that, when using magmatic 147 Sm/144Nd ratios deduced from pristine samples for all samples from 600 Ma to 2035 Ma, the obtained 1Nd at 2035 Ma are quite homogeneous, varying from – 0.8 to þ2.5 for most samples, with one sample at –3.3 and two samples at þ3.1 and þ4.1 (Fig. 6f); with the recalculated values, the mean 1Nd at 2035 Ma for the Zenaga plutons is remarkably determined at þ1.1 + 0.9. This indicates a mainly juvenile source (1Nd of the depleted mantle at 2035 Ma ¼ þ5.5), which is also indicated by the TDM model ages of the samples having 147Sm/144Nd , 0.15 (measured or recalculated)

whose mean is 2159 + 61 Ma. The calculation for the other samples would have implied a three-stage evolution and would have given similar model ages. They denote that major REE fractionation existed in the source, which, coupled with the strongly peraluminous character of the Zenaga granitoids, is consistent with a metasedimentary juvenile continental crustal source combining metasedimentary formations and mafic rocks melted at depth. More details about the nature of the Eburnian orogeny in the Anti-Atlas require further constraints on the Zenaga metamorphic basement whose age, according to the inherited zircons dated in the granitoids, would be around 2.17 Ga (Thomas et al. 2002). Sr isotopes have been largely modified by the Pan-African event: at 2035 Ma, most of the 87 Sr/86Sr initial ratios are much lower than 0.7.

14


(a)

(b)

+20

+20

+20

Azguemerzi

Nd

Nd

Tazenakht TZK3

TZG2

TGR43

TA4 0

0

0

Tamarouft

Nd

+10

+10

+10

(c)

TGR48

TA8 –10

–10

–10

–20

–20

An39

TIM40 –20 ASZ1 –30

–30

–30

Ma

Ma

–40 0

400

800

1200

1600

2035

–40 0

(d)

400

800

1200

1600

2035

(e)

+20

+10

0

0

800

1200

1600

2035

+20 Nd

+10

400

(f)

+20

Ait Daoui

Nd

Ma

–40 0

Assersa

All plutons

Nd

+10 ASRA9 ASRA11

+1.1 ±0.9 0

ASRA112

AD26 –10

–10

–10

–20

–20

AD24 –20 AD28 –30

–40 0

–30

–30

Ma 400

800

1200

1600

2035

–40 0

143

Ma 400

800

144

1200

143

1600

144

2035

Ma

–40 0

400

143

800

1200

1600

2035

144

Fig. 6. Nd isotopes (shown as 1Nd ¼ [( Nd/ Nd)sample-( Nd/ Nd)Bulk-Earth]/( Nd/ Nd)Bulk-Earth, both sample and Bulk Earth Nd isotopic compositions being calculated at the time considered. Full lines are the evolution calculated with the measured 147Sm/144Nd ratio of the sample (single stage evolution). The dashed lines from 600 Ma (age of the Pan-African perturbation) to 2035 Ma (age of the granite crystallization; U– Pb zircon, Thomas et al. 2002) are the evolutions calculated with the 147Sm/144Nd ratio of unaltered samples during the Pan-African orogeny, from the same pluton (two-stage evolution). See text for more explanation. (a) Azguemerzi granodiorite; (b) Tazenakht granite; (c) Tamarouft granite; (d) Ait Daoui granite; (e) Assersa granite; (f) All plutons together with measured magmatic 147Sm/144Nd ratios (single stage evolution) or adopted 147Sm/144Nd ratios (two-stage evolution).

The alignment (with a poor MSWD of 190) determined by the 15 samples gives an ‘age’ of 1467 + 310 Ma (initial 87Sr/86Sr ¼ 0.711 + 0.078). This ‘age’ is the result of the interplay of the Eburnian age of the granitoids and the major Pan-African effect.

Conclusions The chemical and isotopic data for the Zenaga Palaeoproterozoic magmatism indicate that these peraluminous granitoids originated from the partial melting of a juvenile, largely metasedimentary, crustal source but that the Pan-African orogeny has strongly reactivated some plutons or parts of plutons, including REE. The likely cause was the circulation of F-rich fluids circulation linked to the extrusion of the voluminous Ouarzazate Group and the emplacement of associated plutons (c. 580 Ma) along reactivated faults and shear zones. The reactivation event occurred during important vertical movements (the thickness of the

Ouarzazate Group varies from 0 to more than 2500 m) but without major crustal or lithospheric thickening. This is demonstrated by the low-grade character of the Pan-African metamorphism (greenschist facies) and by the excellent preservation of the c. 800 Ma passive margin sediments and the early c. 750–700 Ma ophiolitic complex. This is considered to be typical of a metacratonic evolution: the cratonic boundary was dissected by faults and invaded by magmas but preserved most of its rigidity and primary features. Such a metacratonic reactivation was favourable for fluid circulation, in this case able to mobilize rare-earth elements. Such fluid movements were also an excellent vector for element concentration and genesis of mineralizations, for which the Anti-Atlas is internationally renowned. This book concerns an example of these areas of paramount importance that are the boundaries of craton that suffered partial reactivation, i.e. a metacratonic evolution. The West African craton is a particularly good example because: (1) it became a strong craton during the Mesoproterozoic, a


period of 600 Ma that left no trace on the WAC; (2) all its boundaries intervened as indentors during the Pan-African orogeny leading to situations varying from nearly frontal collision to nearly entirely transcurrent dockings; (3) only its western and northern boundaries were included as foreland in the Variscan collision, allowing fruitful comparison between the eastern and the western WAC boundaries; (4) the WAC boundaries were major suppliers of lithospheric pieces or of sedimentary material for the Peri-Gondwanan terranes now located in Europe or in North America; (5) currently, the WAC boundaries are reactivated by the stress generated by the Africa-Europe convergence and constitutes a key area for studying intraplate deformation submitted to stress; (6) WAC boundaries, or at least a part of them, are known for their mined or potential mineral deposits. A better knowledge of the boundaries of the West African craton is a prerequisite for understanding all these processes. This is the aim of the papers constituting this Special Publication. This is a contribution to the IGCP485. We warmly thank the UNESCO and the IUGS for their financial support and encouragements during these five years. We are happy that this Special Publication comes out during the International Year of Planet Earth, an outstanding outcome from these two organizations. Organizing contacts and field meetings in the desert regions of Mauritania, Mali, Algeria and Morocco was always a challenge but also very stimulating. We would like to thank Khalidou Loˆ (Mauritania), Samba Sacko (deceased), Dramane Dembe´le´, Renaud Caby (Mali), Khadidja Ouzegane, Abla Azzouni-Sekkal (Algeria), Ezzoura Errami Hassan Admou, Hassan Ouanaimani, and Abdelilah Fekkak (Morocco) for their personal support and the support of organizations in their country. We are also grateful to the Universities of Algiers (USTHB) and El Jadida, the Office Mauritanien de Recherches Geólogiques, the Direction Nationale de la Geólogie et des Mines de Bamako, the Algerian COMENA as well as the Ministries of Mines and Energy of Mauritania, Mali and Algeria. We warmly thank Bernard Bonin for judicious remarks on key points of the paper and Kevin Hefferan for detailed and helpful remarks and corrections. Both have suggested significant improvements to this article. Finally, we would like to thank the Publication team of the Geological Society for their support and patience.

References A BDELSALAM , M, L IE´ GEOIS , J. P. & S TERN , R. J. 2002. The Saharan metacraton. Journal of African Earth Sciences, 34, 119–136. A BOUCHAMI , W., B OHER , M., M ICHARD , A. & A LBARE` DE , F. 1990. A major 2.1 Ga event of mafic magmatism in West Africa: an early stage of crustal accretion. Journal Geophysical Research, 95, 17605– 17629.

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A FFATON , P., R AHAMAN , M. A., T ROMPETTE , R. & S OUGY , J. 1991. The Dahomeyide Orogen: tectonothermal evolution and relationships with the Volta basin. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 107– 122. A KO , J. A. & W ELLMAN , P. 1985. The margin of the West African craton: the Voltaian basin. Journal of the Geological Society, London, 142, 625–632. A ZZOUNI -S EKKAL , A., L IE´ GEOIS , J. P., B ECHIRI B ENMERZOUG , F., B ELAIDI -Z INET , S. & B ONIN , B. 2003. The ‘Taourirt’ magmatic province, a marker of the very end of the Pan-African orogeny in the Tuareg Shield: review of the available data and Sr-Nd isotope evidence. Journal of African Earth Sciences, 37, 331– 350. B ATES , R. L. & J ACKSON , J. A. (eds) 1980. Glossary of Geology. American Geological Institute, Falls Church, Virginia. B AU , M. 1996. Controls on the fractionation of isovalent trace elements in magmatic and aqueous systems. Evidence from Y/Ho, Zr/Hf and lanthanide tetrad effect. Contributions to Mineralogy and Petrology, 123, 323– 333. B LACK , R. & L IE´ GEOIS , J. P. 1993. Cratons, mobile belts, alkaline rocks and continental lithospheric mantle: the Pan-African testimony. Journal of the Geological Society, London, 150, 89– 98. B LACK , R., C ABY , R., M OUSSINE -P OUCHKINE , A., B ERTRAND , J. M. L., B OULLIER , A. M., F ABRE , J. & L ESQUER , A. 1979. Evidence for Precambrian plate tectonics in West Africa. Nature, 278, 223–227. B LACK , R., L ATOUCHE , L., L IE´ GEOIS , J. P., C ABY , R. & B ERTRAND , J. M. 1994. Pan-African displaced terranes in the Tuareg shield (central Sahara). Geology, 22, 641– 644. B OHER , M., M ICHARD , A., A LBARE` DE , F., R OSSI , M. & M ILE´ SI , J. P. 1992. Crustal growth in West Africa at 2.1 Ga. Journal Geophysical Research, 97, 345–369. B OUOUGRI , E. H. & S AQUAQUE , A. 2004. Lithostratigraphic framework and correlation of the Neoproterozoic northern West African craton passive margin sequence (Siroua, Zenaga, Bouazzer-Elgraara inliers, Central Anti-Atlas, Morocco): an integrated approach. Journal of African Earth Sciences, 39, 227–238. B OYTON , W. V. 1984. Geochemistry of the rare earth elements: meteorite studies. In: H ENDERSON , P. (ed.) Rare Earth Element Geochemistry. Elsevier, 63–114. B RONNER , G., R OUSSEL , J., T ROMPETTE , R. & C LAUER , N. 1980. Genesis and Geodynamic Evolution of the Taoudeni Cratonic Basin (Upper Precambrian and Paleozoic), Western Africa, Dynamics of Plate Interiors. Geodynamics Series vol. 1, American Geophysical Union, 81– 90. M., C ARITG , S., H ELG , U., B URKHARD , R OBERT -C HARRUE , C. & S OULAIMANI , A. 2006. Tectonics of the Anti-Atlas of Morocco. Comptes Rendus Geoscience, 338, 11– 24. C ABY , R., A NDREOPOULOS -R ENAUD , U. & P IN , C. 1989. Late Proterozoic arc-continent and continentcontinent collision in the Pan-African trans-Saharan belt of Mali. Canadian Journal of Earth Sciences, 26, 1136–1146.

16


C ALVEZ , J. Y. & V IDAL , P. 1978. Two billion years old relicts in the Hercynian belt of Western Europe. Contributions to Mineralogy and Petrology, 65, 395–399. C ARITG , S., B URKHARD , M., D UCOMMUN , R., H ELG , U., K OPP , L. & S UE , C. 2004. Fold interference patterns in the Late Palaeozoic Anti-Atlas belt of Morocco. Terra Nova, 16, 27–37. D AHLQUIST , J. A., G ALINDO , C., P ANKHURST , R. J., R APELA , C. W., A LASINO , P. H., S AAVEDRA , J. & F ANNING , C. M. 2007. Magmatic evolution of the Penõń Rosado granite: Petrogenesis of garnet-bearing granitoids. Lithos, 95, 177– 207. D ECKART , K., B ERTRAND , H. & L IE´ GEOIS , J. P. 2005. Geochemistry and Sr, Nd, Pb isotopic composition of the Central Atlantic Magmatic Province (CAMP) in Guyana and Guinea. Lithos, 82, 282–314. D E LA B OISSE , H. 1979. Pe´trologie et geochronologie des roches cristallophylliennes du bassin de Gourma (Mali), conse´quences pe´trogeńe´tiques. Unpublished PhD thesis, Montpellier. D’L EMOS , R. S., I NGLIS , J. D. & S AMSON , S. D. 2006. A newly discovered orogenic event in Morocco: Neoproterozoic ages for supposed Eburnean basement of the Bou Azzer inlier, Anti-Atlas mountains. Precambrian Research, 147, 65–76. E NNIH , N. & L IE´ GEOIS , J. P. 2001. The Moroccan AntiAtlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton. Precambrian Research, 112, 291–304. E NNIH , N., L ADURON , D., G REILING , R. O., E RRAMI , E., DE W ALL , H. & B OUTALEB , M. 2001. Superposition de la tectonique e´burneénne et panafricaine dans les granitoı¨des de la bordure nord du craton ouest africain, boutonnie`re de Zenaga, Anti-Atlas central, Maroc. Journal of African Earth Sciences, 32, 677– 693. F ABRE , J. 2005. Geólogie du Sahara occidental et central. Se´rie/Reeks: Tervuren African Geosciences Collection, MRAC Tervuren, Belgique. F AIK , F., B ELFOUL , M. A., B OUABDELLI , M. & H ASSENFORDER , B. 2001. Les structures de la couverture neóprote´rozoı¨que terminal et paleózoı¨que de la re´gion de Tata, Anti-Atlas centre-occidental, Maroc: de´formation polyphaseé, ou interactions socle/couverture pendant l’orogene`se hercynienne? Journal of African Earth Sciences, 32, 765– 776. F EYBESSE , J. L. & & M ILE´ SI , J. P. 1994. The Archaean/ Proterozoic contact zone in West Africa: a mountain belt of d6collement thrusting and folding on a continental margin related to 2.1 Ga convergence of Archaean cratons? Precambrian Research, 69, 199– 227. F ROST , B. R., B ARNES , C. G., C OLLINS , W. J., A RCULUS , R. J., E LLIS , D. J. & F ROST , C. D. 2001. A geochemical classification for granitic rocks. Journal of Petrology, 42, 2033–2048. G ASQUET , D., L EVRESSE , G., C HEILLETZ , A., A ZIZI S AMIR , M. R. & M OUTTAQI , A. 2005. Contribution to a geodynamic reconstruction of the Anti-Atlas (Morocco) during Pan-African times with the empasis on inversion tectonics and metallogenic activity at the Precambrian-Cambrian transition. Precambrian Research, 140, 157 –182.

G ASQUET , D., E NNIH , N., L IEGEOIS , J. P., S OULAIMANI , A. & M ICHARD , A. 2008. The Pan-African Belt. In: M ICHARD , A., C HALOUAN , A. & S ADDIQI , O. (eds) Continental Evolution: The Geology of Morocco. Structure, Stratigraphy, and Tectonics of the Africa-Atlantic-Mediterranean Triple Junction. Springer Verlag, Lecture Notes in Earth Sciences, 116. G UIRAUD , R., B OSWORTH , W., T HIERRY , J. & D ELPLANQUE , A. 2005. Phanerozoic geological evolution of Northern and Central Africa: an overview. Journal of African Earth Sciences, 43, 83– 143. H EFFERAN , K. P., A DMOU , H., K ARSON , J. A. & S AQUAQUE , A. 2000. Anti-Atlas (Morocco) role in Neoproterozoic Western Gondwana reconstruction. Precambrian Research, 103, 89– 96. I NGLIS , J. D., M AC L EAN , J. S., S AMSON , S. D., D’L EMOS , R. S., A DMOU , H. & H EFFERAN , K. 2004. A precise U –Pb zircon age for the Bleı¨da granodiorite, Anti-Atlas, Morocco: implications for the timing of deformation and terrane assembly in the eastern Anti-Atlas. Journal of African Earth Sciences, 39, 277–283. J AHN , B. M., C ABY , R. & M ONIE´ , P. 2001. The oldest UHP eclogites of the World: age of UHP metamorphism, nature of protoliths and tectonic implications. Chemical Geology, 178, 143 –158. L AVILLE , E., P IQUE´ , A., A MRHAR , M. & C HARROUD , M. 2004. A restatement of the Mesozoic Atlasic rifting (Morocco). Journal of African Earth Sciences, 38, 145–153. L IE´ GEOIS , J. P., C LAESSENS , W., C AMARA , D. & K LERKX , J. 1991. Short-lived Eburnian orogeny in southern Mali. Geology, tectonics, U– Pb and Rb– Sr geochronology. Precambrian Research, 50, 111–136. L IE´ GEOIS , J. P., L ATOUCHE , L., B OUGHRARA , M., N AVEZ , J. & G UIRAUD , M. 2003. The LATEA metacraton (Central Hoggar, Tuareg shield, Algeria): behaviour of an old passive margin during the Pan-African orogeny. Journal of African Earth Sciences, 37, 161–190. L IE´ GEOIS , J. P., B ENHALLOU , A., A ZZOUNI -S EKKAL , A., Y AHIAOUI , R. & B ONIN , B. 2005. The Hoggar swell and volcanism: reactivation of the Precambrian Tuareg shield during Alpine convergence and West African Cenozoic volcanism. In: F OULGER , G. R., N ATLAND , J. H., P RESNALL , D. C. & A NDERSON , D. L. (eds) Plates, Plumes and Paradigms. Geological Society of America, Special Paper, 388, 379–400. L UGMAIR , G. W. & M ARTI , K. 1978. Lunar initial 143 Nd/144Nd: differential evolution of the lunar crust and mantle. Earth and Planetary Science Letters, 39, 349–357. M ALUSA , M. G., P OLINO , R., F ERONI , A. C., E LLERO , A., O TTRIA , G., B AIDDER , L. & M USUMECI , G. 2007. Post-Variscan tectonics in eastern Anti-Atlas (Morocco). Terra Nova, 19, 481– 489. M ARZOLI , A., R ENNE , P. R., P ICCIRILLO , E. M., E RNESTO , M., B ELLIENI , G. & D E M IN , A. 1999. Extensive 200-million-year-old continental flood basalts of the Central Atlantic Magmatic Province, Science, 284, 616– 618.

THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT M ILLER , C. F., M C D OWELL , S. M. & M APES , R. W. 2003. Hot and cold granites? Implications of zircon saturation temperatures and preservation of inheritance. Geology, 31, 529– 532. M OUSSINE -P OUCHKINE , A. & B ERTRAND -S ARFATI , J. 1978. Le Gourma: un aulacoge`ne du Prećambrien supe´rieur? Bulletin de la Socie´te´ Geólogique de France, 20, 851–857. N AVEZ , J. 1995. De´termination d’e´le´ments en traces dans les roches silicateés par ICP-MS. Rapport Annuel du De´partement de Geólogie et Mine´ralogie 1993–1994 Museé Royal de l’Afrique Centrale, Tervuren, Belgique, 139– 147. N E´ DE´ LEC , A., A FFATON , P., F RANCE -L ANORD , C., C HARRIE` RE , A. & A LVARO , J. 2007. Sedimentology and chemostratigraphy of the Bwipe Neoproterozoic cap dolostones (Ghana, Volta Basin): a record of microbial activity in a peritidal environment. Comptes Rendus Geosciences, 339, 223– 239 and erratum Comptes Rendus Geosciences, 339, 516– 518. N ELSON , B. K. & D E P AOLO , D. J. 1985. Rapid production of continental crust 1.7 to 1.9 b.y. ago: Nd isotopic evidence from the basement of the North Americam midcontinent. Geological Society of America Bulletin, 96, 746–754. P ATIN˜ O D OUCE , A. E. 1992. Calculated relationships between activity of alumina and phase assemblages of silica-saturated igneous rocks. Journal of Volcanology and Geothermal Research, 52, 43–63. P ELLETER , E., C HEILLETZ , A., G ASQUET , D. ET AL . 2007. Hydrothermal zircons: a tool for ion microprobe U–Pb dating of gold mineralization (TamlaltMenhouhou gold deposit—Morocco). Chemical Geology, 245, 135– 161. P OTREL , A., P EUCAT , J. J. & F ANNING , C. M. 1998. Archean crustal evolution of the West African craton: example of the Amsaga area (Reguibat Rise). U–Pb and Sm–Nd evidence for crustal growth and recycling. Precambrian Research, 90, 107–117. S AMSON , S. D., I NGLIS , J. D., D’L EMOS , R. S., A DMOU , H., B LICHERT -T OFT , J. & H EFFERAN , K. 2004.

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Geochronological, geochemical, and Nd-Hf isotopic constraints on the origin of Neoproterozoic plagiogranites in the Tasriwine ophiolite, Anti-Atlas orogen, Morocco. Precambrian Research, 135, 133– 147. S AMSON , S. D. & D’L EMOS , R. S. 1998. U –Pb geochronology and Sm–Nd isotopic composition of Proterozoic gneisses, Channel Islands, UK. Journal of the Geological Society, London, 155, 609– 618. S OULAIMANI , A, E SSAIFI , A, Y OUBI , N. & H AFID , A. 2004. Les marqueurs structuraux et magmatiques de l’extension crustale au Prote´rozoı¨que terminal– Cambrien basal autour du massif de Kerdous (AntiAtlas occidental, Maroc). Comptes Rendu Geoscience, 336, 1433–1441. S TEIGER , R. H. & J A¨ GER , E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Sciences Letters, 36, 359–362. T HOMAS , R. J., C HEVALLIER , L. P., G RESSE , P. G. ET AL . 2002. Precambrian evolution of the Sirwa Window, Anti-Atlas Orogen, Morocco. Precambrian Research, 118, 1–57. T HOMAS , R. J., F EKKAK , A., E NNIH , N. ET AL . 2004. A new lithostratigraphic framework for the Anti-Atlas Orogen, Morocco. Journal of African Earth Sciences, 39, 217– 226. V ERATI , C., B ERTRAND , H. & F ERAUD , G. 2005. The farthest record of the Central Atlantic Magmatic Province into West Africa craton: Precise 40Ar/39Ar dating and geochemistry of Taoudenni basin intrusives (northern Mali). Earth and Planetary Science Letters, 235, 391–407. V EKSLER , I. V., D ORFMAN , A. M., K AMENETSKY , M., D ULSKI , P. & D INGWELL , D. B. 2005. Partitioning of lanthanides and Y between immiscible silicate and fluoride melts, fluorite and cryolite and the origin of the lanthanide tetrad effect in igneous rocks. Geochimica et Cosmochimica Acta, 69, 2847– 2860.

REE patterns, Nd – Sm and U – Pb ages of the metamorphic rocks of the Diagorou–Darbani greenstone belt (Liptako, SW Niger): implication for Birimian (Palaeoproterozoic) crustal genesis A. SOUMAILA1, P. HENRY2, Z. GARBA1 & M. ROSSI2 1

De´partement de Geólogie, Faculte´ des Sciences, Universite´ Abdou Moumouni, BP: 10662, Niamey, Niger (e-mail: [email protected] or [email protected]) 2

De´partement de Geósciences, UFR Sciences et Techniques, Universite´ de Franche-Comte´, 16, route de Gray, 25030 Besançon cedex, France Abstract: The Palaeoproterozoic Diagorou –Darbani greenstone belt in Liptako (Niger) is made up of micaschists, various amphibolites, metaconglomerates, and metabasalts intruded by granodioritic plutons. One of these plutons, the Dargol granodiorite, is dated at 2174 + 4 Ma, this age is comparable with those previously reported by many researchers. The micaschists (Type 1 sediments) and intercalated amphibolites have REE patterns variously enriched in light REE (LREE), suggesting oceanic arc-related rocks. The protolith of these micaschists have calcalkaline affinities, with crystallization ages around 2273– 2278 Ma, and TDM close to 2.3 Ga. This age is suggestive of an early Palaeoproterozoic magmatic event in crustal growth. The metaconglomerates (Type 2 sediments) exhibit REE patterns depleted in heavy REE (HREE) typical of tonalite–trondhjemite–granodiorite (TTG), the protolith of which may have been crystallized at 2187 + 55 Ma. These results, together with the earlier ones, led to a Palaeoproterozoic geodynamic model in which the crustal genesis was completely related to subduction zones, with an early Palaeoproterozoic magmatic event. Partial melting of a mantle slab generated the granitoid rocks of calc-alkaline affinities, whereas those with TTG characters could have been produced by direct partial melting of subducted oceanic crust. The crustal growth may have been the result of a continuous input of crustal materials in the interval time of 2.3–2.15 Ga, corresponding to ages recorded by various detrital zircon grains of micaschists and conglomerates.

The West African craton is bounded by Pan-African and Hercynian mobile belts, and comprises a northern rise (Reguibat rise) and a southern one (Man rise), each of which has a western Archaean domain and an eastern Palaeoproterozoic area (Fig. 1a). The major part of the Man rise is made up of an alternating granitoid batholiths and greenstone belts (Fig. 1b). From recent geochronological studies, it was argued that the Birimian terranes were emplaced in a time interval of 2.2–2.1 Ga (e.g. Boher et al. 1992; Hirdes et al. 1992). The Sm–Nd isotopic data led some workers to consider the Birimian (Palaeoproterozoic) as a major period of creation and growth of juvenile continental crust (Abouchami et al. 1990; Boher et al. 1992). The granitoid batholiths, emplaced from 2.18 to 2.15 Ga (Hirdes et al. 1992; Cheilletz et al. 1994), are slightly younger than the greenstone belts dated between 2.2 and 2.1 Ga (Abouchami et al. 1990; Boher et al. 1992; Hirdes et al. 1992; Hirdes & Davis 1998). These batholiths are interpreted to have been derived from partial melting of garnet-amphibolites (e.g. Pouclet et al. 1990), whereas the volcano-sedimentary and sedimentary

rocks constituting part of the greenstone belts are accepted to be products of erosion of the volcanic and granitoid rocks (e.g. Davis et al. 1994; Bossie`re et al. 1996). The aim of this paper is to present some U –Pb ages from granitoid (zircons of Dargol pluton) and detrital zircons from micaschists and metconglomerates. The REE and Sm–Nd isotopic characteristics of all samples (Soumaila 2000) will be considered to constrain the mechanism of Birimian crust genesis.

Geological context The Liptako, the northeasternmost part of the Man rise (Fig. 1b), is composed of alternating granitoid batholiths and greenstone belts; the Diagorou – Darbani belt is the central one of the area (Fig. 1b). This belt is made up of metamorphic low-grade rocks (various detrital metsediments, talcschists, chloritoschists, metabasalts with pillow structures), and medium-grade rocks (amphibolites,

From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 19–32. DOI: 10.1144/SP297.2 0305-8719/08/$15.00 # The Geological Society of London 2008.

20

A. SOUMAILA ET AL.

Fig. 1. (a) Main geological entities of the West African craton. 1, Archaean rocks; 2, Birimian rocks; 3, Precambrian sedimentary basins; 4, Pan-African chain; 5, Hercynian chains; 6, reactivated basement; 7, Phanerozoic sedimentary basins. (b) Geological map of Man rise (Mile´si et al. 1989, modified) and location of Liptako. DD, Diagorou –Darbani greenstone belt; 1, Phanerozoic cover; 2, Birimian sedimentary and volcano-sedimentary rocks; 3, Birimian greenstone belts; 4, Granitoids; 5, Dextral faults; 6, Sinistral faults.

kyanite–staurolite-bearing micaschists). Locally mafic to ultramafic rocks, dioritic to granitic late to post-kinematic plutons intruded the belt. In some places, volcanic and plutonic rocks of intermediate composition crop out (Fig. 2). In the southwestern part of the region, the Taka pluton, dated at 2137 Ma (Klockner 1991), divides the belt into two branches: an eastern branch, dominantly composed of sedimentary to volcano-sedimentary rocks, and the western branch, mainly of magmatic rocks. The belt is flanked by an eastern pluton (Dargol pluton), and a northern to western one (Te´ra pluton) dated at 2158 Ma (Cheilletz et al. 1994). Field relations, and metamorphic and deformation patterns led Soumaila (2000) to propose an Archaean-type tectonic event, as evoked by Vidal et al. (1996) in Ivory Coast. According to Dupuis et al. (1991) and Pons et al. (1995), the Diagorou– Darbani greenstone belt has recorded a single deformation event, which was related to the interference between emplacement and growth of huge plutons, and a regional NW –SE shortening. In more recent studies, this event has been interpreted as a continuum of deformation that generated a NE– SW-trending schistosity and foliation

(Ama-Salah et al. 1996; Soumaila & Konate´ 2005). The Liptako Birimian rocks are emplaced in an oceanic arc environment (Ama-Salah et al. 1996), or in an oceanic arc and back-arc basin setting, with a strong influence of a metasomatized depleted mantle in the genesis of the metabasalts and amphibolites (Soumaila et al. 2004).

Petrographic features of samples The locations of the samples considered in the present paper are given in Figure 2. Petrographic types are: amphibolites (samples 255, 262, 268, 274, 284), kyanite–staurolite-bearing micaschists (samples 129, 135, 6880), metabasalts (samples 216, 232, 414, 420, 348, 514, 804, 810, 812, 821), granodiorite (sample 6799), and metaconglomerates (samples 868, 784). The petrographic features of these samples are given in Table 1.

Rare earth elements In a previous paper (Soumaila et al. 2004), the geochemical features, including REE patterns

REE, Nd, U– Pb, AND BIRIMIAN CRUSTAL GENESIS

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Fig. 2. Geological map of the Diagorou– Darbani greenstone belt (Soumaila 2000, modified). 1, Granitoid plutons; 2, late to post-tectonic granites; 3, tonalites and quartz-diorites; 4, mafic to ultramafic plutons; 5, metabasalts; 6, talcschists and chloritoschists; 7, low-grade metamorphic sediments and volcano-sediments; 8, volcanic rocks of intermediate to acid composition; 9, amphibolites locally interbedded with micaschists; 10, syenite; 11, supposed fault; 12, major shear zone; 13, sample location.

of garnet-free amphibolites (samples 255, 262), the metabasalt samples (216, 232, 514, 414, 420, 348, 804, 812, 821), and others (amphibolites and metabasalts not mentioned here) have also been discussed. The bulk-rock REE contents are given in Table 2, and the chondrite-normalized REE patterns shown in Figure 3. Two types of metasedimentary rocks are distinguished: Type 1 consisting of

garnet-amphibolites (samples 268, 274 and 287; probably metagreywackes) and micaschists (samples 129, 135, 6836), and Type 2 made up of metamorphosed conglomerates (samples 784 and 868; analyses were performed on the monogenetic matrix). In both groups, the REE patterns are strongly fractionated with values of Lan/Ybn (normalized to chondrite) ranging from 4.58 to 14.39, and also a strong to moderate plagioclase

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Table 1. General features of samples Sample no.

Lithology Dark garnet-amphibolites

268

Dark to greenish garnet-pyroxene amphibolite

129 (6836) 135 (6880)

Grey to brownish kyanite– staurolite-bearing micaschists

868, 784

Conglomerates, with locally dm-sized oblong pebbles of quartz, and of tourmaline-bearing quartzite Grey granodiorite

6799 216, 232, 414, 420, 348, 514, 804, 810, 812, 821 255, 262

Hornblende, plagioclase, garnet, quartz, biotite, chlorite, epidote, titanite, opaque minerals Hornblende, plagioclase, quartz, pyroxene, garnet, epidote, calcite, titanite, apatite, opaque minerals Plagioclase, quartz, biotite, muscovite, staurolite, kyanite, cordierite, + garnet, chlorite, opaque mineral, scarce zircon grains Quartz, sericite, muscovite, opaque minerals

Metamorphic grade

Structural features

Amphibolite facies with N40- to N163-trending foliation superimposed greenschist facies N145-trending foliation Amphibolite facies with superimposed greenschist facies Amphibolite facies with N35 foliation superimposed greenschist facies

Greenschist facies

N50-trending schistosity

Quartz, plagioclase, hornblende, N65-trending foliation biotite, epidote, allanite, apatite, opaque minerals Metabasalts of Te´ra and Tillabe´ry, Quartz, plagioclase, hornblende, Epidote-amphibolite to greenschist Foliation trending N70 – N80 in blackish to greenish with actinote, epidote, chlorite, facies Te´ra basalts and N1750– N145 locally flattened pillow titanite, opaque minerals in Tillabe´ry structures Grey to blackish and locally Quartz, plagioclase, hornblende, Amphibolite facies with N45-trending regular foliation greenish amphibolites, fine to biotite with zircon grains, superimposed greenschist facies with decimetre-scale folds medium grained titanite, apatite, opaque minerals

A. SOUMAILA ET AL.

274, 284

Mineralogical composition


23

Table 2. REE contents of samples Samples:

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

Garnetamphibolites

Garnet-free amphibolites

Micaschists Type 1 sediments

Metaconglomerates Type 2 sediments

255

262

274

287

268

129

135

784

868

127.85 276.46 34.47 147.89 22.61 5.18 13.36 1.34 5.06 0.53 1.52 0.13 0.89 0.13

170.26 355.68 44.03 175.93 24.63 5.86 14.76 1.56 5.87 0.68 1.92 0.20 1.14 0.16

35.31 30.12 23.92 17.53 10.69 9.29 7.90 5.44 5.25 4.83 4.82 5.48 4.93 5.99

49.01 42.38 35.50 28.34 18.59 12.77 11.02 7.56 6.40 5.33 5.28 5.63 5.11 6.25

27.10 23.21 19.22 15.36 10.71 9.62 7.78 5.51 5.03 4.41 4.23 4.83 3.87 4.59

58.71 43.12 35.84 27.32 16.13 12.54 10.88 7.89 7.33 5.65 4.71 4.34 4.08 3.98

32.73 35.81 33.27 25.98 16.50 9.24 10.82 9.15 7.97 7.12 6.35 6.92 7.15 6.60

40.33 24.76 17.15 9.25 3.16 1.95 1.54 2.07 1.97 1.53 1.61 1.67 1.33 1.84

38.96 30.62 19.49 11.91 6.28 4.83 2.22 2.59 2.47 2.12 2.37 3.06 3.02 3.42

fractionation in the source of these rocks. This is suggested by a negative Eu anomaly varying from 0.68 to 0.86. A concave shape of heavy REE (HREE) patterns characterizes the samples of Type 2, which is a typical feature of the Archean TTG

Fig. 3. REE patterns of rocks from the Diagorou– Darbani greenstone belt, compared with those of granitoid (Ama-Salah et al. 1996). Data sources: 1, Soumaila et al. (2004); 2, Ama-Salah et al. (1996); 3, this study.

(Martin 1993). High field strength elements (HFSE) are highly depleted, with Lan/Nbn (normalized to normal mid-ocean ridge basalt, N-MORB) values varying from 5.13 to 7.89. These values .3.5 are suggestive of a subduction zone (Drummond & Defant 1990); this means that the protolith of these conglomerates was subduction-related rocks that left refractory garnet and/or amphibole in the source. In contrast, in Type 1, the values of La/Nb (1.5 – 3.19) ,3.4 are in agreement with those reported for modern basalts from an oceanic back-arc basin environment (Fryer et al. 1990; Monnier et al. 1995). For comparison, the REE patterns of metabasalts, garnet-free amphibolites (Soumaila et al. 2004), and the granodiorite of Toure´ (Ama-Salah et al. 1996) are given in Figure 3. The garnet-amphibolites, the garnet-free amphibolites and the kyanite – staurolite-bearing micaschists are interbedded. However, we note that only garnet-free amphibolites and kyanite – staurolite-bearing micaschists have comparable REE patterns; they are interpreted to be have been derived from the same protoliths. The garnet-free amphibolites have high bulk REE contents, and are highly enriched in both compatible and incompatible elements, with Ni 90 – 320 ppm, Cr 345 – 628 ppm, and Zr 118 – 485 ppm (Soumaila et al. 2004). Two groups of metabasalts are distinguished: a group 1 with flat to slightly depleted REE patterns and a group 2 that is moderately enriched in light REE (LREE). All the chemical compositions of these rocks can be explained by production in an oceanic arc to back-arc-basin

24

A. SOUMAILA ET AL.

and micaschists have values from þ0.3 + 0.4 to þ0.9 + 0.5, which are lower than those of the garnet-free amphibolites (þ1.3 + 0.3 and þ1.7 + 0.3). TDM ranges from 2.07 to 2.22 Ga for metabasalts; however, samples 821 and 348 have lower TDM values (1.83 and 1.75 Ga, respectively); this might be due to a disturbance of the Nd–Sm system. Garnet-free amphibolites and micaschists have slightly higher values of TDM (2.2–2.3).

environment by a slab-derived siliceous melt that had metasomatized a depleted mantle source (Soumaila et al. 2004).

Nd – Sm isotopic data The Nd–Sm isotopic analysis was carried out on metabasalts, amphibolites, and micaschists (Type 1 sediments) and metaconglomerate (Type 2 sediments) at the GEOTOP laboratory (Montreal, Canada) according to a method described by Henry et al. (1998). Samarium and neodymium were loaded on a double Re–Ta filament and analysed in static and dynamic multi-collector mode on a VG Sector 54 mass spectrometer. The La Jolla Nd standard yielded a 143 Nd/144Nd ratio of 0.511849 + 12 and blanks for Nd and Sm were less than 50 pg. The 143Nd/144Nd values were normalized to 146Nd/144Nd ¼ 0.7219, and the CHUR composition used to calculate 1Nd values is 143Nd/144Nd ¼ 0.512638 and 147Sm/144Nd ¼ 0.1967; the decay constant is 6.54 10212, Nd model ages are calculated according to the depleted mantle model of Ben Othman et al. (1984), and the results of Nd and Sm isotopic measurements are reported in Table 3. For comparison, Nd and Sm data for metabasalts and garnet-free amphibolites (Soumaila et al. 2004) are also given. The 1Nd(T) values are back-calculated at 2.15 Ga, corresponding to the crystallization age of the Te´ra pluton (Cheilletz et al. 1994), which allows us to compare, at a given age, the enriched or depleted character of the sources of these rocks. The metabasalts show higher values ranging from þ2.1 + 0.7 to þ3.3 + 0.6, except for samples 216 and 420 (with 1Nd ¼ þ0.9 + 0.6 and þ1.6 + 0.6, respectively). Garnet-amphibolites

U – Pb data Zircon grains of 80–200 mm in size were separated from the Dargol granodiorite (sample 6799) and micaschists (samples 6836 and 6880), and were mounted in epoxy resin together with a standard, then polished and gold coated. Zircons are then observed by back-scattered scanning electron microscopy and cathodoluminescence to characterize their internal structures (Fig. 4). The zircon grains of the Dargol granodiorite are euhedral, locally corroded, with a tiny whitish rim indicating a very slight alteration (Fig. 4a). The detrital zircons of the micaschists are fresh, rounded, ovoid, or subeuhedral, both being slightly cracked and pitted (Fig. 4b and c). Zircon 12 (Fig. 4b) shows a rounded, darker off-centre core, which could be an older zircon. The zircons of the metaconglomerates are euhedral to sub-euhedral, and show thin cracks and rare pits (Fig. 3d). The analyses were performed on the CRPG-CNRS Cameca IMS-1270 ion microprobe. A detailed analytical procedure has been given by Deloule et al. (2002). U– Pb isotopic compositions were determined on single zircons. Data are presented in

Table 3. Sm–Nd concentrations and isotopic compositions Sample no. Sm (ppm) Nd (ppm) 216 232 348 414 420 514 804 810 812 821 255 262 268 287 129 135

3.69 3.21 2.67 2.95 4.01 1.97 1.37 1.42 1 1.31 22.1 25.5 2.57 4.56 3.84 3.82

10.52 9.81 8.23 9.29 16.73 6.04 3.7 4.34 3.52 4.97 149.7 178.1 11.97 22.22 18.94 18.13

147

Sm/144Nd 0.2121 0.198 0.196 0.1919 0.1449 0.197 0.2234 0.198 0.2259 0.2205 0.0892 0.0864 0.1220 0.1239 0.1225 0.1272

143

Nd/144Nd 0.5129 0.5128 0.5128 0.5127 0.5112 0.5128 0.5131 0.5128 0.5132 0.5131 0.5112 0.5111 0.5117 0.5116 0.5116 0.5117

+2s

1Nd (2.15 Ga)

0.000008 0.000011 0.000007 0.000011 0.000010 0.000013 0.000013 0.000014 0.000013 0.000012 0.000007 0.000008 0.000013 0.000008 0.000008 0.000009

0.9 2.6 3.3 2.6 1.6 2.9 2.4 2.4 2.5 2.1 1.7 1.3 0.8 0.4 0.9 0.3

+s TDM (Ga) 0.6 0.6 0.6 0.6 0.6 0.6 0.7 0.7 0.7 0.7 0.3 0.3 0.5 0.4 0.4 0.4

2.17 2.07 1.75 2.07 2.24 1.92 2.14 2.15 2.2 1.83 2.2 2.22 2.27 2.3 2.7 2.28


25

Fig. 4. Selected scanning electron microscopy (SEM) images of zircon grains from the Dargol granodiorite (a), micaschists (b and c), and metaconglomerates (d).

Table 4, and weighted mean ages and discordia lines were determined using the Isoplot program (Ludwig 2000).

The Dargol granodiorite (sample 6799) Five zircon grains are plotted in Figure 5a. Zircon 8 gives a near-concordant age of 2174 + 4 Ma (MSWD ¼ 0.028). This age is interpreted as the crystallization age of the Dargol pluton. The other zircons define a discordia with an upper intercept at 2169 + 37 Ma (MSWD ¼ 18), which is identical to the crystallization age.

the zircon grains of sample 6880 are discordant (Fig. 7a). Zircons 2, 5 and 7 define a discordia with an imprecise upper intercept, with a minimum age of 2208 + 390 Ma (MSWD ¼ 17) (Fig. 7b); this indicates a disturbance of the U –Pb system in relation to the metamorphic event. The ages 2273 + 19 Ma to 2278 + 5 are considered to be the time interval of crystallization of the protoliths of the micaschists. The discordant age of 2400 + 42 Ma may indicate a contribution of early Palaeoproterozoic material in the deposition of sediments from which the micaschists were derived.

The micaschists (samples 6836 and 6880) In sample 6836 (Fig. 6a), zircon 2 yields a concordant age at 2278 + 5 Ma. The other zircon grains are discordant and define two discordia: zircons 7 and 14, and the rim of grain 12 give an upper intersect at 2273 + 19 Ma (MSWD ¼ 2.6) (Fig. 6a), and zircons 4 and 15, and the core of zircon 12, give an upper intersect at 2400 + 42 Ma (MSWD ¼ 0.054) (Fig. 6b). In sample 6880, all

The metaconglomerates (sample 784) Six detrital zircons from sample 784 were analysed; they are all discordant, and define a discordia with an upper intercept at 2187 + 55 Ma (Fig. 5b), which is regarded as the crystallization age of the granitoids reworked by these sediments. This age is identical to that of Dargol granodiorite (2174 + 4 Ma) within analytical error, and to the

26

Table 4. U –Pb isotopic data for the zircon grains analysed Analysis no.

Corrected isotopic ratio (%) 206

238

U+s

207

235

U+s

207

Ages (Ma) 206

Pb + s

206

238

U+s

207

Pb/235U + s

207

Pb/206Pb + s

Pb (ppm)

U (ppm)

181.4 109.9 107.7 126.9 238.7

2034.9 397.4 485.1 1344.9 691.2

0.1038 (0.74) 0.3217 (0.1) 0.2585 (0.26) 0.1098 (0.76) 0.4020 (0.14)

1.3966 (1.83) 5.7961 (1.50) 4.6458 (1.43) 1.5550 (2.94) 7.5241 (1.41)

0.0976 (0.74) 0.1307 (0.09) 0.1304 (0.26) 0.1027 (0.76) 0.1358 (0.15)

637 + 11 1798 + 23 1482 + 18 672 + 18 2178 + 25

888 + 11 1946 + 12 1758 + 12 953 + 18 2176 + 12

1579 + 13 2107 + 1 2103 + 3 1674 + 13 2174 + 2

24.3 18.4 14.5 39.4 24.5 30.1 53.0

66.2 54.7 39.9 110.1 68.7 88.4 168.7

0.4269 (0.96) 0.3918 (0.30) 0.4236 (0.24) 0.416 (0.18) 0.4146 (0.17) 0.3969 (0.44) 0.3658 (0.34)

8.9978 (1.54) 7.7330 (1.46) 8.4198 (1.54) 8.2379 (1.38) 8.1757 (0.76) 8.1682 (0.85) 7.3402 (0.78)

0.1528 (0.96) 0.1432 (0.30) 0.1442 (0.24) 0.1436 (0.18) 0.1430 (0.17) 0.1493 (0.44) 0.1455 (0.34)

2292 + 29 2131 + 26 2277 + 29 2243 + 26 2236 + 14 2155 + 15 2010 + 13

2338 + 16 2200 + 13 2277 + 14 2258 + 12 2251 + 7 2250 + 8 2154 + 7

2378 + 16 2266 + 5 2278 + 3 2271 + 2 2264 + 2 2337 + 7 2294 + 5

39.0 29.2 52.6 32.9 59.6 110.9

120.2 86.8 161.9 99.5 219.6 599.2

0.3775 (0.17) 0.3916 (0.11) 0.37856 (0.09) 0.3854 (0.18) 0.3159 (0.15) 0.2154 (2.28)

7.5819 (0.8) 7.3938 (097) 7.3430 (0.81) 7.8674 (0.83) 6.0338 (0.90) 3.8537 (4.17)

0.1457 (0.17) 0.1369 (0.11) 0.1407 (0.08) 0.1481 (0.17) 0.1386 (0.15) 0.1298 (2.28)

2065 + 14 2130 + 17 2070 + 14 2102 + 14 1770 + 13 1257 + 47

2183 + 7 2160 + 8 2154 + 7 2216 + 7 1981 + 7 1604 + 37

2296 + 2 2189 + 1 2236 + 1 2324 + 3 2209 + 2 2095 + 40

50.6 55.5 41.3 48.6 42.3 61.9

279.7 264.6 185.2 161.0 172.0 383.5

0.2107 (0.05) 0.2439 (0.37) 0.2594 (0.17) 0.3513 (0.28) 0.2863 (0.35) 0.1880 (0.44)

4.0860 (0.83) 4.6490 (0.92) 4.8118 (1.02) 6.6314 (0.87) 5.4509 (1.31) 3.5356 (1.13)

0.1407 (0.05) 0.1382 (0.37) 0.1346 (0.17) 0.1369 (0.28) 0.1381 (0.35) 0.1364 (0.44)

1232 + 9 1407+ 11 1487 + 13 1941 + 14 1623 + 18 1111 + 53

1652 + 6 1758 + 8 1787 + 8 2064 + 8 1893 + 11 1535 + 9

2236 + 0 2205 + 6 2158 + 2 2189 + 4 2203 + 6 2182 + 7

Pb/

Pb/

Pb/

Pb/

A. SOUMAILA ET AL.

Granodiorite of Dargol 3 4 6c 7 8 Micaschist sample 6836 (129) 12b 12c 2 7 14 15 4 Micaschist sample 6880 (135) 1 2 3 4 5 7 Metaconglomerate sample 784 12 19 1 2 3 5

Concentrations


0.5 0.4

2400

0.44

3

0.2

800

0.1

400

7

Age Concordia 2173.8 ± 3.9 Ma MSWD = 0.028

Pb / 238U

4

1600 1200

2320

7

2000

0.3

2

micaschist 6836

a

8

206

206Pb / 238U

a

Granodiorite of Dargol

27

Intercept at 335 ± 82 & 2169 ± 37Ma MSWD = 18

0.42 0.40

2240

14

2160

0.38

12rim

Intercept at 263 ± 890 & 2273 ± 19Ma MSWD = 2.6

0 0

2

4 207

6 Pb / 235U b

Conglomerate 784

0.4

0.36 7.2

2400

7.6

8.0

8.4

8.8

0.46

2000 1600

0.44

1200

0.2

sample 784:

800

discordia Intercept at –31 ± 210 & 2187 ± 55 Ma

0.1 400

MSWD = 71

2

4 207

2300

6

12core

0.42 2200

0.40

8

10

Pb / 235U

Fig. 5. U –Pb zircon data of the Dargol granodiorite (a) and of the conglomerate sample 784 (b).

age of 2171 + 68 Ma reported by Cheilletz et al. (1994) for the Tondia granite (Fig. 8).

Discussion and conclusion The REE patterns, the Nd –Sm isotopic features and the U –Pb ages of the Birimian metamorphic rocks of the Diagorou –Darbani greenstone belt, Liptako, are in part, comparable with those of the Sirba greenstone belt (Ama Salah et al. 1996), and more generally with the results reported from the other Birimian domains in the Man rise (e.g. Abouchami et al. 1990; Boher et al. 1992; Sylvester & Attoh 1992). The tholeiitic metabasalts, the garnetamphibolites, garnet-free amphibolites, and micaschists (Type 1 sediments) are variously depleted or enriched in LREE, and have been interpreted as oceanic arc related (Ama-Salah et al. 1996) or oceanic arc and back-arc basin related rocks (Soumaila 2000) with the participation of a metasomatized depleted mantle slab (Soumaila et al. 2004). The Type 2 sediments, corresponding to weakly metamorphized monogenetic conglomerates, have TTG-type REE patterns, which may indicate derivation from a TTG-like protolith;

15

2100

0.38 0.36

0 0

2400

b

Dargol pluton

Pb / 238U

0.3

206

Pb / 238U

10

207Pb / 235U

0.5

206

8

0.34 6.6

4

7

Intercept at 858 ± 240 & 2400 ± 42Ma MSWD = 0.054

7.4 7.8 8.2 8.6 207Pb / 235U

9

9.4 9.8

Fig. 6. U–Pb data for detrital zircons of the micaschist sample 6836, with two discordia (a) and (b).

that is, a granitoid generated by partial melting of the basaltic crust metamorphosed to garnetamphibolites in subduction zone. This interpretation is different from that of Ama-Salah et al. (1996) for the Sirba area. According to those workers, the chemical character of the granitoid of Sirba area is related to a normal oceanic arc with partial melting of a metasomatized mantle slab. The Nd isotopic data indicate the primary character of these rocks, with whole-rock (2.15 Ga) positive 1Nd values that vary from þ0.3 + 0.4 to þ3.9 + 0.6 for all samples; such values have been reported by Boher et al. (1992) for the West African Birimian rocks. However, we can distinguish the Type 1 sediments, garnet- amphibolites and garnet-free amphibolites on one hand, and the metabasalts on the other hand. The former have lower 1Nd values (þ0.3 to þ1.7) with TDM ¼ 2.2–2.3 Ga, whereas the latter have slightly higher 1Nd values (þ2.1 to þ3.3), and TDM ranging from 2.07 to 2.22 Ga. The Dargol granodiorite is not a pre-Birimian migmatites as described by Machens (1973); the crystallization age (2174 + 4 Ma) belongs to the

206Pb/ 238U

28

A. SOUMAILA ET AL. Micaschist

0.38

sample 6880

2100 1900

0.34 1700

0.30

2 4 31

5

1500

0.26 0.22

2300

a

0.42

1300

7

0.18 0.14 2

4

6

8

207Pb / 235U

0.42 0.38 206Pb / 238U

2200

b

2

0.34

1800 5

0.30 0.26 1400

Intercepts at 217 ± 1900 & 2208 ± 390 Ma MSWD = 17

0.22

7 0.18 0.14 2.5

3.5

4.5 207

5.5

Pb /

235

6.5

7.5

8.5

U

Fig. 7. U–Pb data for all detrital zircons of micaschist sample 6880 (a), and the discordia defined by zircon grains 2, 5 and 7 (b).

previously defined age interval of 2158 + 9 Ma to 2188 + 12 Ma, reported from various granitoids of the Liptako regions (Fig. 8). These are coeval with belt-type plutons dated from 2180 to 2150 Ma elsewhere in West Africa (Hirdes et al. 1992; Davis et al. 1994; Kouamelan 1996; Kouamelan et al. 1997a, b; Doumbia et al. 1998; Hirdes & Davis 1998; Oberthur et al. 1998; Loh & Hirdes 1999; Egal et al. 2002; Lahonde`re et al. 2002). If Type 2 sediments seem to derive from erosion of well-known granitoid dated from 2158 + 9 Ma to 2188 + 12 Ma, the detrital zircon Type 1 sediments, with concordant ages of 2273 + 19 Ma to 2278 + 5 Ma, may indicate a contribution from an early unknown Palaeoproterozoic material, which seems to be emplaced in an oceanic arc environment. This is suggested by the REE patterns of the micaschists. However, these samples reveal a heterogeneous detrital zircon ages from 2.26 Ga to 2.4 Ga, indicating (1) a probable contribution of various sources of early Palaeoproterozoic ages or (2) that some zircons are of Archaean inheritance with lead loss. This latter hypothesis is uncertain because it has been emphasized that Birimian

terranes are juvenile without Archaean participation (Abouchami et al. 1990). The first hypothesis is thus in agreement with TDM ages (2.2–2.3 Ga) close to the ages of crystallization of protoliths (2273– 2278 Ma). In the Liptako area, these early Palaeoproterozoic rocks could have been generated in relation to subduction zones. Such early Palaeoproterozoic ages (2200–2312 Ma) have been recorded in many areas of the Man rise, and in zircon grains from metasedimentary rocks or various granitoids of Ghana, Ivory Coast, Senegal or Guinea (Davis et al. 1994; Bossie`re et al. 1996; Dia et al. 1997; Doumbia et al. 1998; Loh & Hirdes 1999; Lahonde`re et al. 2002; Gasquet et al. 2003). They have already been interpreted by the listed researchers as corresponding to an early Birimian magmatic event, which was the initial Palaeoproterozoic crustal growth (e.g. Kouamelan et al. 1997a, b). In conclusion, the geochemical features of the metamorphic rocks of the Diagorou– Darbani greenstone belt (Soumaila et al. 2004), in addition to the present U –Pb and Nd–Sm isotopic data, emphasize a multi-stage crustal genesis and growth in the Liptako area during Palaeoproterozoic times. Such evolution can be summarized as follows. (1) At about 2300 Ma, oceanic arc related calc-alkaline rocks (garnet-free amphibolites and Type 1 sediments) were generated, and are comparable with those of modern oceanic arc magmas (Fig. 9). This means partial melting of a mantle slab metasomatically enriched by siliceous fluids switching from a subducted oceanic crust (Soumaila et al. 2004). This hypothesis explains the paradoxical features of some garnet-free amphibolites that are highly enriched both in compatible and incompatible elements. (2) Oceanic back-arc environment development occurred around 2200 Ma, with emplacement of thick tholeiitic basalts characterized by flat to slightly depleted REE patterns (Soumaila et al. 2004). (3) The time interval of 2190–2150 Ma was marked by the emplacement and the expansion of large granitoid plutons such as the Dargol pluton, which could be of TTG or normal oceanic arc calc-alkaline granitoid affinity. Granitoids of TTG affinity may have been eroded around 2150– 2100 Ma, resulting in the deposition of Type 2 sediments (Fig. 9). The proposed model could explain the overall range of geochemical, isotopic and geochronological results recorded on the Birimian crust of the West African craton. It also integrates the findings on the existence of an early Palaeoproterozoic magmatic event of crustal growth. We propose that during Palaeoproterozoic times, crustal genesis took place in an oceanic arc environment. This


29

Fig. 8. General map of the Liptako area showing main ages recorded on various rocks. Data sources (superscript numbers): 1, Abdou et al. (1998); 10 , Abdou et al. (1992); 2, Boher et al. (1992); 3, Abouchami et al. (1990); 4, Cheilletz et al. (1994); 5, Ama-Salah et al. (1996); 6, Machens (1973); 7, Le´ger et al. (1992); 8, Klockner (1991); 9, this study. Dating methods: U– Pb; Rb–Sr; K–Ar, Pb–Pb; Evap (evaporation).

means that the granitoid rocks were generated by partial melting of a metasomatically enriched mantle slab (calc-alkaline affinities) or by partial melting of subducted oceanic crust metamorphosed

to garnet-amphibolites (TTG affinities). The process of crustal growth may have been the result of a continuous input of crustal materials of early Palaeoproterozoic age (2.3 Ga) to 2.15 Ga.

30

A. SOUMAILA ET AL.

Fig. 9. (La/Yb)N v. (Yb)N, for rocks of the Diagorou –Darbani greenstone belt compared with Archaean and modern processes (Martin 1993).

B. Bonin and J. Abati are thanked for their helpful reviews of the manuscript, and their pertinent criticisms that improved the final version. We thank UNESCO for financial support during the 3rd IGCP symposium of GAO, Mali, 2005.

References A BDOU , A., B LIN , G., K ADEY , B. D., H IRBEC , Y., R EGNOULT , J. M. & Y OUNFA , I. 1992. Mise en exećution du plan de de´veloppement mine´ral du Niger. Unpublished report, Ministry of Mines and Energy of Niger, Niamey. A BDOU , A., B ONNOT , H., K ADEY , B. D., C HALAMET , D., S AINT M ARTIN , M. & Y OUNFA , I. 1998. Notice explicative des cartes geólogiques du Liptako a` 1/ 100 000 et 1/200 000. Direction de la Recherche Geólogique et Minie`re, Ministe`re des Mines et de la Geólogie, Niamey, Niger. A BOUCHAMI , W., B OHER , M., M ICHARD , A. & A LBARE` DE , F. 1990. A major 2.1 Ga event of mafic magmatism in West Africa: an early stage of crustal accretion. Journal of Geophysical Research, 95, 17605– 17629. A MA -S ALAH , I., L IE´ GEOIS , J. P. & P OUCLET , A. 1996. E´volution d’un arc insulaire oceánique birimien prećoce au Liptako nige´rien (Sirba): geólogie, geóchronologie et geóchimie. Journal of African Earth Sciences, 22, 235 –254.

B EN O THMAN , B., P OLVE´ , M. & A LLE` GRE , C. J. 1984. Nd– Sr isotopic composition of granulites and constraints on the evolution of the lower continental crust. Nature, 307, 510 –515. B OHER , M., A BOUCHAMI , W., M ICHARD , A., A LBARE` DE , F. & A RNDT , N. 1992. Crustal growth in West Africa at 2.1 Ga. Journal of Geophysical Research, 97, 345–369. B OSSIE` RE , G., B ONKOUNGOU , I., P EUCAT , J. J. & P UPIN , J. P. 1996. Origin and age of Palaeoproterozoic conglomerates and sandstones of the Tarkwaian Group in Burkina Faso, West Africa. Precambrian Research, 80, 153–172. C HEILLETZ , A., B ARBEY , P., L AMA , C., P ONS , J., Z IMMERMANN , J. L. & D AUTEL , D. 1994. Age de refroidissement de la crouˆte juveńile Birimienne d’Afrique de l’Ouest. Donneés U–Pb, Rb–Sr et K–Ar sur les formations a` 2.1 Ga du SW Niger. Comptes Rendus de l’Acade´mie des Sciences, 319, 435–442. D AVIS , D. W., H IRDES , W., S CHALTEGGER , U. & N UNOO , E. A. 1994. U–Pb constraints on deposition and provenance of Birimian and gold-bearing Tarkwaian sediments in Ghana, West Africa. Precambrian Research, 67, 89–107. D ELOULE , E., A LEXANDROV , P., C HEILLETZ , A., L AUMONIER , B. & B ARBEY , P. 2002. In-situ U– Pb zircon ages for early Ordovician magmatism in the eastern Pyrenees, France: the Canigou orthogneisses. International Journal of Earth Sciences, 91, 398– 405. D IA , A., V AN S CHMUS , W. R. & K RO¨ NER , A. 1997. Isotopic constraints on the age and formation of a

REE, Nd, U– Pb, AND BIRIMIAN CRUSTAL GENESIS Palaeoproterozoic volcanic arc complex in the Kedougou Inlier, eastern Senegal, West Africa. Journal of African Earth Sciences, 24, 197–213. D OUMBIA , S., P OUCLET , A., K OUAMELAN , A. N., P EUCAT , J. J., V IDAL , M. & D ELOR , C. 1998. Petrogenesis of juvenile type Birimian Palaeoproterozoic granitoids in Central Coˆte d’lvoire (West Africa): geochemistry and geochronology. Precambrian Research, 87, 33– 63. D RUMMOND , M. S. & D EFANT , M. J. 1990. A model for trondhjemite–tonalite –dacite genesis and crustal growth via slab melting: Archean to modern comparisons. Journal of Geophysical Research, 95, 21503– 21521. D UPUIS , D., P ONS , J. & P ROST , A. E. 1991. Mise en place des plutons et caracte´risation de la de´formation birimienne au Niger occidental. Comptes Rendus de l’Acade´mie des Sciences, 312, 769–773. E GAL , E., T HIE´ BLEMONT , D. & L AHONDE` RE , D. 2002. Late Eburnean granitization and tectonics along the western and northwestern margin of the Archean Keńe´ma–Man domain (Guinea, West African Craton). Precambrian Research, 117, 57– 84. F RYER , P., T AYLOR , B., L ANGMUIR , C. H. & H OCHSTAEDTER , A. G. 1990. Petrology and geochemistry of lavas from Sumisu and Torishima back-arc rift. Earth and Planetary Science Letters, 100, 161– 178. G ASQUET , D., B ARBEY , P., A DOU , M. & P AQUETTE , J. L. 2003. Structure, Sr –Nd isotope geochemistry and zircon U–Pb geochronology of the granitoids of the Dabakala area (Coˆte d’Ivoire): evidence for a 2.3 Ga crustal growth event in the Palaeoproterozoic of West Africa. Precambrian Research, 127, 329–354. H ENRY , P., S TEVENSON , R. K. & G ARIE´ PY , C. 1998. Late Archean mantle composition and crustal growth in the Western Superior Province of Canada: neodymium and lead isotopic evidence from the Wawa, Quetico and Wabigoon subprovinces. Geochimica et Cosmochimica Acta, 62, 143–157. H IRDES , W. & D AVIS , D. W. 1998. First U–Pb zircon age of extrusive volcanism in the Birimian Supergroup of Ghana, West Africa. Journal of African Earth Sciences, 27, 291–294. H IRDES , W., D AVIS , D. W. & E ISENLOHR , B. N. 1992. Reassessment of Proterozoic ages in Ghana on the basis of U– Pb zircon and monazite dating. Precambrian Research, 56, 89–96. K LOCKNER , I. A. 1991. Cartographie geólogie du sillon de Te´ra, E´tude geóchronologique. Unpublished report, Ministry of Mines and Energy of Niger, Niamey, 49–59. K OUAMELAN , A. N. 1996. Geóchronologie et geóchimie des formations archeénnes et Prote´rozoı¨ques de la Dorsale de Man en Coˆte d’lvoire. Implications pour la transition Archeén Prote´rozoı¨que. Thesis, Rennes 1 University, Rennes. K OUAMELAN , A. N., D ELOR , C. & P EUCAT , J. J. 1997a. Geochronological evidence for reworking of Archean terrains during the early Proterozoic (2.1 Ga) in the western Coˆte d’Ivoire (Man Rise (West African Craton). Precambrian Research, 86, 177–199. K OUAMELAN , A. N., P EUCAT , J. J. & D ELOR , C. 1997b. Reliques archeénnes (3,15 Ga) au sein du magmatisme

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birimien (2,1 Ga) de Coˆte d’Ivoire, craton ouest-africain. Comptes Rendus de l’Acade´mie des Sciences, 324, 719–727. L AHONDE` RE , D., T HIEBLEMONT , D., T EGYEY , M., G UERROT , C. & D IABATE´ , B. 2002. The first evidence of Pre-Birimian volcanic and associated rocks of the Niani suite. Journal of African Earth Sciences, 35, 417– 431. L E´ GER , J. M., L IE´ GEOIS , J. P., P OUCLET , A. & V ICAT , J. P. 1992. Occurrence of syntectonic alkali–pyroxene granites of Eburnean (2.1 Ga) in W Niger. In: Abstracts, 14e`me Reúnion Annuelle des Sciences de la Terre, Toulouse, France. Socie´te´ Geólogique de France, 96. L OH , G. & H IRDES , W. 1999. Explanatory notes for the geological map of Ghana 1:100 000: Sekondi (0402A) and Axim (0403B) sheets. Geologisches Jahrbuch, B93. L UDWIG , K. R. 2000. Isoplot/Ex 2.49. A geochronological toolkit for Microsoft Excel. Berkeley Geochronological Center, Special Publications, 1a. M ACHENS , E. 1973. Contribution a` l’e´tude des formations du socle cristallin et de la couverture se´dimentaire de l’Ouest de la Re´publique du Niger. Me´moires du BRGM, 82. M ARTIN , H. 1993. The mechanisms of petrogenesis of the Archaean continental crust—comparison with modern processes. Lithos, 30, 373– 388. M ILE´ SI , J. P., F EYBESSE , J. L., L EDRU , P. ET AL . 1989. Les mine´ralisations aurife`res de l’Afrique de l’Ouest. Chronique de la Recherche Minie`re, 497, 3 –98. M ONNIER , C., G IRARDEAU , J., M AURY , R. C. & C OTTON , J. 1995. Back-arc-basin origin for the East Sulawesi ophiolite (eastern Indonesia). Geology, 23, 851–854. O BERTHU¨ R , T., V ETTER , U., D AVIS , D. W. & A MANOR , J. A. 1998. Age constraints on gold mineralization and Palaeoproterozoic crustal evolution in the Ashanti belt of southern Ghana. Precambrian Research, 89, 129– 143. P ONS , J., B ARBEY , P., D UPUIS , D. & L EGER , J. M. 1995. Mechanisms of pluton emplacement and structural evolution of 2.1 Ga juvenile continental crust: the Birimian of southwestern Niger. Precambrian Research, 70, 281– 301. P OUCLET , A., P ROST , A. E., A MA -S ALAH , I. & L APIERRE , H. 1990. Les ceintures birimiennes du Niger occidental (Prote´rozoı¨que infe´rieur), nouvelles donneés pe´trologiques et structurales des formations me´tavolcaniques. Comptes Rendus de l’Acade´mie des Sciences, 311, 333–340. S OUMAILA , A. 2000. E´tude structurale, pe´trographique et geóchimique de la ceinture birimienne de Diagorou– Darbani, Liptako, Niger occidental (Afrique de l’Ouest). Thesis, University of Franche-Comte´, Besançon. S OUMAILA , A. & K ONATE´ , M. 2005. Caracte´risation de la de´formation dans la ceinture birimienne (paleóprote´rozoı¨que) de Diagorou– Darbani (Liptako nige´rien, Afrique de l’Ouest). Africa Geoscience Review, 12, 161– 178. S OUMAILA , A., H ENRY , P. & R OSSY , R. 2004. Contexte de mise en place des roches basiques de la ceinture de roches vertes birimienne de Diagorou– Darbani (Liptako, Niger, Afrique de l’Ouest): plateau oceánique ou environnement d’arc/bassin arrie`re-arc

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oceánique. Comptes Rendus Geósciences, 336, 1137–1147. S YLVESTER , P. J. & A TTOH , K. 1992. Lithostratigraphy and composition of 2.1 Ga greenstone belts of the West African Craton and their bearing on crustal evolution and the Archean– Proterozoic boundary. Journal of Geology, 100, 377– 393.

V IDAL , M., D ELOR , C., P OUCLET , A., S IME´ ON , Y. & A LRIC , G. 1996. E´volution geódynamique de l’Afrique de l’Ouest entre 2.2 Ga et 2 Ga: le style archeén des ceintures vertes et des ensembles se´dimentaires Birimiens du nord-est de la Coˆte d’Ivoire. Bulletin de la Socie´te´ Geólogique de France, 167, 307 – 319.

Two Mesoarchaean terranes in the Reguibat shield of NW Mauritania R. M. KEY1, S. C. LOUGHLIN1, M. GILLESPIE1, M. DEL RIO1, M. S. A. HORSTWOOD2, Q. G. CROWLEY2, D. P. F. DARBYSHIRE2, P. E. J. PITFIELD3 & P. J. HENNEY3 1

BGS, Murchison House, West Mains Road, Edinburgh, EH9 3LA, UK (e-mail: [email protected]) 2

NERC Isotope Geosciences Laboratory, Keyworth NG12 5GG, UK 3

BGS, Keyworth NG12 5GG, UK

Abstract: Two domains have previously been recognized in the Archaean Reguibat shield of NW Mauritania, based primarily on their gross lithological differences. New fieldwork has identified a major ductile shear zone (Taˆçaraˆt –Inemmauˆdene Shear Zone) separating these domains and new geochronological studies show that the two domains record different Mesoarchaean histories. As such, the two domains are redefined as the Choum–Rag el Abiod Terrane and Tasiast– Tijirit Terrane. Previous isotopic studies of metamorphic lithologies of the eastern Choum–Rag el Abiod Terrane indicate a succession of crustal growth from about 3.5– 3.45 Ga to between about 3.2 and 2.99 Ga. Isotopic data presented in this contribution from the Tasiast– Tijirit Terrane indicate that emplacement of major calc-alkaline plutons occurred at c. 2.93 Ga after volcanism (preserved as greenstone belts) that included late felsic eruptive centres dated at c. 2965 Ma. This Mesoarchaean intrusive and extrusive magmatism was confined to the Tasiast–Tijirit Terrane, where it was emplaced through migmatitic orthogneisses that are the oldest lithodemic unit of the Tasiast–Tijirit Terrane. Widespread bimodal, post-tectonic magmatism in both terranes included major granitic magmatism dated at c. 2730 Ma. The north–south- to NNE–SSW-trending curvilinear Taˆçaraˆt –Inemmauˆdene Shear Zone that separates the two terranes records late intense transpressive ductile shearing. It has a flower structure over a horizontal distance of about 70 km across its southern portion with unquantifiable sinistral horizontal offset, and east-directed thrusting on its eastern side where it cuts into the Choum–Rag el Abiod Terrane. A new U– Pb zircon age of 2954+111 Ma is presented for a deformed granite confined within the central part of this shear zone. A minimum age for the shearing is provided by a previously determined c. 2.73 Ga age for a post-tectonic granite that cuts across the easternmost part of the shear zone in the Choum– Rag el Abiod Terrane.

The Reguibat (also spelt R’gueibat) shield (Menchikoff 1949) forms the exposed NW part of the West African craton that underlies much of NW Africa. This craton is a large segment of Precambrian crust (c. 4.5 106 km2) stable since about 1700 Ma and bounded on all sides by Pan-African orogenic belts (Cahen et al. 1984). However, much of the craton is concealed beneath a cover of various unmetamorphosed sedimentary strata, and unconsolidated superficial deposits of the Sahara Desert. The Reguibat shield refers to the northwestern part of the craton west of the Taoudeni basin (Fig. 1), whereas the southern exposed part is referred to as the Leo shield. Dillon & Sougy (1974) and Bessoles (1977) identified two provinces within the Reguibat shield: a SW Province composed of rocks older than c. 2000 Ma (i.e. pre-‘Eburnean cycle’) that includes the rocks described in the present paper, and a central and northeastern province of Eburnean rocks (recently described by Schofield et al. 2006).

Subsequently, Rocci et al. (1991) divided the Reguibat shield into two main parts including an ‘Archaean shield’ in northern Mauritania (part of the SW Province), itself divided into an eastern Amsaga –Tiris– Ouassat domain and a western Tasiast –Lebzenia domain. Our mapping in NW Mauritania in the general area west of Atar (Key 2003; Key & Loughlin 2003; Key et al. 2003; Pitfield et al. 2005), defined the geographical limits of these two domains and their contact relationship. The eastern and western domains are renamed the Choum– Rag el Abiod Terrane and Tasiast –Tijirit Terrane, respectively, after newly defined type areas (Pitfield et al. 2005). The redefinition of the two crustal segments as terranes is because the isotopic evidence from the present and previous studies shows that they preserve different Mesoarchaean geological histories indicative of different geotectonic settings and because they are separated by a major ductile shear zone, now named the Taˆçaraˆt– Inemmauˆdene Shear Zone (TISZ). (The shear zone is named after two


34 R. M. KEY ET AL. Fig. 1. The major geological features of the study area. The granite in the Choum Rag el Abiod column is the Touijenjert Granite. Circled letters show the locations of the dated samples: A, Augen granite gneiss, sample 201401619; B, epidote-tonalite of the Gleibat el Fhoud Suite, sample 201401584; C, Bir Igueni Granite, sample 201401598; D, felsic metavolcanic rock from the Chami Greenstone Belt, sample 201500738.

MESOARCHAEAN SHIELDS, NW MAURITANIA

locations where it is best exposed. There is a nearcontinuous section across the shear zone in the Taˆçaraˆt (a NE –SW trending strip of land within the Akchaˆr dune field to the NW of Chami), and strongly deformed granites within the shear zone form large inselbergs at Inemmauˆdene.)

Choum – Rag El Abiod Terrane Potrel et al. (1998) considered that the Amsaga region of the Reguibat shield (the Choum–Rag el Abiod Terrane) formed through a series of crustal growth events between about 3.50– 3.45 Ga and 2.73 Ga. An early event is recorded in migmatitic orthogneisses at about 3.5–3.45 Ga (Auvray et al. 1992; Potrel 1994; Potrel et al. 1996). Zircons from a single migmatitic gneiss were dated by the sensitive high-resolution ion microprobe (SHRIMP) U –Pb method to give ages of between 3515+15 Ma and 3422+ 10 Ma with Nd model ages of about 3.64 Ga (Potrel et al. 1996). The migmatites are interlayered with mafic gneisses of possible volcanic origin, as well as with undoubted metasedimentary rocks. These include marbles and hyperaluminous gneisses (including K-feldspar–quartz–sillimanite–biotite– hercynite –garnet-gneisses and sillimanite – cordierite –garnet-gneisses). A granulite-facies metamorphism was interpreted to have occurred between about 3.20 and 2.99 Ga based on SHRIMP dating of zircons from charnockites dated at 2986 +8 Ma with a whole-rock Sm–Nd isochron at 3012+ 142 Ma (Potrel et al. 1998) and Nd model ages of about 3.23–3.10 Ga (Barre`re 1967; Potrel et al. 1996). This metamorphism generated new gneissic fabrics with partial melting and emplacement of minor anatectic granites including the Ioulguend Granite (Fig. 2; Potrel et al. 1998). Later granitic magmatism dated at 2726+7 Ma coincided with the emplacement of gabbroic intrusions dated at 2706+54 Ma and about 2740 Ma (Auvray et al. 1992; Potrel et al. 1998). Negative 1Nd values at 2.7 Ga for one of the ‘late’ granites (Potrel et al. 1996, 1998) imply that it was, at least in part, derived from pre-existing crust. Fieldwork undertaken as part of the present study has confirmed the polyphase sequence of metamorphic and intrusive events documented by previous work.

Metamorphic lithologies The characteristic feature of the terrane is the preponderance of high-grade (granulite-facies) metamorphic rocks lacking primary (sedimentary and igneous) textures (Fig. 2). The various gneisses and charnockitic rocks described by previous

35

workers are tectonically interlayered. Migmatitic gneisses underlie a flat regolith surface with a gravel or sand veneer and are the dominant lithology in the south. Here, other lithologies form parallel lenses, layers or sheets within the migmatitic gneisses (Key et al. 2003). Charnockitic rocks and thick, variably garnetiferous quartzofeldspathic gneiss units are the dominant northern lithologies. These lithologies are noticeably absent from the Tasiast–Tijirit Terrane. Thinner bands of other lithologies including quartzites, banded ironstones and calc-silicate rocks form strike-parallel low ridges throughout the Choum–Rag el Abiod Terrane. The migmatitic gneisses have coarse-grained, biotite-bearing, grey tonalitic gneiss palaeosomes cut by several generations of felsic veins including invasive partial melt patches, that form up to 50% of the rock volume (phlebitic to stromatic migmatite textures; Fig. 3a). Petrographically, neosome phases range in composition from monzo- to syenogranite. Garnets commonly form in clusters, and cordierite is locally present as relatively large grains in the gneissic groundmass. The whole-rock chemical compositions of these migmatitic gneisses are similar to those of the migmatitic gneisses from the Tasiast–Tijirit Terrane. The Choum–Rag el Abiod Terrane migmatitic gneisses have SiO2 contents of 63.3–72.0% and total alkali contents of 5.7–9.3%. Tasiast–Tijirit Terrane migmatitic gneisses have more restricted SiO2 contents (72.2–74.7%) and also have a more restricted range of total alkali contents (6.9–8.4%). In general, the Tasiast–Tijirit Terrane migmatitic gneisses appear more fractionated than those of the Choum–Rag el Abiod Terrane (Pitfield et al. 2005). Massive, variably garnetiferous quartzofeldspathic gneisses (previously referred to as ‘leptynites’ by Barre`re 1967) and hypersthene-bearing charnockitic gneisses can occur in stacked sequences of tabular units each up to about 40 m in thickness (see also Barre`re 1967). Gneissosity in the quartzofeldspathic gneisses is defined by millimetre- to centimetre-thick mafic stringers and seams with parallel quartz–garnet bands up to 30 cm thick that may mimic a primary layering. These rocks may also carry sillimanite and hercynite spinel along with the biotite in mafic clots and aggregates. Coarse-grained sillimanite – garnet–cordierite– K-feldspar-gneiss bands are up to several hundreds of metres in thickness and can be traced for several kilometres along strike. Other metasedimentary rocks are uncommon and include quartz-rich lithologies (garnetiferous quartzites and quartz-schists) as well as calc-silicate rocks and banded ironstones. The quartzites appear to form ‘restite’ seams and pods in migmatites. Amphibolites are common as small, elongate pods parallel to gneissosity in surrounding migmatitic gneisses.

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Fig. 2. A simplified geological map of the Choum–Rag el Abiod Terrane showing the distribution of the main lithological components. The TISZ is extrapolated under the Akchaˆr Dune Field.

There are minor meta-igneous intrusions within the migmatitic gneisses that are found only in the Choum–Rag el Abiod Terrane. Numerous isolated small pods of dunite with criss-crossing carbonate vein networks are present along the western margin of the terrane as well as within the sheared part of this terrane along the eastern part of the TISZ. The pods comprise angular pieces of altered dunite (antigorite after olivine with magnetite trails) embedded in a crosscutting network of carbonate veins up to several centimetres in thickness. Individual mounds are up to about 200 m long and 10 m in height. The tectonic significance of these intrusions is not clear. Barre`re (1967) mapped a series of anorthositic sheets that are also not present in the Tasiast –Tijirit Terrane. These sheets are concordant to the regional NE–SW to north–south structural grain of the host gneisses. Single sheets can be traced for several kilometres and are less than 100 m in thickness. The southern exposed half of the Choum–Rag el Abiod Terrane is intensely sheared within the TISZ (Fig. 1). Individual lithological units (composed of all the lithologies described above) in the shears are deformed into lenses from several centimetres to several hundred metres in length, and from less than 1 cm to tens of metres in thickness. The lenses are tectonically juxtaposed in a quartz-mylonite matrix (Fig. 3b). Transposed planar fabrics are dominant, locally with a down-dip lineation.

Post-TISZ c. 2730 Ma plutonic rocks Barre`re (1967) mapped a large granite in the centre –east part of the Atar area that he referred to as the Touijenjert –Modreı¨gue Granite. He recognized a central porphyritic phase (Touijenjert Granite) surrounded by a gneissic biotite phase (Touyerma Granite). A third xenolithic phase was noted during the present fieldwork. The NNW – SSE-trending Touijenjert –Modreı¨gue Granite cuts across major NE –SW-trending shears that branch off the TISZ (Figs 1 and 2) as well as cutting obliquely across the gross layering and tectonic fabric of its metamorphic country rocks. However, the granite is itself cut by a new set of shears with zones of intense deformation (notably in the NW part of the intrusion). Potrel et al. (1998) obtained a SHRIMP U –Pb age of 2726+7 Ma for the Touijenjert Granite that confirms an earlier single zircon age of 2715+ 11 Ma obtained by Auvray et al. (1992). The southernmost part of the Touijenjert – Modreı¨gue Granite comprises a xenolithic, medium-grained, equigranular granite. The characteristic feature of the granite is the presence of angular blocks of different types of country rocks, most notably migmatitic gneisses. The blocks are chaotically organized and vary in size from several centimetres in length to huge blocks that are tens of cubic metres in volume. It is suggested that the granite was emplaced into a


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Fig. 3. (a) Blocks of massive migmatitic tonalite gneisses in the southern part of the Choum– Rag el Abiod Terrane away from the major shear zones at W13.48280 N20.34860. (Note the small basalt dyke to the right of the men.) BGS photo P513290. (b) Lenticular fabric typical of shear zones in the Choum–Rag el Abiod Terrane at W13.38115 N20.24778. BGS photo P513193. (c, d) Polished (thick) section images (crossed polarizers and plane-polarized light, respectively) of biotite –quartz– K-feldspar–garnet –sillimanite–cordierite gneiss (sample 20131250), in which trails and elongate crystals of garnet, sillimanite and biotite define a strong planar fabric. Field of view is 3 mm. BGS photos P1010168 and P1010169, respectively.

pre-existing me´lange within the NE – SW-trending shear zone. Barre`re (1967) also identified and described a major gabbro that underlies a range of large hills at Iguilid. This gabbro has an arcuate NNE – SSW shape with a strike length of about 10 km within the NE – SW-trending shear zone through the middle of the Choum – Rag el Abiod Terrane (Fig. 1). The gabbro is weakly metamorphosed with foliated margins and is now referred to as the Iguilid Metagabbro. Several generations of quartzofeldspathic pegmatites cut the Choum–Rag el Abiod Terrane including late muscovite–tourmaline–garnet –beryl-bearing veins common within the Iguilid Metagabbro. Auvray et al. (1992) dated a granulitic gabbro to the south of the Iguilid body (at Guelb el Azib) at c. 2.74 Ga, which led Potrel et al. (1998) to conclude that

there was a second high-grade metamorphic event in this area.

Tectonothermal events Tectonothermal events recognized in the Choum– Rag el Abiod Terrane are divided into (1) early (pre-TISZ) high-grade events, (2) events synchronous with the development of the TISZ zone and its NE– SW-trending offshoot, and (3) post-TISZ events. An early polyphase tectonothermal history (equivalent of the Precambrian one event of Barre`re 1967) is recorded that produced structures subsequently flattened by shearing associated with the development of the TISZ. Gneissic fabrics are axial planar to tight to isoclinal folds of early vein phases, best seen in the migmatitic orthogneisses.

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The gneissosity is itself tightly folded with the development of an axial planar foliation and intrafolial, rootless folds. Ptygmatic folds and complex interference fold patterns characterize the migmatitic rocks. Larger-scale folds best defined by individual lithologies such as amphibolite mimic the small-scale structures. The early polyphase metamorphic evolution culminated in a granulite-facies event that included charnockites dated at 2986+8 Ma and terminated prior to the emplacement of post-tectonic granites and gabbroic bodies at about 2.73 Ga (Potrel et al. 1998). Granulite-facies mineral assemblages define gneissic fabrics and include various garnet– sillimanite –cordierite –K-feldspar ( –hercynite spinel) assemblages in paragneisses (Fig. 3c and d). The presence of K-feldspar and absence of muscovite suggests that these rocks have experienced significant partial melting. Small pockets of symplectite seen in thin sections of these paragneisses also suggest some incipient melting. The garnet – sillimanite –cordierite assemblage documented in these samples is typically stable at conditions of about 800 8C at pressures of about 6 kbar (e.g. White et al. 2001) as recorded by Potrel et al. (1998) for this region. Lower P–T estimates are based on microprobe data for two samples (one is shown in Fig. 3c and d) of garnet –sillimanite – cordierite gneiss with minor biotite, K-feldspar, plagioclase and green hercynite spinel, from the central part of the Choum –Rag el Abiod Terrane. The microprobe data were run through GPT, an Excel spreadsheet for thermobarometric calculations in metapelitic rocks (Reche & Martinez 1996). The key minerals are not significantly zoned and are apparently well suited to quantitative P –T estimation. GPT allows the simultaneous solution of two two-variable expressions through iteration, removing the need to assume a value for one parameter (i.e. pressure or temperature) to estimate the other. Iterative calculation using the calibration for garnet –biotite equilibrium of Hodges & Spear (1982) with the garnet –plagioclase –biotite –quartz geobarometer of Hoisch (1990) produces temperature estimates of 674 8C and 612 8C, and pressure estimates of about 4.6 and 2.5 kbar, respectively. However, iterative calculation using the calibration for garnet – cordierite equilibrium of Bhattacharya et al. (1988) with the garnet –plagioclase –biotite– quartz geobarometer of Hoisch (1990) produces slightly higher P –T conditions (698 and 705 8C and 5.0 and 3.8 kbar, respectively). Although a substantial margin of error must be attached to the results, they are consistent with mineral equilibration in a high-temperature, low- to intermediatepressure (up to 15 km depth) metamorphic environment. The estimated P –T conditions, although lower than those recorded by Potrel et al. (1998), lie close

to the reaction line for biotite þ sillimanite ¼ garnet þ cordierite þ water (þK-feldspar) in the KFMASH system (Spear & Cheney 1989). Spear (1993) noted that all the mineral phases involved in this reaction are found to coexist, implying that the reaction is divariant. Figure 3c and d shows that sillimanite growth followed cordierite, as individual sillimanite laths cut across cordierite grains (and twinning in cordierite) to define (with biotite) a strong planar fabric. Garnet grains are fragmented and are also locally overgrown by sillimanite needles. Slight pinitization of many cordierite grains indicates some retrogression. The NE–SW-trending shears through the southern exposed half of the terrane are part of the Precambrian 2 event of Barre`re (1967). Shear fabrics trend roughly NE –SW and dip both to the NW and SE (Fig. 4). The sheared rocks have strong planar fabrics in which all fabrics (including veins) are parallel to a new foliation. Existing folds are tightened and refolded, and new structures include transposed gneissosity and ductile, westerly dipping thrusts and subvertical faults. Quartzmylonites commonly infill thrust and fault planes. Individual lithological units were deformed into lenses during this event by anastomosing shears (Fig. 3b). The development of these ductile shears predated the emplacement of the Touijenjert – Modreı¨gue Granite, as this granite truncates major NE– SW-trending shears. A third period of ductile shearing occurred either during emplacement of the Touijenjert –Modreı¨gue Granite or at a later date, as the porphyritic phase of this granite is locally deformed into augen gneiss. New shears formed and individual shears within the NE– SW-trending shear zone of event 2 were reactivated. The growth of amphibole after pyroxene, widespread growth of epidote along joint planes, and alteration of biotite to chlorite, and feldspar to sericite indicate that there was at least one regional retrogressive event after the emplacement of the late granites. Open folding succeeded ductile shearing. Brittle deformation (equivalent of the Precambrian 3 event of Barre`re 1967) produced numerous faults and fractures with NE–SW, east–west and north–south trends. Many of these are infilled by basic dykes and secondary epidote coats joint faces. Repeated brecciation of quartz-mylonites occurred during brittle reactivation of ductile shear zones. All of these ‘late’ tectonic and magmatic events affected the whole of the Reguibat shield area of NW Mauritania.

Tasiast – Tijirit Terrane Previous geological mapping in parts of the Tasiast– Tijirit Terrane, notably by Maurin et al. (1997),


Fig. 4. Lower hemisphere equal area stereonet plots of poles to foliation and gneissosity for the Choum –Rag el Abiod Terrane and eastern and western domains of the Tasiast– Tijirit Terrane. 39

40

R. M. KEY ET AL.

identified migmatitic gneisses and major greenstone belts surrounded by younger, voluminous tonalitic or granodioritic plutons that commonly core domelike structures. The greenstones and surrounding plutons are absent from the Choum –Rag el Abiod Terrane. Conventional Rb/Sr whole-rock and mineral ages of between 2.34 and 2.84 Ga were previously recorded from this terrane (Bonnici & Giraudon 1963; Rocci et al. 1991). Chardon (1997) obtained one zircon evaporation age of c. 2.97 Ga from migmatitic gneisses adjacent to the Chami Greenstone Belt. All the examined gneisses had Nd model ages in a restricted range of 3.05– 3.10 Ga. A similarly determined age of c. 2.93 Ga was obtained from a granodiorite pluton also close to the Chami Greenstone Belt and all the plutons have similar Nd model ages to the examined gneisses. Felsic volcanic rocks from the Chami Greenstone Belt had Nd model ages from 3.05 to 3.60 Ga, suggesting involvement of older crustal material in their genesis. Our mapping (Key 2003; Key & Loughlin 2003) confirmed the earlier work and identified the oldest rocks as the variably migmatized tonalitic gneisses, one phase of which was dated by Chardon (1997) at c. 2.97 Ga. These gneisses are cut by all the other ‘granitic’ phases and tectonically as well as unconformably underlie the greenstone belts. There are a number of small, elongate, predominantly amphibolitic units, interpreted as older greenstone belt remnants, within the tonalitic gneisses. The lithologies that make up the Tasiast –Tijirit Terrane can therefore be divided into the following three major groups (Fig. 5): (1) migmatitic gneisses with amphibolite lenses (oldest unit) and minor gneissic granite; (2) greenstone belt lithologies; (3) post-greenstone belt granitoid intrusions that include major biotite –tonalites or granodiorites (including plutons with abundant secondary epidote), large granites and small pegmatitic muscovite-granites that are the youngest intrusions. The charnockites, massive quartzofeldspathic gneisses and sillimanite –garnet –cordierite – Kfeldspar-gneisses that are major components of the Choum– Rag el Abiod Terrane are noticeable by their absence.

Metamorphic lithologies There are two major textural types of early migmatitic gneiss: phlebitic and stromatic (Mehnert 1968). Migmatitic gneisses with phlebitic textures (Fig. 6a) comprise a grey tonalitic gneiss host cut by at least four generations of felsic leucosome comprising intersecting felsic veins and diffuse (partial melt) zones all cut by minor metamafic

dykes. The various leucosome phases make up over 20% of the rock volume. Thick quartzofeldspathic veins define ptygmatic and convoluted folds and mafic layers form surreitic (dilatational) structures. The migmatitic gneisses preserve a complex tectonothermal history that predates deposition of the volcano-sedimentary greenstone belts sequences. Migmatites with stromatic textures (Fig. 6b) form two linear, north–south belts on either side of the western (Kreidat and Chami) greenstone belts. These gneisses have a steeply dipping north– south gneissosity on a millimetre to centimetre scale accentuated by parallel white felsic veins up to about 30 cm in thickness and that make up between about 10% and 50% of the rock volume. The gneissosity and veins can be followed for tens of metres across large rock pavements parallel to the mean trends of the two outcrop belts. Tight intrafolial folds within the gneissosity show that it is a transposed fabric that effectively obliterated earlier structures. The gneissosity forms an axial planar fabric to cuspate folds defined by discordant fine-grained amphibolite sheets. The migmatitic gneisses tectonically (and locally unconformably for the Ahmeyim Greenstone Belt) underlie the various greenstone lithologies to confirm the schematic cross-sections on the published Chami Sheet (Maurin et al. 1997). Larger migmatitic gneiss exposures have antiformal or dome-like aspects. There are numerous isolated and small amphibolite pods in the migmatitic gneisses. The amphibolites are invaded by neosome phases of the surrounding migmatites and are older than the metamafic volcanic rocks of the overlying greenstone belts. The main greenstone belts in the Tasiast – Tijirit Terrane are named as follows, from east to west (Fig. 1): Tijirit Greenstone Belt; Ahmeyim Greenstone Belt; Sebkhet Nich Greenstone Belt in the SE; Kreidat Greenstone Belt in the north – centre; Chami Greenstone Belt in the centre; two small greenstone belts in the west –centre referred to as the Hudeibt Agheyaˆne and Hadeibt Lebtheinıˆye´ greenstone belts (collectively referred to formerly as the Lebzenia greenstone belts). The characteristic feature of the western greenstone belts is their linear aspect parallel to the north – south fabric of surrounding stromatic migmatitic gneisses. The eastern greenstones trend NNE – SSW parallel to the strike of the TISZ, which defines the eastern margin of the Tasiast – Tijirit Terrane. These linear trends are clearly controlled by major tectonic events described below. All of the greenstone belts


41

Fig. 5. The distribution of igneous rocks within the Tasiast– Tijirit Terrane and adjoining parts of the TISZ.

comprise interlayered mafic and ultramafic metavolcanic rocks with minor uppermost felsic metavolcanic rocks confined to specific eruptive centres (including dated sample D in Fig. 1; Tasiast drillcore sample in Table 1). Chemical analyses of the various greenstone belt metavolcanic rocks (Pitfield et al. 2005) suggest that early mafic and ultramafic rocks were emplaced during initial rifting of continental crust. Siliciclastic metasedimentary rocks as well as prominent banded ironstones (Fig. 6c) are either interbanded with the mafic and ultramafic rocks (e.g. in the Ahmeyim, Hadeibt Agheyaˆne and Hadeibt Lebtheinıˆye´ greenstone belts) or confined within rifts in the centres of the greenstone belts (e.g. in the Chami Greenstone Belt).

Plutonic rocks A sequence of intrusive events is identified, with most pluton-forming magmatism postdating the volcanism and associated sedimentation recorded in the greenstone belts. Pre-greenstone belt plutonism is confined to small, early gneissic granites intruding into migmatites in the western part of the Tasiast –Tijirit Terrane. The planar fabric in the early granites is defined by biotite lenses (c. 6 cm long) and a parallel quartz leaf fabric with flattened feldspars. The commonest intrusions in the Tasiast –Tijirit Terrane are biotite-tonalites (of the Ndaouaˆs Suite), which form a number of large intrusions west of the Ahmeyim Greenstone Belt. The plutons cut across fabrics in the migmatitic gneisses and also intrude

42

R. M. KEY ET AL.

Fig. 6. (a) Migmatitic tonalite gneiss cut by ptygmatically folded granitic veins in the Tasiast– Tijirit Terrane at W14.94257 N20.99558. BGS photo P513260. (b) Stromatic or straight gneisses in the western part of the Tasiast–Tijirit Terrane at W15.61730 N21.00081. BGS photo P522082. (c) Isoclinal folds, locally with sheared-out limbs in BIF of the Hadeibt Lebtheinıˆye´ Greenstone Belt at W15.86750 N20.95300. BGS photo P521983. (d) Float of L-tectonites (talc-phyllites) with characteristic arrowhead shapes from the southern part of the Chami Greenstone Belt at W15.58889 N20.49333. BGS photo P521989.

through folded greenstone belt successions. Small country-rock xenoliths generally form less than 1% of the rock volume but are locally common, notably immediately adjacent to greenstone belts where abundant metamafic volcanic xenoliths are present. The tonalites are locally foliated and gneissic (in discrete shear zones), and are cut by minor granitic vein phases. Abundant secondary epidote as veins and as a pervasive replacement phase (of feldspar and biotite) is present in some of the tonalitic plutons. The largest of these altered plutons occurs at Gleibat el Fhoud, west of the Ahmeyim Greenstone Belt (zircon sample B in Fig. 1, sample 201401584 of Table 1, is from this pluton). It is a medium- to coarse-grained, equigranular leucocratic rock, locally with euhedral feldspar phenocrysts (partly replaced by epidote) that are up to about 5 cm in length. The modal composition is close to or on the tonalite –granodiorite boundary. The

granodioritic to tonalitic plutons of the Ndaouaˆs Suite are calc-alkaline granitoids with SiO2 contents from 65.8 to 77.5% and Na2O þ K2O values in the interval 5.8– 9.1% that range from medium potassic through high potassic to shoshonitic (Pitfield et al. 2005). Less common are granitic intrusions of the Bir Igueni Suite (zircon sample C in Fig. 1, sample 201401598 of Table 1, is from one granitic pluton) found in the centre of the Tasiast –Tijirit Terrane. These plutons have abundant rafts of migmatitic gneiss in white– grey, variably gneissic leucogranite that forms a network of criss-crossing veins. The veins form up to 50% of individual large exposures. A characteristic feature of the veins is the presence of tiny red garnets and large magnetite grains. The gneiss rafts are variably assimilated and up to tens of cubic metres in volume. There is a transitional contact between the granite and its host migmatitic gneiss with

Table 1. U –Pb zircon data for Mauritanian samples Sample name & fraction code

Sample wt (mg)

Cm-Pb (P9)

Concentrations* Pb (ppm)

U (ppm)

Atomic ratios 206

208

207

204

206

206

Pb/ Pb†

Pb/ Pb†

Pb/ Pb‡

2SE%

206

Age (Ma)

Pb/ U‡

2SE%

207

238

Pb/ U‡

235

2SE% Rho

§

207

Pb/ Pb age

2s

206

2.8 1.3 0.6 0.9

0.4 0.0 0.6 3.0

81 137 510 162

172 304 1173 296

23444 377440 22184 2075

0.1906 0.2195 0.1697 0.1989

0.20816 0.20839 0.20738 0.20933

0.094 0.186 0.110 0.112

0.38951 0.36527 0.36610 0.44166

0.329 0.377 0.370 0.409

11.179 10.496 10.468 12.747

0.344 0.423 0.387 0.432

0.96 0.90 0.96 0.97

2891.2 2893.1 2885.2 2900.3

1.5 3.0 1.8 1.8

Sample 201401598 12 10 X 7

0.5 0.6 0.8 1.8

21.2 11.3 4.1 0.5

494 279 174 53

752 397 229 99

512 637 1349 9122

0.1061 0.1536 0.2919 0.1090

0.21245 0.21379 0.21290 0.21084

0.095 0.143 0.144 0.109

0.52872 0.55763 0.56675 0.46578

0.303 0.372 0.831 0.530

15.488 16.437 16.636 13.541

0.320 0.403 0.847 0.547

0.96 0.93 0.99 0.98

2924.3 2934.5 2927.7 2912.0

1.5 2.3 2.3 1.8

Tasiast drill-core; three grains from Z-1 4.7 Z-3 17.1 Z-4 0.2

50 kg of rock 1.3 21 0.5 3 2.0 93

29 4 132

3184 0.2252 0.21753 4733 0.2313 0.21828 397 0.1536 0.20461

0.192 0.125 0.599

0.57991 0.56615 0.54077

0.253 0.376 0.970

17.393 17.039 15.256

0.321 0.399 1.168

0.80 0.95 0.86

2962.5 2968.1 2863.3

3.1 2.0 9.7

Sample 201401619 Z-15 Z-16 Z-17 Z-13

5.3 6.9 7.9 3.3

1104 946 1054 3871

0.099 0.082 0.113 0.065

0.50464 0.55342 0.55382 0.52952

0.381 0.351 0.342 0.225

14.895 16.466 16.455 15.672

0.398 0.363 0.364 0.235

0.97 0.97 0.95 0.96

2936.6 2949.5 2947.3 2941.0

1.6 1.3 1.8 1.0

0.2 0.4 1.0 3.3

645 371 186 84

1077 547 279 136

0.1081 0.1406 0.1228 0.1284

0.21407 0.21579 0.21549 0.21465


Sample 201401584 7 2 1 6

All fractions were nonmagnetic at 208 tilt angle on a Frantz LB-1 separator at 1.7 A. Abbreviations: rbrown, red– brown; pbrown, pink– brown; transl., transluscent; transp., transparent; fr, fragment, subcirc., subcircular; ra, aspect ratio. Standard zircon separation techniques were used including abrasion following Krogh (1982). Samples were spiked with a mixed 205Pb/235U tracer and dissolved in HF–HNO3 using the method of Parrish (1987). Chemical separations followed Krogh (1973) with modifications after Corfu & Noble (1992). All data were obtained by single-collector peak jumping (Noble et al. 1993) on a VG354 mass spectrometer fitted with a WARP filter and Philips ion-counting Daly detector. Ages were calculated by using decay constants of Jaffey et al. (1971). The laboratory blank Pb composition was 206Pb/204Pb ¼ 18.19, 207Pb/204Pb ¼ 15.58, and 208Pb/204Pb ¼ 38.5. Quoted errors are 2s (% for atomic ratios, absolute for ages). *Maximum errors are +20%. Weights were measured on a Cahn C32 microbalance. † Measured ratio corrected for fractionation. ‡ Corrected for fractionation and spike. §207 Pb/235U – 206Pb/238U error correlation coefficient calculated following Ludwig (1980).

43

44

R. M. KEY ET AL.

a gradual decrease in the proportion of the granite vein phase. One such granite intrudes into the northeastern part of the Sebkhet Nich Greenstone Belt. A distinctive, coarse-grained to pegmatitic muscovite-granite forms a series of very small intrusions throughout the western part of the Tasiast –Tijirit Terrane. Its pegmatitic phase (which usually forms sheets and veins cutting all other lithologies of the Tasiast –Tijirit Terrane) is characterized by a range of unusual lithium, strontium and beryllium minerals as well as opaque minerals, tourmaline, biotite and red garnet. Muscovite books in the pegmatites are up to 1 m in length. These micaceous granites may be coeval and comagmatic with the (late) muscovitegranites occurring in the Choum –Rag el Abiod Terrane and in the TISZ.

Tectonothermal events The following sequence of deformation events and associated metamorphism is preserved in the rocks of the Tasiast –Tijirit Terrane. (1) An early set of small-scale folds of gneissosity and ductile shears in the various migmatitic gneisses that predate the deposition of the greenstone belt sediments and volcanic rocks. An initial high-grade event produced gneissic fabrics that are axial planar to folds defined by felsic veins. The gneissosity (as well as felsic veins and dykes) was then sheared and folded into tight to isoclinal folds. Structures include ptygmatic folds of quartzofeldspathic veins and isoclinal folds associated with ductile shears. A lack of indicator minerals precludes accurate assessments of the metamorphic grades of these early events. However, partial melting occurred, as indicated by phlebitic textures (Fig. 6a). Subsequent (post-greenstone belt) tectonic events have strongly modified the original orientation of the early structures. (2) Early tight to isoclinal folding in the greenstone belt lithologies was accompanied by the generation of a penetrative foliation and accompanying lineation (Fig. 6d). Greenschist- to amphibolitefacies metamorphism accompanied the folding. Ductile shearing in the eastern part of the Tasiast –Tijirit Terrane imposed a strong NNE– SSW to NE–SW tectonic grain (Fig. 4). In contrast, the western part of the terrane has a strong north– south tectonic grain imposed by shearing and transposed planar fabrics. The temporal relation between the two sets of shears is not known. (3) The emplacement of biotite-tonalite plutons modified the structure of the greenstone belts and post-dated the tectonic events described above, producing open domical structures cored by the plutons

(e.g. in the western part of the Kreidat Greenstone Belt). Discrete ductile shears locally cut the plutons, notably close to the TISZ. This period(s) of ductile deformation may have been contemporaneous with the late ductile shears seen in the Choum –Rag el Abiod Terrane. Regional retrogression is indicated by the widespread growth of epidote, chlorite and carbonate minerals. Epidote occurs along brittle fractures and replaces mafic minerals in all of the gneissic and granitic rocks. Secondary growth of chlorite and carbonate is widespread in the greenstone belts, notably in metabasalts. (4) Several generations of brittle fractures can be recognized including NE–SW- to NNE – SSW-trending fractures infilled by a subcontinental mafic dyke swarm (which extends eastwards across the TISZ and the Choum –Rag el Abiod Terrane). Late brittle fractures are locally infilled by east – west-trending, brecciated quartz reefs associated with late hydrothermal alteration. Breccias along faults in the greenstone belts locally contain clasts of chlorite schist.

Taˆçaraˆt– Inemmauˆdene Shear Zone (TISZ) Fieldwork identified a major curviplanar, NNE – SSW- to NE– SW-trending ductile shear zone approximately 70 km wide, which separates the Choum –Rag el Abiod and Tasiast– Tijirit terranes and also cuts into the southern exposed part of the Choum –Rag el Abiod Terrane. Anastomosing shears divide the TISZ into five major segments as follows from west to east. (1) There is a foliated granite with mafic xenoliths aligned in its strong planar fabric along the western margin adjacent to the Tijirit Greenstone Belt (Fig. 7a). (2) This is bounded on its eastern side by the Taˆçaraˆt Suite of strongly foliated granites and augen granite gneisses. (3) A central zone is dominated by metamorphic rocks with transposed planar fabrics (Fig. 7b). This elongate segment has a core of steeply dipping, foliated to mylonitic, flaggy, biotite-bearing quartzofeldspathic (tonalitic) gneisses and quartz-mylonites flanked by less steeply dipping sheared lithologies including augen gneisses. The central gneisses have porphyroclastic textures with tiny rounded and rotated quartz grains. Amphibolite lenses in the gneisses range in length from several centimetres to hundreds of metres. Less common are lenses of migmatitic gneiss with (transposed) stromatic textures and isoclinal folds of gneissosity and intrafolial folds within the gneissosity.


45

Fig. 7. (a) Xenolithic granite with mafic xenoliths aligned in the foliation. BGS photo P513309. (b)0 Transposed gneissic fabrics in migmatitic gneisses in the western part of the TISC. BGS photo P513321. (c) SC fabric in augen granite gneiss in the northern part of the shear zone at W13.56990 N21.19645. BGS photo P522133. (d) Close-up of 7(c) to show s clasts of K-feldspar, indicating sinistral offset.

(4) A wedge of foliated granitic rocks including porphyritic phases of the Aoutitilt Suite (dated sample A in Fig. 1; sample 201401619 of Table 1) forms an eastern core to the TISZ. The porphyritic granites contain up to 5% by volume of angular, dark mafic xenoliths up to about 30 cm in length and are chemically distinct from other analysed granites from the adjoining terranes with flat chondrite-normalized REE profiles with a positive Eu anomaly (Pitfield et al. 2005). These granites are mostly strongly foliated (biotite fabric) with euhedral feldspar phenocrysts in a medium- to coarse-grained groundmass of quartz, feldspar and biotite, and less common hornblende. Strongly lineated augen gneisses develop where the deformation is most intense. Quartz lenses help define the gneissosity with lozenge-shaped mafic layers less than 10 cm in thickness in the gneissic

0

fabric. SC fabrics (Passchier & Trouw 1996) are common, with an early foliation cut by shear bands (Fig. 7c). Feldspar augen are rotated and are 1 –7 cm in length, set in a generally coarse-grained matrix. Some augen are deformed veins rather than single feldspar grains. Anastomosing ductile shears locally infilled by pegmatite veins are ubiquitous. Tails to augen (s clasts; Fig. 7d) imply sinistral rotation in the horizontal plane on the eastern side of the ‘flower structure’. A similar sinistral sense of movement is also indicated by0 the arrangement of single shears and by SC fabrics in augen gneisses (Fig. 7c). (5) The eastern part of the TISZ, as well as a NE –SW-trending offshoot, comprises tectonically interleaved rocks derived from the Choum –Rag el Abiod Terrane.0 Small-scale movement indicators (including SC fabrics in augen granite gneiss)

46

R. M. KEY ET AL.

Fig. 8. Cross-section across the southern part of the Taˆçaraˆt –Inemmauˆdene Shear Zone.

confirm sinistral horizontal offset along, and eastdirected thrusting across, the eastern part of this transpressive, fundamental deformation zone. An east –west cross-section across the southern part of the shear zone demonstrates the presence of a ‘flower structure’ (Fig. 8). Further north, foliation planes dip consistently westwards across the TISZ at between about 358 and 808. The grade of the metamorphism that accompanied the shearing is inferred to be in the greenschist to amphibolite facies based on the presence of hornblende in 0 sheared mafic rocks and biotite (C ) fabrics in augen granite gneisses. Epidote coats joint surfaces, notably in granitic rocks.

Geochronology and isotope geochemistry All U –Pb ages reported in this paper are singlezircon analyses conducted by isotope-dilution thermal ionization mass spectrometry (ID TIMS) at the NERC Isotope Geosciences Laboratory, UK (Table 1). Zircon separates were prepared following standard separation techniques including density separation using heavy liquids and magnetic separation to select the least magnetic, least included, non-metamict zircons. The zircons were hand picked under alcohol using a high-quality binocular microscope to select the best grains for analysis, discriminating populations by virtue of morphology, size, visible cores, etc. Selected grains were rigorously abraded using the air abrasion technique of Krogh (1982) to reduce the likelihood of Pb loss and rim phases causing discordance and mixing of age components. U –Pb analyses on single zircons were carried out using a 205Pb/235U mixed spike solution following the procedures of Krogh (1973) and Parrish (1987) with modifications after Corfu & Noble (1992). Analyses were conducted on a VG 354 multi-collector TIMS system

equipped with a WARP filter, axial Daly photomultiplier and Ortec ion counting detection system, or using a Thermo-Finnegan Triton multi-collector TIMS instrument fitted with an axial electron multiplier. Ages and errors were determined using the Isoplot 3 macro of Ludwig (2003) using the uranium decay constants of Jaffey et al. (1971), and common Pb was corrected using a Stacey & Kramers (1975) two-stage model lead evolution curve. The residue remaining after ion exchange separation of U and Pb, was evaporated to incipient dryness and redissolved in 2% HNO3 –0.1M HF ready for Hf isotope analysis by multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS) following a procedure modified after Nowell & Parrish (2000). A ThermoElemental Axiom MC-ICP-MS system was used, coupled to a Cetac Technologies Aridus desolvating nebulizer. Corrections for Lu and Yb isobaric interferences were determined and applied following Nowell & Parrish (2001). Data were normalized to equivalent washes of the 91500 zircon standard prepared at the same time, according to 176Hf/177Hf and 176Lu/177Hf values of 0.282284 and 0.000288 (Wiedenbeck et al. 1995) respectively. Data are presented in Table 2. Sm and Nd concentrations were obtained by isotope dilution using a mixed 149Sm– 150Nd spike solution. Double filament assemblies comprising tantalum (evaporation) and rhenium (ionization) filaments were used to run the samples, with analyses conducted on a Finnigan-MAT 262 multicollector TIMS instrument. Further details of methods have been given by Darbyshire & Sewell (1997). Data are presented in Table 3.

Data interpretation Despite efforts to improve the concordance of the analysed zircon crystals through careful hand picking and air abrasion techniques, the majority of the zircons were poor in quality, which is reflected in the largely discordant datasets displayed below. This is not atypical of Archaean and Proterozoic terranes. At the time of writing a chemical abrasion method (Mattinson 2005) is now employed at NIGL, which helps improves concordance of ancient and high-U zircons. Zircons from a sample of augen granite gneiss (sample 201401619 in Table 1) within the TISZ were red–brown in colour, variably cracked and included, and possessed obvious cores under an optical microscope. Care was taken to avoid visible cores and pick the least included and cracked grains for single-grain U –Pb analysis. Despite this, all analysed fractions exhibited Pb loss (up to 10%) but still formed a reasonable

MC-ICP-MS Hf isotope data were corrected for instrumental mass bias assuming 179Hf/177Hf ratio of 0.7325. Data were normalized to JMC475 assuming 176Hf/177Hf ¼ 0.282160. Data were corrected for 176 Yb and 176Lu isobaric interferences using 173Yb/176Yb ¼ 0.795249 and 176Lu/175Lu ¼ 0.026533. Reproducibility of the 176Hf/177Hf and 176Lu/177Hf ratios for 91500 zircon run over the analytical session were 108 ppm and 5% (2SD), respectively. Depleted mantle model ages (TDM Hf) were calculated according to values of Griffin et al. (2004). 1Hf was calculated relative to CHUR values of Blichert-Toft & Albare`de (1997).

70 67 87 3029 3075 3062 1.9 1.8 2.4 3.5 1.8 3.0 0.280927 0.280892 0.280902 0.0094 0.0088 0.0114 0.280964 0.280930 0.280959 0.0391 0.0434 0.0992 0.028014 0.025146 0.041055 0.1408 0.2735 0.0863 2933 2912 2954 201401598 201401584 201401619

0.000633 0.000655 0.000972

Hf/177Hf 176

+1SE% Yb/177Hf 176

Lu/177Hf +1SE% (normalized to 91500) 176

Age (Ma) Sample

Table 2. MC-ICP-MS Lu– Hf data for Tasiast–Tijirit zircons analysed for U – Pb geochronology

+1SE%

176

Hf/177Hfi

1Hft (abs)

+2s

TDM Hf (Ma)

+ 2s


47

regression with upper intercept of 2954+11 Ma (MSWD ¼ 6.1, Fig. 9a). The collinearity of these four discordant data points argues against the variable inclusion of inherited cores. Instead, a magmatic interpretation of this age provides a maximum age for the ductile shearing within the TISZ. Because all four fractions were interpreted to be of the same age, all four wash fractions were combined for Hf isotope analysis. The data (Hf TDM ¼ 3062+87 Ma and 1Hft ¼ 3.1+2.4 (2s)) suggest derivation from a depleted mantle source with little if any contamination, contrasting with the cored nature of the zircons and the Nd isotope data (Nd TDM ¼ 3204 Ma and 1Ndt ¼ 20.7), which reflect an inherited component. This apparent decoupling of the Hf and Nd isotope data can be explained by the fact that the dominant sink for Hf in a melt is zircon. The refractory nature of zircon means that the Hf signature of a contaminant melt is largely ‘locked’ into its zircons, which may not be resorbed into the host melt, especially if this melt is saturated with respect to Zr. Crystallization of new zircon will therefore reflect the Hf isotope composition and source of the host melt rather than the melt–contaminant mix, whereas the Nd isotope signature will reflect the bulk-rock mixing with the contaminant. As none of the cored zircons were selected for U–Pb/Hf analysis, the two Hf components can be resolved as different from the whole-rock Nd data, and the Hf isotope data are considered to indicate the source of the 2954 Ma melt. The zircons from an epidotized tonalite of the Ndaouaˆs Suite (sample 201401584 in Table 1) were pink to yellow –brown in colour with a prismatic (classic tetragonal) to elongate prismatic habit. No obvious cores or inclusions were apparent. Four U– Pb TIMS data points were strongly discordant (19–31%), forming a discordia with an upper intercept of 2912+35 Ma (MSWD ¼ 10.7, Fig. 9b). The strong discordance of these data means that caution should be exercised in interpreting the upper intercept age and as such it is simply noted that the age of this sample is within error of sample 201401598 (see below). The same similarity exists in the Hf (Hf TDM ¼ 3075 Ma, 1Hft ¼ 1.8+ 1.8 (2s)) and Nd (Nd TDM ¼ 3079 Ma, 1Ndt ¼ 0.6) isotope geochemistry, suggesting a common origin for these two intrusions. A granite pluton from the Bir Igueni Suite of the Tasiast –Tijirit Terrane (sample 201401598 in Table 1) yielded mostly elongate prismatic grains with aspect ratio c. 4:1, pink to brown –red in colour with some smaller, stubbier multi-faceted grains. Milky, high-U overgrowths were apparent on some grains; these were avoided in preference for the clearer, more crystalline pink-coloured grains. Four data points form a discordia with an upper intercept of 2933+16 Ma with

48

R. M. KEY ET AL.

Table 3. Nd isotope data for Tasiast–Tijirit samples Sample

Age (Ma)

Sm (ppm)

Nd (ppm)

2968 2933 2912 2954

1.590 1.604 3.627 4.090

10.20 10.72 26.50 18.94

Tasiast drill-core 201401598 201401584 201401619

147

Sm/144Nd

143

0.0941 0.0904 0.0827 0.1305

Nd/144Nd

0.510757 0.510635 0.510479 0.511310

TDM*

1Ndt

2996 3066 3079 3204

2.5 1.1 0.6 20.7

*Calculated for the age of the rock. Nd errors are 2SD from measured or calculated values. Analytical uncertainties are estimated to be 1.0% for 147Sm/144Nd ratios. Measured 143 Nd/144Nd values were corrected for mass fractionation relative to 146Nd/144Nd ¼ 0.7219. Also shown are the corresponding calculated values for 1Nd together with the depleted mantle Nd model ages (TDM). The latter have been calculated according to DePaolo et al. (1991). 1Nd is calculated relative to a chondritic reservoir with 143Nd/144Nd of 0.512638 and 147Sm/144Nd of 0.1967.

MSWD ¼ 12 (Fig. 9c). The regression has a very high MSWD, indicating excess scatter, but this is pinned relatively near concordia (1–2% discordant) by two data points, lending confidence to the final upper intercept age. All four Hf fractions were

combined to maximize the amount of Hf analysed. The data (Hf TDM ¼ 3029+70 Ma, 1Hft ¼ 3.5+1.9 (2s)) indicate a depleted mantle origin for the melt whereas the Nd isotope data (Nd TDM ¼ 3066 Ma, 1Ndt ¼ 1.1) indicate a small

(b)

(a) 0.57

0.60 2880

2950

0.56

2840

2750

2800

206Pb/238U

206Pb/238U

0.55

2850

0.53

0.52

2650 2550

0.48

2450

0.44 0.51 Intercepts at 325 ± 220 & 2950 ±8.5 [± 11] Ma MSWD = 6.1

Intercepts at 118 ± 200 & 2912 ±34 [± 35] Ma MSWD = 10.7

0.40

data-point error ellipses are 2σ

0.49 14.4 14.8 15.2 15.6 16.0 16.4 16.8

0.36

9

11

207Pb/235U

13

15

17

207Pb/235U

(d)

(c) 0.60

0.60 2900

0.58

0.52

2700

0.48

2940 Z3

0.56

2860

0.54

Z4

Intercepts at 240 ± 300 & 2933 ±14 [± 16] Ma MSWD = 12 vv



0.44 12.5

13.5

14.5

15.5

207Pb/235U

16.5

17.5

Z1

0.52 14.5

15.5

16.5

98001868

2800

206Pb/238U

206Pb/238U

0.56

17.5

207Pb/235U

Fig. 9. (a) U– Pb (TIMS) concordia plot for sample 201401619 (augen granite gneiss). (b) U –Pb (TIMS) concordia plot for sample 201401584 (tonalite). (c) U– Pb (TIMS) concordia plot for sample 201401598 (granite). (d) U– Pb (TIMS) concordia plot for sample 201500738 (felsic metavolcanic rock).


amount of crustal assimilation. This is in agreement with the Nd model age data of Chardon (1997) for a 2.93 Ga granodioritic pluton from the same area. Analysis of zircons from a felsic metavolcanic rock (Tasiast drillcore in Table 1) from the (upper) Aoueóua Formation of the Chami Greenstone Belt in the Tasiast –Tijirit Terrane yielded data from only three small, prismatic, mauvecoloured zircon crystals after processing of more than 50 kg of drill core. These crystals were isolated after removing plentiful pyrite present in the sample. All three data points are discordant (Fig. 9d) with one subconcordant. Because of the lack of collinearity, the best estimate for the age of the rock is given by the 207Pb/206Pb age of the most concordant point. This represents a minimum age for the rock at c. 2965 Ma. The small size of the zircons provided little Hf for isotopic analysis and as such no data are presented for this sample in Table 2. The Nd isotope results (Nd TDM ¼ 2996 Ma; 1Ndt ¼ 2.5) suggest that this volcanic rock was derived directly from a depleted mantle source with little or no crustal contamination. As field and petrographic evidence for the origin of this volcanic rock is ambiguous, the possibility exists that these zircons may all be inherited and that none represents the age of the rock. However, a crosscutting relationship with sample 201401598 (see below) constrains the minimum age of this rock at 2933 Ma. Chardon (1997) also obtained a similar (Pb –Pb) age (c. 2.97 Ga) and Nd isotope data for associated migmatitic gneisses in this area, which unconformably underlie the greenstone belts. The weight of evidence therefore strongly suggests that the interpretation of an eruption age of c. 2965 Ma for this Chami Greenstone Belt volcanic rock is correct. This also provides an upper

49

age limit on the greenstone belt volcanism and associated sedimentation.

Discussion and conclusion The Reguibat shield in NW Mauritania comprises two terranes that are lithologically different and preserve different Mesoarchaean geological histories indicative of different geotectonic settings (Table 4). Both terranes have migmatitic orthogneisses as their oldest components. These gneisses have previously been dated at 3.5 –3.45 Ga in the Choum –Rag el Abiod Terrane, with a single zircon evaporation age of 2.97 Ga for migmatitic gneisses of the Tasiast –Tijirit Terrane with Nd model ages of 3.05–3.10 Ga. However, there is evidence for a c. 3.6 Ga crustal component based on a Nd model age signature for the dated felsic volcanic rock from the Chami Greenstone Belt emplaced through migmatitic gneisses. A siliciclastic supracrustal sequence of quartzofeldspathic gneisses that includes 2.99 Ga charnockite sheets dominates the Mesoarchaean geology of the Choum– Rag el Abiod Terrane. In contrast, a classic c. 2.93 – 2.97 Ga granite–greenstone belt sequence characterizes the Mesoarchaean geological record of the Tasiast –Tijirit Terrane. In this respect the Reguibat shield of NW Mauritania resembles the Superior province of the Canadian shield, where Archaean subprovinces and terranes characterized by granite –greenstone assemblages are tectonically juxtaposed with Archaean metasedimentary subprovinces and terranes (Card & Ciesielski 1986; Davis et al. 2005). The granite –greenstone belt sequences of the Tasiast –Tijirit Terrane preserve voluminous calc-alkaline magmatism sourced from a depleted

Table 4. A summary of the Archaean development of the Choum – Rag el Abiod and Tasiast– Tijirit terranes Reguibat shield 2.73 Ga granitoids and gabbros with crustal contamination Development of the Taˆçaraˆt–Inemmauˆdene Shear Zone with transpressive ductile shearing dated at between c. 2.73 and 2.95 Ga Choum –Rag el Abiod Terrane

Tasiast– Tijirit Terrane 2.93– 2.97 Ga granite – greenstones with .3.2 Ga (up to 3.6 Ga) contamination intruding orthogneisses

2.99 Ga charnockite sheets (with .3.2 Ga contamination or origin) within a siliciclastic paragneiss sequence c. 3.5 Ga migmatitic orthogneisses with interlayered amphibolites and metasedimentary lithologies

Evidence for c. 3.6 Ga crustal component based on Nd model age signature in Chami greenstones; .2.97 Ga migmatitic orthogneisses with .3.1 Ga origin and with amphibolite rafts

50

R. M. KEY ET AL.

mantle with minor crustal contamination and rift-related sedimentation. The Mesoarchaean siliciclastic supracrustal rocks of the Choum–Rag el Abiod Terrane may originally have been deposited as passive margin (accretionary prism) sediments and as such would support an interpretation that the two terranes were allochthonous to each other prior to their amalgamation along the TISZ. Granulite-facies Mesoarchaean paragneisses are exposed in the Choum –Rag el Abiod Terrane whereas the exposed Mesoarchaean rocks of the Tasiast –Tijirit Terrane are of greenschist-facies metamorphic grade. These two assemblages represent different (lower v. mid- to upper) crustal levels. However, the strongly contrasting Mesoarchaean geotectonic setting of the two terranes precludes an interpretation that they represent delamination between lower and middle crustal levels of a single terrane with the Taˆçaraˆt–Inemmauˆdene Shear Zone (TISZ) as an intra-terrane structure. The measured transpressive character of the TISZ reflects post-amalgamation movement and in this respect is a typical post-collisional transpressive shear zone (Lie´geois et al. 1998). Cross-sections of the TISZ show an east– west ‘flower structure’ across its southern part with sinistral horizontal offset and east-directed thrusting on its eastern side. Further north, shear-related fabrics dip consistently westwards, indicating that the Tasiast –Tijirit Terrane lies tectonically above the Choum –Rag el Abiod Terrane. Widespread Neoarchaean granitic (and less common basic) magmatism throughout the Reguibat shield followed transpressive movement along the TISZ. The c. 2730 Ma age for the post-TISZ Touijenjert–Modreı¨gue Granite (Potrel et al. 1998) of the Choum –Rag el Abiod Terrane provides a minimum age for the shearing. A maximum age for the shearing is provided by the 2954+11 Ma magmatic age for an Aoutitilt Suite granite from the eastern core of the TISZ. The original size of the two terranes is not known. The present western limit of the Tasiast –Tijirit Terrane is defined by the Pan-African (late Proterozoic –early Phanerozoic) NW Mauritanides, which lie about 230 km west of the TISZ. Only about 50–70 km of the Choum –Rag el Abiod Terrane is exposed east of the TISZ before it is concealed beneath unconformably overlying strata of the Taoudeni Basin. Elsewhere on the African plate, Archaean terranes that are up to about 800 km by 200 km in area are recognized within the Kaapvaal craton (Eglington & Armstrong 2004; de Wit & Tinker 2004, and references therein). The work described in this paper formed part of the World Bank-funded PRISM Project of the Government of Mauritania and the authors wish to thank Mr Samory

Ould Souedatt, the Project Coordinator, for all his logistical support. Several referees are thanked for their detailed comments on earlier drafts of this paper. This paper is published by permission of the Executive Director, British Geological Survey, Natural Environmental Research Council (NERC) and is NIGL Publication 670.

References A UVRAY , B., P EUCAT , J. J., P OTREL , A., B URG , J. P., D ARS , R. & L O , K. 1992. Donneés geóchronologiques nouvelles sur l’Archeén de l’Amsaga (dorsale Reguibat, Mauritanie). Comptes Rendus de l’Academie des Sciences, 315, 63–70. B ARRE` RE , J. 1967. Le groupe precambrien de l’Amsaga entre Atar et Akjoujt (Mauritanie). Etude d’un me´tamorphisme profond et de ses relations avec la migmatisation. Me´moires du BRGM, 42. B ESSOLES , B. 1977. Geólogie de l’Afrique. Le craton Ouest Africain. Me´moires du BRGM, 88. B HATTACHARYA , A., M AZUMDAR , A. C. & S EN , A. C. 1988. Fe–Mg mixing in cordierite; constraints from natural data and implications for cordierite– garnet geothermometry in granulites. American Mineralogist, 73, 338–344. B LICHERT -T OFT , J. & A LBARE` DE , F. 1997. The Lu–Hf isotope geochemistry of chondrites and the evolution of the mantle–crust system. Earth and Planetary Science Letters, 148, 243–258. B ONNICI , J. P. & G IRAUDON , R. 1963. Le groupe du Tasiast, nouvelle unite´ lithostratigraphique du socle antećambrien de la Mauritanie occidentale. Bulletin de la Socie´te´ Geólogique de France, 5, 1118–1123. C AHEN , L., S NELLING , N. J., D ELHAL , J., V AIL , J. R., B ONHOMME , M. & L EDENT , D. 1984. The Geochronology and Evolution of Africa. Clarendon Press, Oxford. C ARD , K. D. & C IESIELSKI , A. 1986. DNAG#1 Subdivisions of the Superior province of the Canadian shield. Geosciences Canada, 13, 5– 13. C HARDON , D. 1997. Les de´formations continentales archeénnes. Exemples naturels et mode`lisation thermomećanique. Me´moires de Geósciences, Rennes, 76. C ORFU , F. & N OBLE , S. R. 1992. Genesis of the southern Abitibi greenstone belt, Superior Province, Canada: Evidence from zircon Hf-isotope analyses using a single filament technique. Geochimica et Cosmochimica Acta, 56, 2081–2097. D ARBYSHIRE , D. P. F. & S EWELL , R. J. 1997. Nd and Sr isotope geochemistry of plutonic rocks from Hong Kong: implications for granite petrogenesis, regional structure and crustal evolution. Chemical Geology, 143, 81–93. D AVIS , D. W., A MELIN , Y., N OWELL , G. M. & P ARRISH , R. R. 2005. Hf isotopes in zircon from the western Superior province, Canada: Implications for Archean crustal development and evolution of the depleted mantle reservoir. Precambrian Research, 140, 132–156. D E P AOLO , D. J., L INN , A. M. & S CHUBERT , G. 1991. The continental crustal age distribution: methods of determining mantle separation ages from Sm–Nd

MESOARCHAEAN SHIELDS, NW MAURITANIA isotopic data and application to the southwestern United States. Journal of Geophysical Research (Solid Earth and Planets), 96, 2071–2088. D E W IT , M. & T INKER , J. 2004. Crustal structures across the central Kaapvaal Craton from deep-seismic reflection data. South African Journal of Geology, 107, 185–206. D ILLON , W. P. & S OUGY , J. M. A. 1974. Geology of west Africa and Canary and Cape Verde Islands. In: N AIRN , A. E. M. & S TEHL , F. G. (eds) The Ocean Basins and Margins, Vol. 2, The North Atlantic. Plenum, New York, 315–390. E GLINGTON , B. M. & A RMSTRONG , R. A. 2004. The Kaapvaal Craton and adjacent orogens, southern Africa: a geochronological database and overview of the geological development of the craton. South African Journal of Geology, 107, 13– 32. G RIFFIN , W. L., B ELOUSOVA , E. A., S HEE , S. R., P EARSON , N. J. & O’ REILLY , S. Y. 2004. Archaean crustal evolution in the northern Yilgarn Craton: U–Pb and Hf isotope evidence from detrital zircons. Precambrian Research, 131, 231– 282. H ODGES , K. V. & S PEAR , F. S. 1982. Geothermometry, geobarometry and the Al2SiO5 triple point at Mt. Moosilake, New Hampshire. American Mineralogist, 67, 1118– 1134. H OISCH , T. D. 1990. Empirical calibration of six geobarometers for the mineral assemblage quartz þ muscovite þ biotite þ plagioclase þ garnet. Contributions to Mineralogy and Petrology, 104, 225–234. J AFFEY , A. H., F LYNN , K. F., G LENDENIN , L. E., B ENTLEY , W. C. & E SSLING , A. M. 1971. Precision measurements of half-lives and specific activities of 235 U and 238U. Physics Reviews, C4, 1889–1906. K EY , R. M. 2003. 1:200 000 geological map of the Chami Sheet (2015). Government of Mauritania, Nouakchott. K EY , R. M. & L OUGHLIN , S. C. 2003. 1:200 000 geological map of the Ahmeyim Sheet (2014). Government of Mauritania, Nouakchott. K EY , R. M., L OUGHLIN , S.C. & W ATERS , C. N. 2003. 1:200 000 geological map of the Atar Sheet (2013). Government of Mauritania, Nouakchott. K ROGH , T. E. 1973. A low contamination method for the hydrothermal decomposition of zircon and extraction of U and Pb for isotopic age determinations. Geochimica et Cosmochimica Acta, 37, 485–94. K ROGH , T. E. 1982. Improved accuracy of U– Pb zircon ages by the creation of more concordant systems using an air-abrasion technique. Geochimica et Cosmochimica Acta, 46, 637–649. L IE´ GEOIS , J.-P., N AVEZ , J., H ERTOGEN , J. & B LACK , R. 1998. Contrasting origin of post-collisional high-K calc-alkaline and shoshonitic versus alkaline and peralkaline granitoids. The use of sliding normalization. Lithos, 45, 1– 28. L UDWIG , K. R. 1980. Calculation of uncertainties of U–Pb isotope data. Earth and Planetary Science Letters, 46, 212 –220. L UDWIG , K. R. 2003. User’s manual for Isoplot 3.00: a geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication No. 4. M ATTINSON , J. M. 2005. Zircon U– Pb chemical-abrasion (CA-TIMS) method: combined annealing and

51

multi-step dissolution. Chemical Geology, 220, 47–66. M AURIN , G., B RONNER , G., L E G OFF , E. & C HARDON , D. 1997. Notice explicative de la carte geólogique au 1/200 000 de la feuille Chami (Mauritanie) – Prospection aurife`re dans le Tasiast–Tijirit. Rapport, BRGM, 2459. M EHNERT , K. R. 1968. Migmatites and the Origin of Granitic Rocks. Elsevier, Amsterdam. M ENCHIKOFF , N. 1949. Quelques traits de l’histoire geólogique du Sahara occidental. Livre jubilaire Charles Jacob, Annales Heˆbat et Haug, 7, 303– 325. N OBLE , S. R., T UCKER , R. D. & P HARAOH , T. C. 1993. Lower Palaeozoic and Precambrian igneous rocks from eastern England, and their bearing on late Ordovician closure of the Tornquist Sea: constraints from U– Pb and Nd isotopes. Geological Magazine, 130, 835– 846. N OWELL , G. M. & P ARRISH , R. R. 2001. Simultaneous acquisition of isotope compositions and parent/ daughter ratios by non-isotope dilution solutionmode plasma ionisation multi-collector mass spectrometry (PIMMS). In: H OLLAND , G. & T ANNER , S. D. (eds) Plasma Source Mass Spectrometry: The New Millennium. Royal Society of Chemistry, Gateshead, Special Publication, 267, 298–310. P ARRISH , R. R. 1987. An improved micro-capsule for zircon dissolution in U– Pb geochronology. Chemical Geology (Isotope Geoscience Section), 66, 99–102. P ASSCHIER , C. W. & T ROUW , R. A. J. 1996. Microtectonics. Springer, Berlin. P ITFIELD , P. E. J., K EY , R. M., W ATERS , C. N., H AWKINS , M. P. H., S CHOFIELD , D. I., L OUGHLIN , S. C. & B ARNES , R. P. 2005. Notice explicative des cartes geólogiques et gıˆtologiques au 1/200 000 et 1/500 000 du Sud de la Mauritanie. DMG, Ministe`re des Mines et de l’Industrie, Nouakchott, Mauritania. P OTREL , A. 1994. Evolution tectono-me´tamorphique d’un segment de crouˆte continentale archeénne. Exemple de l’Amsaga (R.I. Mauritanie), Dorsale Re´guibat (Craton Ouest Africain). Me´moire de Geósciences Rennes, 56. P OTREL , A., P EUCAT , J. J., F ANNING , C. M., A UVRAY , B., B URG , J. P. & C ARUBA , C. 1996. 3.5 Ga old terranes in the West African Craton, Mauritania. Journal of the Geological Society, London, 153, 507– 510. P OTREL , A., P EUCAT , J. J. & F ANNING , C. M. 1998. Archean crustal evolution of the West African Craton; example of the Amsaga area (Reguibat Rise); U –Pb and Sm–Nd evidence for crustal growth and recycling. Precambrian Research, 90, 107– 117. R ECHE , J. & M ARTINEZ , F. J. 1996. GPT; an EXCEL spreadsheet for thermobarometric calculations in metapelitic rocks. Computers and Geosciences, 22, 775– 784. R OCCI , G., B RONNER , G. & D ESCHAMPS , M. 1991. Crystalline basement of the West African Craton. In: D ALLMEYER , R. D. & L E´ CORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 31– 61.

52

R. M. KEY ET AL.

S CHOFIELD , D., H ORSTWOOD , M. S. A., P ITFIELD , P. E. J., C ROWLEY , Q. G., W ILKINSON , A. F. & S IDATY , H. C. O. 2006. Timing and kinematics of Eburnean tectonics in the central Reguibat Shield, Mauritania. Journal of the Geological Society, London, 163, 549– 560. S PEAR , F. S. 1993. Metamorphic Phase Equilibria and Pressure–Temperature– Time Paths. Mineralogical Society of America, Reviews in Mineralogy. S PEAR , F. S. & C HENEY , J. T. 1989. A petrogenetic grid for pelitic schists in the system SiO2 –Al2O3 –FeO– MgO– K2O–H2O. Contributions to Mineralogy and Petrology, 101, 149– 164.

S TACEY , J. S. & K RAMERS , J. D. 1975. Approximation of terrestrial lead isotope evolution by a two-stage model. Earth and Planetary Science Letters, 26, 207–221. W IEDENBECK , M., A LLE´ , P., C ORFU , F. ET AL . 1995. Three natural zircons standards for U–Th– Pb, Lu–Hf, trace element and REE analyses. Geostandards Newsletter, 19, 1– 23. W HITE , R. W., P OWELL , R. & H OLLAND , T. J. B. 2001. Calculation of partial melting equilibria in the system Na2O – CaO – K2O – FeO – MgO – Al2O3 – SiO2 – H2O (NCKFMASH). Journal of Metamorphic Geology, 19, 139–153.

Geological setting of the Guelb Moghrein Fe oxide –Cu – Au– Co mineralization, Akjoujt area, Mauritania JOCHEN KOLB1,5, F. MICHAEL MEYER1, TORSTEN VENNEMANN2, RADEGUND HOFFBAUER3, AXEL GERDES4 & GREGORI A. SAKELLARIS1 1

Institute of Mineralogy and Economic Geology, RWTH Aachen University, D-52056 Aachen, Germany

2

Institute of Mineralogy and Geochemistry, University of Lausanne Anthropole, CH-1015 Lausanne, Switzerland 3

Institute for Mineralogy and Petrology, Bonn University, Poppelsdorfer Schloss, 53115 Bonn, Germany

4

Institute of Geosciences, Johann Wolfgang Goethe-University, Altenho¨ferallee 1, D-60438 Frankfurt am Main, Germany

5

Present address: Geological Survey of Denmark and Greenland, Department of Economic Geology, Øster Voldgade 10, DK-1350 Copenhagen, Denmark (e-mail: [email protected]) Abstract: The Guelb Moghrein Fe oxide– Cu–Au–Co deposit is located at the western boundary of the West African craton in NW Mauritania. The wall rocks to the mineralization represent a meta-volcanosedimentary succession typical of Archaean greenstone belts. Two types of metavolcanic rocks are distinguished: (1) volcanoclastic rocks of rhyodacite–dacite composition (Sainte Barbe volcanic unit), which form the stratigraphic base; (2) tholeiitic andesites– basalts (Akjoujt meta-basalt unit). The trace element signature of both types is characteristic of a volcanic arc setting. A small meta-pelitic division belongs to the Sainte Barbe volcanic unit. A metacarbonate body, which contains the mineralization, forms a tectonic lens in the Akjoujt metabasalt unit. It can be defined by the high XMg (¼36) of Fe– Mg carbonate, the REE pattern and the d13C values of 218 to 217‰ as a marine precipitate similar to Archaean banded iron formation (BIF). Additionally, small slices of Fe– Mg clinoamphibole–chlorite schist in the metacarbonate show characteristics of marine shale. This assemblage, therefore, does not represent an alteration product, but represents an iron formation unit deposited on a continental shelf, which probably belongs to the Lembeitih Formation. The hydrothermal mineralization at 2492 Ma was contemporaneous with regional D2 thrusting of the Sainte Barbe volcanic unit and imbrications of the meta-carbonate in the upper greenschist facies. This resulted in the formation of an ore breccia in the meta-carbonate, which is enriched in Fe, Ni, Co, Cu, Bi, Mo, As and Au. Massive sulphide ore breccia contains up to 20 wt% Cu. The ore fluid was aqueous–carbonic in nature and either changed its composition from a Mg-rich oxidizing to an Fe-rich reducing fluid or the two fluid types mixed at the trap site. All lithologies at Guelb Moghrein were deformed by D3 thrusting to the east in the lower greenschist facies. The mobility of REE in the retrogressed rocks explains the formation of a second generation of hydrothermal monazite, which was dated at c. 1742 Ma. Archaean rocks of the West African craton extend to the west to Guelb Moghrein. The active continental margin was deformed and mineralized in the Late Archaean–Early Proterozoic and again reactivated in the Mid-Proterozoic and Westphalian, showing that the western boundary of the craton was reactivated several times.

The Guelb Moghrein Fe oxide–Cu–Au–Co (IOCG) deposit with a total resource of 23.7 Mt at 1.88% Cu and 1.41 gt21 Au is situated near the town of Akjoujt about 250 km NE of Nouakchott, the capital of Mauritania. Commercial production in an open pit operation commenced in October 2006, with a projected annual production of c. 30 000 t of copper and 70 000 ounces of gold (Anonymous 2006). The Guelb Moghrein IOCG mineralization was previously considered to have formed as integral

part of the Late Proterozoic evolution of the Mauritanides (Pouclet et al. 1987; Clauer et al. 1991; Strickland & Martyn 2002; Martyn & Strickland 2004; Kolb et al. 2006). This view, however, is not supported by new chronological data, which point to an Archaean age of the host lithologies and a Late Archaean–Early Proterozoic age of the mineralization (Meyer et al. 2006). Most deposits of the IOCG family (Australia: Olympic Dam, Tennant Creek, Mt. Isa; Sweden: northern Baltic;


54

J. KOLB ET AL.

Canada: Great Bear; Brazil: Salobo) are also of Proterozoic age and have formed in a broad range of crustal and tectonic environments displaying a great variety of structural and host rock controls and styles of mineralization (Hitzman 2000; Partington & Williams 2000). Current genetic models for Mesozoic to Cenozoic IOCG-type mineralization suggest that the major fluid and metal source lies in relatively primitive intrusive rocks, which can be as deep as 10 km and, thus, may not be exposed (Sillitoe 2003). The primary migration paths of the ore fluids are extensional to transpressional major fault zones, which are related to subduction (slab roll-back), anorogenic magmatism and orogenic collapse (Sillitoe 2003). Therefore, the study of the genesis of the Guelb Moghrein deposit is regarded as important for information about the geological evolution and will aid in distinguishing between tectonic models for the Late Archaean–Early Proterozoic of the western margin of the West African craton. In this paper, we present the first geochemical and isotopic dataset for host rocks and mineralized lithologies of the Guelb Moghrein deposit. These data allow a detailed discussion of the tectonomagmatic setting of the country rocks, the nature of the hydrothermal IOCG mineralization and the tectonometamorphic evolution of the region. It will be shown that the deposit is hosted by a suite of rocks that resemble Archaean greenstone belt successions in a volcanic arc and marine continental shelf setting.

Regional geology The Guelb Moghrein deposit in the Akjoujt area lies in the Mauritanides belt, which is generally regarded as a pile of allochthonous terranes thrust eastward towards the West African craton (Fig. 1; Lećorche´ et al. 1989; Villeneuve 2005). About 30 km NE of Akjoujt, a thrust zone marks the boundary to the gneisses and granulites of the Amsaga Basement forming part of the Archaean Reguibat shield and the Taoudeńi basin, which is characterized by rocks of Neoproterozoic to Devonian age. Radiometric dates for the supracrustal rocks in the Akjoujt area and the central part of the Mauritanides are scarce. Thus, their secular evolution was previously inferred from geochronological work carried out in the southern Mauritanides by 40Ar – 39Ar and K –Ar dating of muscovite and amphibole (Ponsard et al. 1988; Dallmeyer & Lećorche´ 1989; Lećorche´ et al. 1989). These data indicate three significant events: (1) at 680 –620 Ma, referred to as Pan-African I orogeny; (2) at 600– 550 Ma, the Pan-African II orogeny; (3) at 330 –270 Ma, related to the Variscan

orogeny (Villeneuve 2005). However, K –Ar and Ar – 39Ar muscovite dating by Clauer et al. (1991) near Akjoujt yielded ages between 315 and 305 Ma and revealed that the supracrustal rocks in this area show no record of any Pan-African ages. Furthermore, laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) U –Pb dating of hydrothermal monazite and xenotime from the Guelb Moghrein deposit points to two hydrothermal events: (1) at 2492 + 9 Ma, signifying the main stage of the hydrothermal IOCG mineralization; (2) at 1742 + 12 Ma, marking the age of lower greenschist-facies retrogression (Meyer et al. 2006). The Akjoujt area was affected by two tectonometamorphic events in the Precambrian and was emplaced to the current position at the western edge of the West African craton during the Variscan orogeny. Whether the Akjoujt terrane shares a common geological evolution with the West African craton or whether it represents an exotic terrane accreted at about 300 Ma remains unresolved at this stage (Meyer et al. 2006). 40

Stratigraphy The Akjoujt area features a supracrustal stratigraphy that is currently believed to consist of two distinct lithological groups separated by an unconformity (Fig. 1). The lower Eizzene Group consists of basalt flows, the Raoui meta-basalt unit, overlain by fine-grained clastic rocks and quartz–magnetite banded iron formation (BIF) of the Khmeiyat Formation. Plutonic activity is indicated by small plagiogranite stocks (Strickland & Martyn 2002; Martyn & Strickland 2004). The lower Eizzene Group is overlain unconformably by a volcanic and clastic succession, named the Oumachoue¿ma Group. This lithological pile has a basal quartzite, the Atilis quartzite, followed by fine-grained meta-greywacke and siltstone of the Irarche`ne el Hamra Formation. This is followed by fine-grained intermediate to mafic volcaniclastic rocks including BIF units up to tens of metres thick, the Atomai Formation. A massive BIF at the top of the Atomai Formation is capped by increasingly felsic lavas and volcaniclastic rocks of andesite to rhyodacite composition, the Sainte Barbe volcanic unit. The volcanic rocks are capped by a widespread BIF–chert marker, the Lembeitih Formation, which, in turn, is overlain by thick and extensive basalt flows with dolerite intrusive rocks of the Akjoujt meta-basalt unit (Strickland & Martyn 2002; Martyn & Strickland 2004).

Structure and metamorphism The Akjoujt area is characterized by a complex set of folded and stacked thin-skinned thrust sheets

GUELB MOGHREIN MINERALIZATION

55

Fig. 1. Schematic geological map of the Akjoujt area (modified after Martyn & Strickland 2004).

(Fig. 1). Five deformation events (D1 –D5) are distinguished (Pouclet et al. 1987; Martyn & Strickland 2004; Kolb et al. 2006). The D1 deformation formed open folds and a weak regional S1 foliation. Thrusting to the NNW created a layerparallel S2 foliation, which formed a crenulation cleavage with the S1 foliation and recumbent folds during D2. A mylonitic S2 foliation is observed in numerous D2 shear zones at all scales. At this stage, the Sainte Barbe volcanic unit was thrust onto the Akjoujt meta-basalt unit at Guelb Moghrein (Figs 1 and 2a). The meta-carbonate body is enveloped by two, up to 40 m wide D2 shear zones in the hanging wall and footwall, respectively (Fig. 2a). Almost orthogonal fabrics were created during the D3 deformation event in all units (Fig. 2a). D3 shear zones are characterized by a closely spaced S3 foliation, which formed during ENE-directed thrusting. Outside the tens of metre-scale shear zones, a crenulation cleavage is

frequently developed in the rocks. Locally, the S2 foliation is folded into upright, open F3 folds with almost horizontal, north–south-trending fold axes (Kolb et al. 2006). Gentle to moderate folds with ENE –WSW-trending fold axes deformed the thrust sheets during D4 and D5, which is correlated with thrusting along the sole thrust during the Westphalian (Fig. 1; Martyn & Strickland 2004). In the Guelb Moghrein deposit, D4 created a conjugate set of S4 foliation and D5 is characterized by a set of NNE–SSW-trending faults with minor offset (Fig. 2a; Kolb et al. 2006). A peak metamorphic, amphibolite-facies hornblende –plagioclase paragenesis is developed parallel to the S1 foliation in the rocks of the Akjoujt meta-basalt unit. Hornblende–plagioclase thermometry indicates 580 + 40 8C for this metamorphic stage (Fig. 2b; Kolb et al. 2006). In contrast, a peak metamorphic, upper greenschistfacies garnet–biotite paragenesis formed parallel

56

J. KOLB ET AL.

Fig. 2. (a) Schematic geological map of the Guelb Moghrein open pit showing the major structural features and the alteration halo surrounding the meta-carbonate, which hosts the IOCG mineralization (modified after Kolb et al. 2006). The almost orthogonal fabrics of the D2 and D3 deformation stages should be noted. (b) P– T– t– D path for the Sainte Barbe volcanic and the Akjoujt meta-basalt units.

to the S2 foliation in the rocks of the Sainte Barbe volcanic unit. Garnet –biotite geothermobarometry records 410 + 30 8C and 2–3 kbar for this metamorphic stage (Fig. 2b; Kolb et al. 2006). The rocks of the Akjoujt meta-basalt unit were retrogressed at this metamorphic stage and a biotite –actinolite paragenesis formed parallel to the S2 foliation. The D2 thrusting assembled the Sainte Barbe volcanic and the Akjoujt meta-basalt units during this second regional metamorphic stage, which was retrograde in the Akjoujt meta-basalt unit, but peak metamorphic in the Sainte Barbe volcanic unit (Fig. 2b; Kolb et al. 2006). We interpret the supracrustal suite in the Akjoujt area as a volcano-sedimentary succession of Archaean age that underwent deformation and peak metamorphism under amphibolite-facies conditions prior to hydrothermal mineralization and retrograde upper greenschist-facies metamorphism (D2/M2) at 2492 Ma. A third deformation event coupled with hydrothermal fluid flow occurred at 1742 Ma in the lower greenschist facies (D3/M3). Final emplacement at the current position on the West African craton by thrusting took place at c. 300 Ma as a result of the collision of Gondwana and Laurentia (Meyer et al. 2006).

Hydrothermal IOCG mineralization Coarse-grained carbonate bodies include the IOCG mineralization at Guelb Moghrein and two prospects, namely Masse and El Joul (Figs 1 and 2a; Strickland & Martyn 2002; Kolb et al. 2006). The Guelb Moghrein ore bodies, Occidental and Oriental, crop out as two abrupt hills west of the Akjoujt town site (Fig. 1; Strickland & Martyn 2002; Kolb et al. 2006). They are 250–500 m in strike length and about 150 m thick. Copper- and gold-rich zones are structurally controlled and interpreted to occur as multiple, about 30 m wide, coalescing lenses that are broadly elongate in the direction of discrete D2 shear zones (Fig. 2a; Strickland & Martyn 2002; Kolb et al. 2006). These mineralized shear zones are developed at the hanging wall and footwall contacts of the carbonate bodies with the surrounding rocks of the Akjoujt meta-basalt unit (Fig. 2a). Furthermore, mineralized shear zones crosscut the carbonate bodies and are controlled by Fe –Mg clinoamphibole–chlorite schist layers within the carbonate bodies (Kolb et al. 2006). Similar mineralization textures are found in Masse and El Joul, which are located in a similar structural and lithological setting (Fig. 2a).


The origin of the carbonate bodies is strongly debated in the literature: Ba Gatta (1982) suggested a synvolcanic origin, based on the nature of the surrounding rocks and the metal content of Guelb Moghrein, which is similar to volcanogenic massive sulphide (VMS) deposits. A synsedimentary origin of the carbonate bodies, precipitated as chemical sediments in a volcano-sedimentary basin, was proposed by Pouclet et al. (1987). Similarly, Kolb et al. (2006) and Meyer et al. (2006) interpreted the carbonates as meta-sediments, based on the following fact: (1) the carbonate is replaced by the hydrothermal mineralization; (2) the Fe –Mg carbonate minerals are deformed by the D2 deformation stage; (3) the carbonates are closely associated with Fe–Mg clinoamphibole– chlorite schists, which represent Fe-rich metasediments. In contrast, a metasomatic origin was suggested by Strickland & Martyn (2002) and Martyn & Strickland (2004), based on the association of the carbonate bodies with carbonate altered volcanic rocks and hydrothermally altered thrust zones and folds. El Khader, Sainte Barbe and Tabrinkout represent smaller IOCG occurrences in the area (Fig. 1; Strickland & Martyn 2002). The mineralization in these prospects is, however, associated with extensive Fe–Mg–Ca carbonate and quartz vein systems and complex stockwork zones (Strickland & Martyn 2002). These fracture-hosted mineralizations are interpreted to be related to later deformation events than the D2 deformation associated with the mineralization at Guelb Moghrein (Strickland & Martyn 2002; Martyn & Strickland 2004).

Analytical procedure Bulk-rock major and trace element analyses were performed on powdered wall rocks as well as mineralized samples by X-ray fluorescence analysis (Institute of Mineralogy and Economic Geology, RWTH Aachen University), instrumental neutron activation analysis and ICP-MS (Activation Laboratories Ltd., Canada). The different techniques were used to obtain a dataset of main and trace elements with low detection limits for the REE and Cu, Au and Co. The mineral chemistry of Fe–Mg carbonate was determined using a JEOL-JXA-8900R Electron Microprobe Analyser at the Institute of Mineralogy and Economic Geology, RWTH Aachen University. Operating conditions were 15 kV and 2 mA. Trace element analyses of Fe– Mg carbonate were performed by LA-ICP-MS with the ThermoFinnigan Element II sector field ICP-MS system coupled to a Merkantek/New Wave UP 213 nm

57

UV laser system at the Institute of Geosciences, Johann Wolfgang Goethe-University Frankfurt. Handpicked Fe –Mg carbonate was analysed for 18 O/16O and 13C/12C isotopic ratios. Mass spectrometric measurements were performed on a SIRA 9 triple-collector VG-Isogas instrument at the University of Bonn, Germany and on a Finnigan MAT Delta Plus XL mass spectrometre at the University of Lausanne, Switzerland. In the latter laboratory the samples were reacted at 90 8C for 1 h with 100% H3PO4 and analysed according to a method adapted from that of Spo¨tl & Vennemann (2003). The data are given in the usual d-notation v. VSMOW and PDB, respectively (see Table 5). The geochemical and isotopic data for the Fe–Mg carbonate was determined to better characterize its sedimentary or hydrothermal origin.

Lithology and lithogeochemistry of the Guelb Moghrein area Detailed open pit mapping revealed that the main lithologies at Guelb Moghrein are metasedimentary and meta-volcanic rocks of the Sainte Barbe volcanic unit and meta-volcanic rocks of the Akjoujt meta-basalts as well as the metacarbonate body that contains the ore bodies. All rocks are characterized by two almost orthogonal oriented foliations, which trend ESE–WNW (S2) and north–south (S3; Fig. 2a; Kolb et al. 2006).

Sainte Barbe volcanic unit The Sainte Barbe volcanic unit comprises quartz– sericite schists and biotite– garnet–quartz schists. These rocks are situated at about 150 m in the hanging wall of the orebody and are not affected by hydrothermal alteration. The quartz–sericite schist is a light grey to greenish, well-foliated, fine-grained rock, made up of sericite, biotite, quartz, plagioclase and chlorite. Rectangular to rounded quartz porphyroclasts enclose numerous, micrometre-scale vitreous inclusions, emphasizing the volcanic origin (Kolb et al. 2006). A representative suite of six samples from the quartz–sericite schists was chemically analysed. Relatively high SiO2 contents from 70 to 75 wt%, K2O concentrations between 2 and 4 wt% and minor Na2O (,0.7 wt%) underline the felsic nature of the deformed meta-volcanic rock (Table 1). Al2O3 and Fe2O3 vary between 11 and 13 wt% and 4 and 9 wt%, respectively, with TiO2 contents being relatively uniform around 0.8 wt%. The transition metals are all very low (,17 ppm); only V reaches values around 63 ppm. Large ion lithophile element (LILE) distributions are more inconsistent, reflecting variable modal composition

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J. KOLB ET AL.

Table 1. Major and selected trace element data for the rocks from the Sainte Barbe volcanic and the Akjoujt meta-basalt units Sample no.

(wt%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total (ppm) Rb Ba Zr Y

Quartz–sericite schist

Biotite –garnet – quartz schist

GM15

GM17

GM14

GM16

GM18

GM20

GM21

GM24

GM27

GM29

69.3 0.8 12.3 8.7 bdl 1.5 0.6 bdl 4.0 0.1 2.7 100.0

72.1 0.8 11.7 6.8 0.1 1.1 1.3 0.5 2.6 0.1 2.3 99.7

71.0 0.7 11.1 8.3 0.1 2.9 0.4 0.1 2.5 0.1 2.8 100.0

70.8 0.8 11.6 8.0 0.1 1.6 1.0 bdl 3.5 0.1 2.5 100.2

74.6 0.8 12.0 4.4 0.1 0.7 0.9 0.7 2.9 0.1 2.3 99.5

69.5 0.8 11.5 9.1 0.1 1.6 0.8 0.2 3.4 0.1 2.9 99.8

53.9 2.3 13.2 15.6 0.2 4.0 1.8 1.3 2.0 0.3 4.2 98.8

58.3 1.1 11.3 15.7 0.3 3.4 0.9 0.7 3.5 0.1 3.5 98.9

40.8 1.2 15.6 21.3 0.1 12.5 0.2 0.6 bdl 0.1 7.1 99.4

56.0 1.0 12.0 17.7 0.1 7.3 0.2 0.6 bdl 0.1 4.7 99.7

146 307 286 26.7

77 242 272 42.8

68 230 197 25

126 332 210 27

57 285 209 37

100 407 205 27

85 576 202 24

153 557 135 22

bdl 257 73 21

bdl 143 99 bdl

bdl, below detection limit; na, not analysed.

(e.g. biotite: Rb, Sr; see Tables 1 and 2). Silica is strongly negative correlated with Fe, whereas Mg shows negative correlations with Na and Ca. This is regarded a typical feature reflecting fractionation during magma crystallization. Based on the SiO2 v. Zr/TiO2 discrimination diagram, most of the quartz –sericite schists plot in the rhyodacite– dacite field, with one sample plotting in the rhyolite field (Fig. 3a). The primitive mantle normalized spider plot shows most of the trace elements to be enriched 10 –100 times (Table 2; Fig. 3b). The distribution pattern can be explained by magmatic differentiation typical for felsic rocks, with an enrichment of incompatible elements such as Rb, Th and U. The depleted nature of Sr, Sc, V, Zn, Cu, Ni and Cr is related to differentiation and crystallization of olivine, chromite and plagioclase (troctolite) in a magma chamber prior to melt extrusion or by partial melting of the mantle source, which is also suggested by the small negative Eu anomaly (Fig. 3b). The biotite –garnet –quartz schist is a yellowish brown, well-foliated, fine-grained rock comprising biotite, garnet, quartz and chlorite. Samples GM 27 and GM 29 are strongly retrogressed in the lower greenschist facies. Four samples were selected for chemical analysis. The rocks have relatively low SiO2 ,58 wt% and contain between 1 and 2.3 wt% TiO2 (Table 1). The Fe, Al, Mg, K and Mn-rich nature of the rocks owing to the dominant biotite-garnet assemblage is characteristic of

meta-pelitic rocks (Bauluz et al. 2000). Variations in the Rb, Sr and Ba content are due to variable plagioclase and biotite abundance. Compared with North Atlantic Shale Composite (NASC), the Y content of 2 ppm is fairly high, which appears to be a typical feature of Archaean meta-siliciclastic rocks (Bauluz et al. 2000). Relatively high V and low Cr concentrations are thought to reflect the characteristics of the source area. The metamorphic retrogression resulted in an enrichment of Fe and Mg and a depletion of Na, Ca, K, Ba and Si.

Akjoujt meta-basalt unit The amphibolite is a dark green massive to slightly foliated rock, comprising amphibole, plagioclase, ilmenite, chlorite, titanite and quartz. A suite of five representative samples was analysed. SiO2 contents vary between 54 and 59 wt% (Table 1), and Al2O3 and Fe2O3 concentrations range from 12 to 14 wt% and 12 to 19 wt%, respectively. TiO2 is relatively uniform with concentrations around 1.3 wt%. K2O contents are low (0.2–0.5 wt%), but owing to the pargasitic composition of amphibole and the albitic composition of plagioclase, Na2O values range from 1.5 to 5 wt%. This relationship is also shown by a strong correlation of Na with Al. The amount of CaO varies between 5 wt% in albite-rich and 7 wt% in hornblende-rich amphibolites. Among the transition metals, V shows relatively high values of up to 448 ppm. LILE and


Amphibolite

59

Biotite – actinolite schist

GM5

GM9

GM13

GM25

GM12

GM1

GM4

GM7

GM10

GM23

GM32

8801

55.9 1.3 13.0 15.4 0.2 3.3 4.9 3.9 0.2 0.2 0.3 98.9

59.0 1.2 12.0 11.8 0.1 5.2 4.8 5.0 0.2 0.1 0.7 100.1

54.0 1.2 12.1 18.8 0.2 2.9 7.1 1.5 0.3 0.1 0.6 99.2

56.3 1.2 13.3 15.1 0.1 3.0 5.2 4.0 0.3 0.1 0.7 99.4

55.1 1.2 13.4 15.2 0.1 3.5 5.6 3.5 0.5 0.1 0.5 98.9

63.5 1.4 15.3 6.9 0.1 2.1 1.7 7.3 0.2 0.2 0.2 98.9

56.2 1.2 13.7 16.2 0.1 3.0 1.2 5.1 1.1 0.1 1.8 99.8

65.5 0.7 12.2 9.5 0.1 2.4 0.5 3.6 1.7 0.2 2.4 98.8

45.7 1.1 16.1 17.0 0.1 7.2 3.6 3.4 2.2 0.1 2.9 99.3

35.4 1.1 18.7 20.5 0.2 9.1 5.6 2.9 0.2 0.1 5.7 99.6

53.8 1.9 15.3 14.5 0.1 4.2 3.0 5.4 1.3 0.2 0.3 100.1

57.8 1.3 12.6 16.9 0.1 3.1 5.6 2.5 1.0 0.1 0.8 101.8

4 28 162 25.9

5 20 142 29.6

4 44 145 31.8

3 41 150 28.3

89 297 193 34

65 152 150 32.6

na na 126 30

8 31 178 32.5

92 218 118 23

64 160 212 39

139 390 60 bdl

bdl 500 108 bdl

(Continued)

high field strength elements (HFSE) are characterized by moderate to low values. The chondrite (C1)-normalized REE plot displays a 10 –100 times enriched, almost flat or gently inclined pattern for the heavy REE (HREE) (GdN/ YbN ¼ 1.16– 1.56) and a steeper slope for the light REE (LREE) (LaN/SmN ¼ 2.18–3.17) (Fig. 4a). The LREE are enriched, with LaN/YbN ratios between 3.15 and 7.48. Most of the amphibolites display a moderate negative Eu anomaly (Eu/ Eu* ¼ 0.85–0.97). Major elements such as Al, Fe, Mg and Ca are negatively correlated with Si, which is a typical feature for magma fractionation. The biotite– actinolite schist is a fine-grained, green to dark brown rock comprising biotite, actinolite, albite, quartz, epidote, grunerite, chlorite, titanite and relic hornblende and ilmenite. It is characterized by a closely spaced S2 foliation and the replacement of hornblende by biotite, actinolite, epidote, chlorite and locally grunerite (sample 8802). The deformation and alteration is contemporaneous with the mineralization and occurs in the upper greenschist facies (Kolb et al. 2006). Geochemical analysis was performed on 10 representative samples collected in the pit and from drill core (8801, 8802, 8832). Geochemical changes compared with the amphibolites are minor and involve slight depletion in Ca and variable enrichments in Si, Na, Ba and K (Table 1). Sample 8802 contains traces of chalcopyrite, which explains Cu contents of 542 ppm. The mineralized sample 8802 is conspicuous by a pronounced Eu anomaly (Fig. 4a), most probably the result of hydrothermal alteration and replacement of plagioclase.

The chlorite schist has a closely spaced S3 foliation and represents the retrogressed equivalent of the amphibolite, comprising chlorite, quartz and plagioclase as well as minor sericite, ilmenite and magnetite. Retrogression of the rocks is contemporaneous with D3 thrusting to the east under lower greenschist-facies conditions (Kolb et al. 2006). Geochemical signatures were studied in 10 samples from the open pit. Most obvious is a strong depletion in Ca, Na, K and LREE, and enrichment in Mg and loss on ignition (LOI) compared with the amphibolites (Tables 1 and 2). Si is also slightly depleted in samples with high modal chlorite. The LREE are strongly depleted (LaN/ YbN ¼ 0.63 –1.69; GdN/YbN ¼ 0.49– 0.85) as a result of the complete replacement of feldspar and hornblende (Fig. 4a). Amphibolites, biotite–actinolite schists and chlorite schists plot in a close cluster in the Zr/ TiO2 v. Nb/Y discrimination diagram (Winchester & Floyd 1977), which shows that their precursor rock has an andesitic to basaltic nature (Fig. 4b) and alteration has not greatly affected these elements. In contrast, the pattern in Figure 4c is more scattered because of geochemical variations in the main elements, suggesting a tholeiitic composition for the volcanic precursor rocks (Irvine & Baragar 1971). Using the tectonomagmatic discrimination diagram with the relative immobile elements Zr, Nb and Y (Meschede 1986), the samples plot in a cluster in the volcanic arc basalt (VAB) field, which points to their origin in a volcanic setting at an active continental margin (Fig. 4d).

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J. KOLB ET AL.

Table 1. (Continued) Sample no.

Biotite– actinolite schist

Chlorite schist

8802 8832 DDGM1 GM2 GM38 GM3 GM6 GM28 GM33 GM34 GM36 GM37 GM39 (wt%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total (ppm) Rb Ba Zr Y

57.9 58.7 1.3 1.4 12.9 12.8 17.2 15.6 0.2 0.2 2.6 3.1 2.6 3.1 3.4 3.2 0.3 1.7 0.1 0.1 1.5 0.6 100.0 100.5 11 74 31 320 148 135 23.3 25

54.9 1.2 12.7 15.6 0.1 3.3 4.0 3.4 1.7 0.2 1.6 98.6 92 na 142 30

54.2 1.1 15.2 12.6 0.1 3.6 4.5 6.7 0.1 0.1 1.5 99.8

50.2 55.0 1.2 1.4 14.2 12.7 14.6 16.3 0.1 0.1 11.9 9.0 bdl 0.4 bdl 0.3 bdl bdl 0.1 0.1 5.9 5.0 98.1 100.3

4 2 bdl 20 81 109 144 142 131 31.1 16.2 bdl

Meta-carbonate The meta-carbonate is a massive, very coarsegrained, dark grey rock, which is mainly composed of euhedral to anhedral Fe –Mg carbonate grains up to 5 cm in diameter and accessory magnetite and graphite. The undeformed and unaltered Fe –Mg carbonate is not zoned and of pistomesite composition with an average XMg ¼ 36 and low Ca and Mn contents (Table 3). Post-Archaean Australian average shale (PAAS)-normalized REE þ Y patterns (McLennan et al. 1990) of four of these Fe–Mg carbonate grains show U-shapes for the LREE and successive enrichments in the HREE (NdN/YbN ¼ 0.03–0.05) with a positive Eu anomaly (Fig. 5a; Table 4). The oxygen isotope composition has d18O values of between 9 and 11‰. The carbon isotope composition is characterized by negative values with d13C ranging from 218 to 217‰ (Table 5). These very low values of carbon isotopes are characteristic of diagenetic siderite, which formed by bacterially mediated reactions in the marine environment and is difficult to be explained by a modification during hydrothermal overprint and/ or metamorphism (Moore et al. 1992; Mozley & Wersin 1992; Mozley & Burns 1993; Faure et al. 1995; Mortimer & Coleman 1997). The earlier researchers listed suggested that the negative carbon isotope values are generated by ferric iron and sulphate reduction accompanying oxidation of organic matter in suboxic conditions, which is restricted to marine settings. Furthermore, the Mg-rich composition of Fe–Mg carbonate is a typical feature of the marine environment (Mozley

49.3 1.4 13.5 20.9 0.1 6.6 0.2 bdl 0.2 bdl 5.8 98.0 na 16 148 na

54.8 55.1 55.7 55.6 54.5 1.4 1.3 1.4 1.2 1.5 11.9 11.8 12.4 12.4 12.9 16.1 17.5 12.2 10.2 10.6 0.1 0.1 bdl bdl bdl 9.2 9.1 12.2 14.4 13.9 0.3 0.2 0.2 0.2 0.2 bdl 0.1 0.1 bdl 0.2 0.1 bdl bdl bdl bdl 0.1 0.1 0.1 0.1 0.2 5.7 4.9 5.6 6.2 6.2 99.7 100.2 100.0 100.3 100.1 bdl 133 133 37

bdl 100 121 21

bdl 81 131 25

bdl 62 96 bdl

bdl 83 139 bdl

54.2 1.4 12.8 16.0 0.1 9.5 0.3 0.5 0.1 0.1 5.2 100.2 bdl 92 128 bdl

1989; Mozley & Carothers 1992). The PAASnormalized REE pattern, together with low Th, is consistent with a seawater signature and typical of Archaean carbonates and iron-formations (Fig. 5a; Table 4; Bau & Dulski 1996; van Kranendonk et al. 2003). The positive Eu anomaly is explained by a contribution of a hydrothermal fluid, which would be typical of an Archaean iron formation proximal to the hydrothermal vent site (Bau & Dulski 1996), and/or by later modification during metamorphism and alteration.

Fe – Mg clinoamphibole – chlorite schists Locally within the meta-carbonate body, Fe –Mg clinoamphibole–chlorite schists occur as narrow, up to 1 m wide layers. The rocks generally comprise up to 90 vol% chlorite together with minor Fe–Mg clinoamphibole (mainly grunerite), magnetite, ilmenite, apatite, calcite, quartz, monazite and allanite. In places, the Fe –Mg clinoamphibole may dominate the mineral assemblage. In this case, its chemical make-up is characterized by very low Al2O3 (,1 wt%) and TiO2 (,0.1 wt%) contents (Table 6). The composition of the schists in general is rather heterogeneous, with SiO2 between 21 and 55 wt% and Al2O3 between 0.7 and 21 wt%, reflecting the variable modal composition and the dominance of either chlorite or Fe – Mg clinoamphibole. This is also indicated by the systematic variation of Fe and Mg (Table 6). Na and K contents are generally low, whereas Ca concentrations vary according to elevated calcite and apatite contents in some samples. High Co, Cu, Bi, As and Au contents point to the fact that the

Table 2. Trace element data for the rocks from the Sainte Barbe volcanic and the Akjoujt meta-basalt units Sample no.

GM15

GM17

bdl 8 6 15.5 58 0.27 bdl 6 5 5.5 24.3 8 0.7 0.5

bdl 2 11 16.7 41 0.17 4 7 6 3.9 29.2 7 1.1 0.4

6 2.5 14 19 2.7 7.3 17.3 15.84 3.18 24.29 56.5 6.08 25.37 5.09 1.400 4.33

13 0.9 60 18 1.8 7.3 16.9 16.31 4.11 32.48 40.4 7.17 30.06 6.57 1.459 6.88

Amphibolite GM5 bdl 7 30 34.8 422 bdl bdl bdl 17 0.9 6.1 6 0.3 0.4 bdl 0.2 34 22 2.0 4.1 9.6 7.73 2.12 14.69 33.9 3.39 14.60 3.65 1.111 3.82

GM9

GM13

6 21 29 34.4 373 bdl 35 bdl 27 0.8 9.9 4 0.5 0.4

bdl 6 36 34.1 350 0.06 bdl bdl 18 1.0 12.5 5 0.6 bdl

5

4

0.4 39 15 1.6 3.7 8.3 7.25 1.54 13.34 30.8 3.63 15.79 3.85 1.234 4.35

0.2 16 21 2.3 3.9 8.5 7.97 1.83 37.98 106.5 9.47 38.26 7.54 2.003 5.26

Biotite – actinolite schist GM25

GM1

8801

8802

bdl 6 10 30.4 363 0.07 99 bdl 8 28.7 6.7 11 0.3 0.4

bdl 6 44 39.2 494 0.31 157 5 19 0.8 3.8 7 0.2 0.4

13 7 25 36.0 476 0.09 542 bdl 18 2.8 3.9 7 0.6 0.5

bdl

bdl

bdl

0.2 24 18 1.8 4.1 8.9 8.39 1.66 27.79 51.8 6.06 26.06 5.59 1.537 5.41

0.5 31 15 1.5 4.5 9.9 8.64 2.24 20.98 41.3 4.48 19.35 4.32 1.251 4.52

3.7 31 22 2.1 3.8 9.1 7.00 2.74 28.23 52.0 5.85 25.34 5.61 1.909 5.80

11 9 40 29.7 348 bdl bdl bdl 16 bdl 24.6 5 0.4 0.4

Chlorite schist GM2 bdl 7 29 27.9 253 bdl 1 bdl 16 2.5 11.1 4 0.3 0.4

21

16

0.9 51 21 2.7 3.8 9.3 7.09 2.42 16.29 32.4 3.91 17.84 4.88 1.118 5.37

0.3 21 17 1.6 3.7 8.2 5.81 1.63 6.79 8.9 1.25 5.58 1.53 0.566 2.88

GM38 11 44 34 35.5 382 bdl 42 4 15 2.4 31.9 4 0.8 0.6 bdl 0.2 bdl 21 1.9 3.6 8.2 6.82 2.03 2.00 2.5 0.43 2.03 0.67 0.309 1.31

61

(Continued)


(ppm) Cr Ni Co Sc V Tl Cu Pb Zn W As Sn Sb Ag (ppb) Au (ppm) Cs Sr Ga Ge Hf Nb Th U La Ce Pr Nd Sm Eu Gd

Quartz –sericite schist

62

Table 2. Continued. Sample no.

GM15

GM17

0.76 4.71 1.02 3.37 0.558 3.71 0.596 1.53

1.18 7.40 1.56 4.74 0.712 4.32 0.630 1.52

bdl, below detection limit.

GM5 0.71 4.23 0.91 2.81 0.414 2.67 0.421 0.76 1.2 2.6 3.9 0.9

GM9 0.79 5.04 1.09 3.21 0.459 2.86 0.415 0.70 1.3 2.2 3.3 0.9

GM13 0.88 5.18 1.13 3.60 0.543 3.43 0.519 0.72 1.3 3.3 7.9 1.0

Biotite – actinolite schist GM25 0.89 5.18 1.06 3.21 0.461 2.81 0.423 0.75 1.6 3.2 7.1 0.9

GM1 0.81 5.25 1.15 3.54 0.531 3.29 0.486 0.88 1.1 3.1 4.6 0.9

8801 1.00 5.76 1.17 3.43 0.471 2.88 0.428 0.73 1.7 3.2 7.0 1.0

Chlorite schist

8802 0.88 4.57 0.86 2.50 0.356 2.09 0.330 0.73 2.1 2.2 5.6 0.7

GM2 0.66 4.76 1.07 3.20 0.463 2.72 0.410 0.64 0.9 2.9 1.8 0.8

GM38 0.30 2.22 0.54 1.84 0.308 2.15 0.352 0.65 0.5 1.9 0.7 1.0

J. KOLB ET AL.

Tb Dy Ho Er Tm Yb Lu Ta GdN/YbN LaN/SmN LaN/YbN Eu/Eu*

Amphibolite

Quartz –sericite schist


(a)

(b) 80

1000 b

sample/primitive mantle

Rhyolite

SiO2 [wt.%]

70 Rhyodacite–Dacite Trachyte

60

Andesite Phonolite

50

40 .001

63

100

10

1

GM15 GM17

0.1 0.01

0.1 Zr/TiO2* 0.0001

1

10

K Rb C Tl Pb Ba Th U Nb La Ce Sr Pr Nd Zr Sm Eu Gd Tb Ti Dy Y Ho Er Tm Yb Lu Sc

Fig. 3. Sainte Barbe volcanic unit. (a) Zr/TiO2 v. SiO2 discrimination diagram showing that the quartz–sericite schists have a rhyodacitic –dacitic composition, with one sample plotting in the rhyolite field. (b) Primitive mantle normalized spider diagram. (Note the strong negative Sr anomaly, suggesting partial melting or fractionated crystallization in a magma chamber.)

sample/C1

Amphibolite

Biotite–actinolite schist

100

Chlorite schist

GM5 GM9 GM13 GM25 GM1 8801 8802 GM2 GM38

(b)

5 Amphibolite

Zr/ TiO2* 0.0001

(a) 1000

1

Com/Pant Phonolite

Biotite–actinolite schist Chlorite schist

Rhyolite Trachyte

Rhyodacite/Dacite

0.1

TrachyAnd Andesite

10 0.01

Andesite/Basalt Alk–Bas SubAlkaline Basalt

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

0.1

1

10

Nb/Y

(c)

(d)

FeOt

Nb*2 Amphibolite Biotite–actinolite schist Chlorite schist

TH

PB

W

M

Amphibolite

O RB

CA

Biotite–actinolite schist

VA B

Chlorite schist

Na2O + K2O

MgO

Zr/4

Y

Fig. 4. Akjoujt meta-basalt unit. (a) Chondrite-normalized REE plot showing the typical LREE-enriched pattern of mafic volcanic rocks. It should be noted that the pattern of amphibolites and biotite actinolite schist are similar and that chlorite schists are strongly depleted in LREE. (b) In the Nb/Y v. Zr/TiO2 0.0001 discrimination diagram, all samples form a cluster in the andesite–basalt field. (c) The samples scatter in the FeOt (total Fe)–Na2O þ (K2O)–MgO plot in the tholeiitic composition field, as a result of geochemical variations from later tectonometamorphic overprints (TH, tholeiitic; CA, calc-alkaline). (d) In the tectonic discrimination diagram (Zr/4 –Y –Nb 2) all samples plot in a trend characteristic of volcanic arc basalts.

64

J. KOLB ET AL.

Table 3. Electron microprobe data for Fe–Mg carbonate from the meta-carbonate Sample no.

27140-6

27123-6

(wt%) FeO 36.3 35.8 MnO 1.9 2.2 MgO 17.9 18.3 CaO 0.5 0.5 Recalculated based on cation composition 43.4 43.6 CO2 Total 100.0 100.4 33.0 33.8 XMg

27123-7

27123-8

27134

27134

27134-9

34.1 2.3 18.9 0.4

34.0 2.1 19.3 0.3

32.5 2.1 20.7 0.3

32.6 2.1 20.8 0.3

34.7 2.1 18.8 0.7

43.3 99.1 35.6

43.5 99.2 36.2

44.1 99.7 39.0

44.3 100.3 39.0

43.7 100.2 35.1

Fe–Mg clinoamphibole –chlorite schists are, locally, strongly mineralized and hydrothermally altered. It is not possible to distinguish altered and least altered samples by their silicate mineralogy however, Fe–Mg clinoamphibole-rich samples are more often mineralized (Kolb et al. 2006). The PAAS-normalized REE þ Y pattern (McLennan et al. 1990) of the Fe–Mg clinoamphibole–chlorite schists shows for seven out of the 10 samples a U-shape for the LREE with a distinct positive Eu anomaly (Eu/Eu* ¼ 1.91 –1.24) whereas the HREE follow a flat trend pattern closely resembling the PAAS composition. Sample 8834 is significantly enriched in LREE as a result of an unusually high monazite content (Table 6; Fig. 5b). Samples 8833 and 8809 belong to the Fe –Mg clinoamphibole-rich variety and are conspicuous by strongly depleted

LREE. The U-shape pattern of the LREE together with the flat distribution of the HREE observed in most samples closely resembles Archaean iron formation patterns with a seawater signature (Bau & Dulski 1996; Khan et al. 1996; van Kranendonk et al. 2003). The low Y/Ho ratios and the high Th and Sc contents point to a significant contribution of detrital material (Bau & Dulski 1996; van Kranendonk et al. 2003). The chemical signature as well as the occurrence of the rock as small slices within the meta-carbonate indicates that the Fe –Mg clinoamphibole–chlorite schists represent ironrich marine sediments. The positive Eu anomaly may be due to a hydrothermal fluid overprint. It is, however, not clear if this signature is synsedimentary or related to the later hydrothermal fluid infiltration during mineralization.

Fig. 5. (a) PAAS-normalized REE þ Y plot for Fe–Mg carbonate of the meta-carbonate, with the typical U-shape of the LREE suggesting deposition in a marine environment typical of Archaean iron formations. (b) PAAS-normalized REE þ Y plot for the Fe–Mg clinoamphibole–chlorite schists, with a weak U-shape of the LREE and PAAS composition of the HREE suggesting deposition in a marine environment and a significant contribution by continental detritus.


Table 4. LA-ICP-MS trace element data for Fe–Mg carbonate from the meta-carbonate Sample no. 26286-1 26286-2 (ppb) Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Pb Th U Ce/Ce* Pr/Pr* Nd/Yb Eu/Eu* Gd/Gd* Er/Er*

871 1723 1105 93 7 7170 179 266 27 139 84 72 223 22 241 60 211 44 421 121 1649 3 5 0.87 0.82 0.03 2.11 1.01 0.87

855 1516 1109 69 6 5950 201 330 37 217 138 55 207 25 164 52 197 36 340 80 1474 33 6 0.88 0.80 0.05 1.48 1.06 0.98

26286-3 26286-4 954 2028 840 73 7 8230 156 203 28 125 73 33 153 18 190 36 161 42 424 116 1227 1 3 0.70 1.01 0.02 1.34 1.22 0.81

911 1471 732 67 7 5790 127 181 24 150 78 44 180 20 142 33 147 32 356 98 1143 6 5 0.75 0.82 0.04 1.54 1.17 0.93

Lithology and lithogeochemistry of the hydrothermal IOCG mineralization Copper-, Co- and Au-rich zones occur in breccia zones developed in the meta-carbonate and form multiple, up to 30 m wide, coalescing lenses that dip moderately SW. They are exclusively related to discrete D2 reverse shear zones located in the Fe– Mg clinoamphibole–chlorite schists enveloped by the meta-carbonate host. Single shear zones are between 5 cm and 1 m wide and form an undulating,

Table 5. Stable isotope composition of Fe–Mg carbonate

97.03 97.08 97.12 97.13 97.16

d18O (SMOW)

d13C (PDB)

12.70 10.94 8.94 12.69 10.42

218.56 216.59 217.05 218.51 217.92

65

broadly tabular network subparallel to the S2 foliation (Kolb et al. 2006). Two types of monomict breccias are distinguished (Kolb et al. 2006). (1) Puzzle-like breccias, locally, form lenses up to 5 cm wide and up to 20 cm long in the Fe– Mg clinoamphibole –chlorite schists. Angular fragments of the schists are up to 2 cm in diametre within a massive sulphide matrix. The fragments are not rotated, resulting in the puzzle-like appearance of the breccia. (2) A pebble-like breccia, consisting of rounded, pebblelike Fe– Mg carbonate clasts within a matrix of a complex arsenide–sulphide –gold, magnetite and Fe–Mg clinoamphibole assemblage, forms tabular bodies up 30 m wide. These bodies are mineralogically zoned and situated about 20 cm –1 m from the lithological contact with the biotite– actinolite and Fe –Mg clinoamphibole– chlorite schists. Idiomorphic magnetite and Fe – Mg clinoamphibole predominate in the matrix (Fig. 6a). The arsenide– sulphide–gold paragenesis comprises pyrrhotite, chalcopyrite, Fe–Co– Ni arsenides, arsenopyrite, cobaltite, uraninite and Bi –Au–Ag– Te minerals. The distal ore breccia, dominated by magnetite and Fe–Mg clinoamphibole, has a high Fe2O3 content between 35 and 60 wt% and MgO between 13 and 18 wt% with generally low SiO2 ,4 wt% (Table 7). Locally elevated Ca is explained by the occurrence of calcite in the alteration assemblage. Trace elements such as Ni, Co, Cu, Mo and As are incorporated into the arsenide– sulphide– gold minerals and are variable, reflecting the variable clast– matrix ratio of the analysed bulk-rock samples. Elevated Cr and V correlate with high magnetite contents, whereas Rb, Ba, Sr and Y correlate with the modal occurrence of Fe –Mg carbonate clasts. Samples 8820 and 8822 represent the central sulphide-dominated ore-breccia and have, therefore, ore-grade Ni, Co, Cu and As in the range of 2 –3 wt% (Table 7). Massive chalcopyrite breccias are locally developed and record up to 20 wt% Cu. The deformed and brecciated Fe–Mg carbonate grains are zoned and become more Fe-rich towards the rims (Table 8; Fig. 6b). The core of this Fe– Mg carbonate variety has a higher XMg than the original undeformed Fe –Mg carbonate of the meta-carbonate body. The brecciated Fe –Mg carbonate differs also significantly from the undeformed variety in the trace element composition (Table 9; Fig. 6c), particularly in the significant depletion of Rb, Sr and Ba (Table 9). Although the PAAS-normalized REE þ Y patterns are similar, the brecciated Fe– Mg carbonate is enriched in the HREE and shows a wider scatter in the LREE. The positive Eu anomaly typical for the undeformed Fe –Mg carbonate is locally absent. Some of the brecciated

66

Table 6. Major and trace element data for the Fe–Mg clinoamphibole – chlorite schists Sample no. 8016

8803

8804

24.8 1.3 20.4 34.0 0.2 10.6 bdl bdl 0.1 0.0 7.6 99.0

48.2 1.3 14.5 16.3 0.1 14.1 0.5 0.1 0.0 0.1 6.1 101.5

48.0 0.9 11.5 16.5 0.2 5.5 6.8 2.2 0.7 0.1 8.1 100.5

bdl bdl 34 na 444 na bdl na na na na na na na na

9 10 31 33.6 315 0.23 6 bdl 15 bdl 3.9 4.2 4 0.5 0.4

8806 54.5 0.6 6.9 28.4 0.6 6.0 0.5 3.1 0.9 0.1 20.4 101.2 13 14 30 42.1 84 0.16 bdl bdl 12 bdl 41.9 26.3 4 0.4 0.4

8808 43.9 0.9 11.4 31.6 0.3 8.2 0.2 0.2 0.7 0.1 4.2 101.7 11 22 47 24.0 227 0.31 bdl 4 10 bdl 29.2 29.2 4 0.6 0.4

8816 20.7 1.3 17.2 25.2 0.1 14.7 0.0 bdl 0.0 0.1 18.3 97.5

8827

8828

Chlorite– Fe – Mg 8830

8831

8834

54.2 1.3 12.5 15.8 0.1 10.2 0.4 0.2 0.3 0.1 5.1 100.2

38.9 1.4 13.2 24.3 0.2 12.6 4.9 0.5 0.2 0.1 4.4 100.8

43.6 0.9 9.5 20.8 0.2 13.2 8.3 0.6 0.2 0.1 3.5 100.8

45.1 0.5 5.6 29.4 0.5 15.4 1.4 0.3 0.0 0.3 2.5 101.0

36.3 1.9 14.9 20.2 0.2 16.2 5.3 0.2 0.1 0.1 5.7 101.2

7 10 29 34.0 453 0.06 14 bdl 10 bdl 65.4 2.8 3 bdl 0.3

10 42 82 25.0 298 0.06 bdl bdl 13 5.64 1.5 61.2 8 bdl 0.5

14 62 172 32.0 191 0.14 2 bdl 12 9.26 6.4 239.0 8 0.2 0.4

14 82 76 35.0 151 bdl 1692 bdl 12 4.60 38.0 111.9 5 0.4 0.4

bdl 59 62 43.1 531 bdl bdl 5 14 13.41 2.1 5.4 5 0.3 0.3

9728

8809

8813

8817

8833

30.9 1.4 15.1 32.0 0.2 12.8 0.0 bdl 0.1 0.1 6.0 98.7

48.6 0.1 0.7 38.0 0.7 12.7 0.3 0.4 0.0 0.1 20.3 101.2

46.9 0.0 0.1 37.9 0.5 14.1 0.3 0.3 0.1 0.1 1.1 101.3

48.6 0.0 0.1 34.5 0.5 15.0 0.1 0.3 0.0 0.1 1.4 100.6

54.8 0.0 0.5 22.7 0.5 17.9 3.6 0.1 0.0 0.1 0.4 100.8

22 57 74 41.6 109 bdl 1252 bdl 11 2.86 4.3 374.0 3 0.4 0.4

na na 1039 na 96 na bdl na na na na adl na na na

na na 115 na 26 na adl na na na na bdl na na na

6 35 43 20.3 28 bdl bdl bdl 13 8.31 bdl 4.4 4 0.3 0.4

(Continued)

J. KOLB ET AL.

(wt%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total (ppm) Cr Ni Co Sc V Tl Cu Pb Zn Bi W As Sn Sb Ag

Fe – Mg clinoamphibole– chlorite schist

Table 6. Continued. Sample no.

Fe – Mg clinoamphibole– chlorite schist 8016

8804

na

bdl

bdl na 54 bdl na na na 106 na 26 na na 54 na na na na na na na na na na na na na na

38 4.4 56 31 20 1.8 3.0 118 7.1 20.6 5.16 1.63 19.19 35.2 4.00 16.41 3.39 1.034 3.48 0.59 3.51 0.73 2.26 0.336 2.11 0.328 0.54

8806 4 36 2.7 64 8 10 2.4 3.9 153 12.8 36.6 9.06 2.96 27.05 50.8 5.71 23.81 4.94 1.517 4.97 0.92 5.85 1.27 4.10 0.611 3.88 0.588 1.04

8808 6 42 3.9 52 bdl 17 2.2 2.9 114 6.7 31.7 4.97 1.79 31.97 56.1 6.26 25.60 5.41 2.172 5.29 0.87 5.27 1.13 3.47 0.516 3.21 0.477 0.50

8816

8827

8828

8830

bdl

bdl

bdl

12 1.1 12 2 23 2.2 3.8 146 9.1 20.1 7.73 2.41 25.98 50.8 5.87 25.69 5.38 1.334 4.75 0.73 3.44 0.66 2.12 0.327 2.19 0.341 0.75

9 0.8 12 bdl 22 1.9 4.4 167 9.7 48.9 7.77 2.77 31.32 57.5 6.51 27.97 6.32 2.014 7.72 1.46 9.02 1.78 5.10 0.718 4.16 0.611 0.75

8 0.8 8 3 16 2.1 3.0 116 6.9 46.5 3.78 1.78 32.66 56.6 6.30 26.49 5.53 1.873 6.77 1.36 8.38 1.75 5.11 0.722 4.14 0.592 0.48

8831 56 2 0.3 bdl bdl 9 1.9 1.6 60 3.9 37.5 3.29 1.44 20.15 36.3 4.14 18.21 4.55 1.059 6.01 1.18 6.90 1.42 3.98 0.558 3.05 0.407 0.25

8834 7 2 0.2 5 3 24 2.2 5.7 210 12.6 31.6 11.97 5.02 121.72 223.1 25.88 112.23 21.96 4.836 15.21 1.72 7.17 1.20 3.51 0.463 3.16 0.478 1.08

9728

8809

8813

8817

8833

247

na

na

bdl

4 0.2 6 bdl 3 2.4 0.3 16 1.0 5.3 0.49 0.39 2.14 3.8 0.45 2.02 0.55 0.183 0.78 0.15 1.00 0.21 0.68 0.120 0.82 0.129 0.06

bdl na na 35 58 na na na na na na na na na na na na na na na na na na na na na na

bdl na 63 33 na na na 25 na na na na na na na na na na na na na na na na na na na

bdl bdl bdl bdl 1 2.3 0.2 16 0.9 13.3 0.11 0.08 0.21 0.5 0.09 0.65 0.46 0.188 1.26 0.34 2.29 0.49 1.43 0.208 1.15 0.157 0.02


(ppb) Au (ppm) Rb Cs Ba Sr Ga Ge Hf Zr Nb Y Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Ta

8803

Chlorite– Fe – Mg

bdl, below detection limit; adl, above detection limit; na, not analysed.

67

68

J. KOLB ET AL.

Fig. 6. (a) Schematic drill core sketch with the strongly S2 foliated Fe– Mg clinoamphibole– chlorite schist at the base, followed by a massive sulphide ore breccia and a more Fe– Mg clinoamphibole– magnetite-dominated breccia and the undeformed Fe– Mg carbonate at the top. Ductile D2 shearing occurred in the schists, whereas Fe– Mg carbonate was brecciated. (b) Electron microprobe profile from centre to rim of a brecciated Fe–Mg carbonate, showing the characteristic zoning from an Mg-rich core to an Fe-rich rim. CO2, MnO and CaO form relatively flat curves. The shaded grey bars represent the FeO and MgO composition of the undeformed Fe–Mg carbonate, respectively. It should be noted that XMg is lower than in the brecciated Fe– Mg carbonate. (c) The PAAS-normalized REE þ Y pattern of the brecciated Fe–Mg carbonate is similar to the original pattern shown in grey. The LREE are, however, more scattered and the HREE are slightly enriched in most samples.

Fe–Mg carbonate developed a negative Ce anomaly (defined by Ce/Ce* , 1 and Pr/ Pr* . 1; see Bau & Dulski 1996).

Discussion Two major tectonometamorphic events overprinted the rocks at Guelb Moghrein (Kolb et al. 2006; Meyer et al. 2006). (1) During D2/M2 at c. 2492 Ma, the Sainte Barbe volcanic unit was thrust to the NNW onto the Akjoujt meta-basalt unit, under upper greenschist-facies conditions. Contemporaneously, shear zones developed within and at the contact of the Guelb Moghrein meta-carbonate bodies that acted as conduits for the hydrothermal fluids responsible for the IOCG mineralization. (2) Eastward thrusting, under retrograde lower greenschist-facies conditions, at about 1742 Ma resulted in transposition of D2 fabrics and the displacement of the ore body during D3/M3. Some geochemical compositions of the rocks from Guelb Moghrein appear to be unmodified during deformation, metamorphism and hydrothermal alteration and, thus, allow for a discussion of the environment during the formation of the protoliths of the IOCG mineralization. Furthermore, the geochemical changes imposed by the hydrothermal alteration can be related to the composition of the protoliths and be discussed in terms of fluid composition.

Geochemical composition of the protoliths The stratigraphic base (structural high) is formed by the rocks of the Sainte Barbe volcanic unit comprising meta-volcanoclastic rocks (quartz–sericite schist) of rhyodacite–dacite composition and metapelites (garnet–biotite–quartz schist). The composition of the felsic meta-volcanoclastic rocks is typical for a continental arc or island arc setting. Similarly, the various rocks of the Akjoujt metabasalt unit also have a composition typical for volcanic arc rocks, but with a more mafic tholeiitic basalt –andesitic composition (Fig. 4). Their chemical composition, as demonstrated by the primitive mantle-normalized spider diagrams, is rather similar, with enrichments of incompatible elements such as Th, U and the depletion of Ni and Cr (Figs 3 and 7). This pattern is explained by fractional crystallization of olivine, chromite and plagioclase and, because of similarities, the compositional trends may imply a similar source for the volcanic rocks of the Sainte Barbe volcanic and the Akjoujt metabasalt units. An alteration effect by hydrothermal fluids can be ruled out because fresh unaltered rocks such as the amphibolite and quartz–sericite schist have the same composition as the altered and retrogressed lithologies. The meta-carbonate unit, including the massive Fe –Mg carbonate and the Fe –Mg

Table 7. Major and trace element composition of the ore breccia in the meta-carbonate Sample no. 8810

8811

3.1 bdl 0.7 46.2 1.2 15.6 2.4 0.8 bdl bdl 30.4 100.5

2.1 bdl 0.8 49.6 1.2 15.3 0.4 0.9 bdl bdl 29.0 99.3

77 71 104 175 adl bdl 41 42 bdl 64 44 32 bdl 194

66 21 92 96 265 bdl 43 135 bdl 129 40 35 bdl 195

8818 2.6 bdl 0.6 45.0 1.2 16.4 0.3 1.0 bdl 0.1 30.6 97.8 61 197 538 60 adl bdl 41 1134 bdl 147 40 39 bdl 251

8823 3.7 bdl 0.3 35.4 1.1 16.7 9.7 0.5 0.1 bdl 33.6 101.0 35 177 150 36 adl bdl 27 337 bdl 57 48 20 73.0 149

DDGM3 2.3 bdl 0.7 47.7 1.1 16.7 0.6 0.4 0.1 bdl 30.0 99.7 59 25 64 64 1447 bdl 30 bdl 35 na 32 38 bdl na

Massive sulphide ore breccia

DDGM5

DDGM7

DDGM8

0.8 bdl 0.7 45.4 1.1 17.3 0.3 0.4 0.1 bdl 33.1 99.1

0.5 bdl 0.5 47.6 1.5 18.0 0.7 0.4 0.1 bdl 29.9 99.1

1.3 bdl 0.6 59.9 1.3 12.8 0.5 0.4 0.1 bdl 22.0 98.8

70 75 135 149 1537 bdl 28 40 bdl na 29 106 26.0 na

70 30 113 169 847 bdl 34 191 bdl na 37 31 23.0 na

78 35 149 335 204 bdl 52 137 21 na 47 61 bdl na

8820 6.0 bdl 0.8 64.7 0.7 8.0 0.2 1.3 bdl 0.1 15.4 97.3 63 2429 819 224 adl 22 102 1055 24 139 89 73 bdl 382

8822 1.3 bdl 0.2 54.8 1.0 12.2 0.6 0.9 bdl bdl 26.1 97.2 53 1351 553 95 adl bdl 76 610 27 189 66 49 bdl 293


(wt%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total (ppm) Cr Ni Co V Cu Zn Mo As Rb Ba Sr Zr Y La

Magnetite-dominated ore breccia

bdl, below detection limit; adl, above detection limit; na, not analysed.

69

70

J. KOLB ET AL.

Table 8. Representative electron microprobe data for Fe–Mg carbonate from the ore breccia (sample 27138-10) Analysis no.

102 core

103

104

FeO 28.2 28.9 30.0 MnO 1.7 1.7 1.6 MgO 24.0 23.1 23.2 CaO 0.4 0.4 0.4 SrO bdl bdl bdl Recalculated based on cation composition CO2 44.9 44.3 45.0 Total 99.2 98.4 100.2 XMg 46.0 44.5 43.6

105

106

107

108

109 rim

30.1 1.6 23.2 0.3 bdl

33.4 1.8 19.8 0.4 bdl

34.3 1.8 19.9 0.4 bdl

36.5 1.6 18.6 0.3 bdl

41.0 2.4 14.4 0.4 bdl

45.0 100.3 43.5

43.6 99.1 37.3

44.2 100.6 36.7

43.8 100.8 33.7

42.6 100.8 26.1

bdl, below detection limit.

clinoamphibole–chlorite schists, has geochemical compositions typical for marine sediments (Fig. 5), where the Fe –Mg carbonates may have been deposited on a continental shelf area. The fact that the positive Eu anomaly is erased in some altered Fe–Mg carbonate samples (Fig. 6) may point to a contribution by high-temperature

hydrothermal fluids during the precipitation process (Bau & Mo¨ller, 1992; Khan et al. 1996; Hecht et al. 1999). The Fe –Mg clinoamphibole– chlorite schists similarly have a distinct positive Eu anomaly but resemble typical shales, especially in their HREE distribution (Figs 6 and 7). These rocks are interpreted to represent Fe-rich marine

Table 9. LA-ICP-MS trace element data for Fe–Mg carbonate from the ore breccia Sample no. 26284-1 26284-2 26284-3 26284-4 26284-5 26285-1 26285-2 26285-3 26285-4 26285-5 (ppb) Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Pb Th U Ce/Ce* Pr/Pr* Nd/Yb Eu/Eu* Gd/Gd* Er/Er*

126 549 2156 52 10 1901 304 352 66 173 147 57 207 32 335 99 392 115 1308 357 1110 2 10 0.57 1.57 0.01 1.49 0.95 0.71

108 431 2414 129 14 1411 189 198 32 120 62 32 103 28 407 112 481 152 1678 448 943 2 7 0.58 1.21 0.01 1.76 0.68 0.69

386 624 1505 152 25 2409 228 212 39 210 98 35 128 30 250 68 259 66 664 172 1483 7 9 0.52 0.98 0.03 1.42 0.77 0.79

229 493 714 101 10 2214 119 176 16 99 67 27 94 12 159 37 121 36 448 120 1109 1 5 0.90 0.71 0.02 1.57 1.00 0.68

247 627 2412 47 29 1920 384 615 71 396 197 84 299 52 399 112 408 101 1048 297 850 2 8 0.85 0.83 0.03 1.56 0.89 0.79

122 1248 1321 250 10 2206 131 162 30 106 83 34 76 20 149 58 305 95 1227 313 1067 4 12 0.59 1.33 0.01 2.03 0.56 0.73

60 383 1547 142 11 2070 70 92 16 44 69 16 70 18 229 65 344 102 1316 342 720 6 12 0.64 1.43 0.00 1.09 0.80 0.76

81 965 1766 39 7 923 48 94 15 95 65 20 133 20 168 85 435 123 1353 345 327 2 5 0.80 0.85 0.01 0.94 1.26 0.78

266 807 3980 54 4 2435 98 169 26 157 78 35 257 40 410 167 961 263 2828 707 744 4 7 0.76 0.88 0.00 0.93 1.32 0.83

61 1596 1356 142 8 1747 115 181 27 185 102 40 115 13 168 66 304 96 1263 349 692 2 6 0.75 0.80 0.01 1.71 0.93 0.70


(b)

(a)

71

100

1000

10 1 Amphibolite

0.1 0.01

Biotite-actinolite schist Chlorite schist

0.001

GM5 GM9 GM13 GM25 GM1 8801 8802 GM2 GM38

K Rb Cs Tl Pb Ba Th U Nb Sr Zr Ti

1

Sample/NASC

Sample/primitive mantle

10 100

0.1

8804 8806

0.01

8808 8827

0.001

8828 8830 8831

Y Sc V Zn Cu Ni Cr

0.0001 Na Al K Ca Sc Ti Cr Mn Fe Co Ni As Rb Sr Zr Sb Cs Ba La Ce Nd Sm Eu Tb Yb Lu Hf Ta Y Th U

Fig. 7. (a) Primitive mantle normalized spider diagram for the rocks of the Akjoujt meta-basalt unit. (Note the strong negative Sr, Ni and Cr anomalies.) Cu is enriched in the mineralized samples. Nb, Zr, Ti and Sc display only a small scatter irrespective of the altered nature of the rocks, which suggests that these elements were largely immobile during the hydrothermal overprints of the biotite–actinolite schists and the chlorite schists. (b) NASC-normalized spider diagram for the Fe– Mg clinoamphibole– chlorite schists with the NASC signature of the REE and the positive Y anomaly, which is typical of Archaean meta-pelites. Negative anomalies of K, Cr, Rb, Sr and Ba are explained either by characteristics of the source region of the sediments or by mobilization during the hydrothermal overprint.

shales with a terrigeneous detritus component and a hydrothermal component, which were deposited as interlayer sediments in the Fe– Mg carbonate body. The NASC-normalized trace element pattern (McLennan & Taylor 1991) has significant negative anomalies in K, Cr, Rb, Sr, Sb and Ba. In particular, K, Cr, Sr and Ba are similar in relative abundance to those of the meta-volcanic rocks from the Sainte Barbe volcanic and the Akjoujt meta-basalt units. This feature could represent the effect of hydrothermal modifications (Figs 3 and 7). However, the volcanic rocks and meta-basalts may also represent the source of the terrigeneous input for the Fe– Mg clinoamphibole –chlorite schists. Our interpretation of the nature of the carbonate body at Guelb Moghrein is in contrast to a previous proposal by other workers, who suggested an origin by hydrothermal alteration and by replacement of the amphibolites (Strickland & Martyn 2002; Martyn & Strickland 2004). Our explanation is based on the following observations: (1) Fe –Mg carbonate is replaced during the hydrothermal mineralization and alteration; (2) the Fe–Mg carbonate grains are deformed and the hydrothermal mineral paragenesis forms the matrix of the breccia; (3) no relics of unaltered meta-volcanic rock are found within the meta-carbonate body; (4) marine sediments (i.e. the Fe–Mg clinoamphibole–chlorite schist) form layers in the meta-carbonate body (this study; Kolb et al. 2006). Except for these layers, sedimentary textures are absent, as a result of the metamorphic recrystallization of Fe–Mg carbonate forming grains up to 5 cm in diametre. The geochemical composition of the Fe –Mg carbonate grains remained relatively unaffected and still preserves a marine-diagenetic composition with a seawater-like REE pattern and carbon isotopic

compositions typical of such settings. Typical hydrothermal Fe– Mg carbonate has a markedly different isotopic composition with d18O values of about 18‰ (compared with 9–11‰ at Guelb Moghrein) and d13C values of about 25‰ (compared with 217 to 218‰ at Guelb Moghrein) (see Radvanec et al. 2004). The positive Eu anomaly observed in the REE pattern of the Fe –Mg carbonate and the Fe–Mg clinoamphibole –chlorite schists is typical for hydrothermal Archaean carbonate- and silicatefacies banded iron formations (Gnaneshwar Rao & Naqvi 1995; Bau & Dulski 1996; Khan et al. 1996). As has been shown before, the Akjoujt metabasalt unit contains only minor thin banded iron formations (Strickland & Martyn 2002; Martyn & Strickland 2004) so that the meta-carbonate cannot be grouped stratigraphically into this unit. We, therefore, propose to group the mineralized meta-carbonate of Guelb Moghrein into the Lembeitih formation, which represents a regional iron formation marker and is stratigraphically positioned between the Sainte Barbe volcanic and the Akjoujt meta-basalt units (Strickland & Martyn 2002; Martyn & Strickland 2004). The metacarbonate now is an imbricated slice within the Akjoujt meta-basalt unit and was tectonically emplaced into the current position during regional D2 thrusting (Kolb et al. 2006). The rocks of the Akjoujt area, consequently, represent a typical metamorphosed and deformed Archaean greenstone assemblage of mafic to felsic meta-volcanic and meta-volcanoclastic rocks with minor pelitic sediments and a thick iron formation (see de Wit 1998). The close association of BIF with mafic or intermediate volcanic rocks in this previously believed Proterozoic succession was

72

J. KOLB ET AL.

also noted by Strickland & Martyn (2002) and Martyn & Strickland (2004). The radiometric ages and the geochemical interpretations of the rocks of the Akjoujt area suggest that this terrane underwent a different tectonometamorphic history from the southern Mauritanides and from that previously believed (this study; Kolb et al. 2006; Meyer et al. 2006). Significant tectonometamorphic events occurred at 2492 Ma and 1742 Ma (Fig. 2), but similar ages are not recorded from the crystalline basement in the area (see Clauer et al. 1991). One could, therefore, conclude that the rocks of the Akjoujt area were not part of the West African craton before the Variscan orogeny. Because of the lack of data and of detailed geological studies a definite statement on whether the Akjoujt area represents a part of the West African craton or not cannot be made. However, characteristic features of a Variscan or earlier suture zone, such as ophiolitic rocks or highpressure metamorphic rocks, are absent in the central Mauritanides (see Martyn & Strickland 2004). Based on this, it is proposed that the Akjoujt area possibly represents a parautochthonous terrane, which was deformed and thrust eastward during the Westphalian as a result of a Variscan collision elsewhere.

Geochemical composition of the hydrothermal IOCG mineralization at 2492 Ma The hydrothermal IOCG mineralization is controlled by D2 shear zones transecting the metacarbonate and by enveloping shear zones at the contact of meta-carbonate and biotite-actinolite schists of the Akjoujt meta-basalt unit (Fig. 2b). P–T conditions are those of the regional M2 metamorphism at 410 + 30 8C and 2–3 kbar (Kolb et al. 2006), which was dated at c. 2492 Ma (hydrothermal monazite U –Pb concordia age; Meyer et al. 2006). Massive sulphide ore developed in breccias within the meta-carbonate as a result of the brittle nature of Fe –Mg carbonate under upper greenschist-facies conditions as a result of structurally controlled hydrothermal fluid flow in zones of high permeability (Kolb et al. 2006). The resulting ore mineral paragenesis comprises magnetite, pyrrhotite, chalcopyrite, Fe– Co –Ni arsenides, arsenopyrite, cobaltite, uraninite and Bi–Au–Ag –Te minerals, and consequently is characterized by an enrichment of Fe, Ni, Co, Cu, Mo and As, as is recorded by bulk-rock analysis of the ore breccia (Table 7). Elements such as Ti, Cr and V, which are commonly incorporated in magnetite, show very low values because of the pure chemical nature of the magnetite. The

brecciated Fe–Mg carbonate variety is zoned, with a Mg-rich core and Fe-rich rim, and differs from the original undeformed Fe –Mg carbonate. This provides further evidence for the suggested process of brecciation associated with intense fluid –rock interaction. The brecciated Fe –Mg carbonate additionally shows a marked depletion in Rb, Sr and Ba (Table 9) and a distinct variation in the REE (Fig. 6). The positive Eu anomaly and the negative Ce anomaly is interpreted as a result of interaction between the brecciated Fe–Mg carbonate and a hydrothermal ore fluid of a relatively oxidizing nature at a temperature .250 8C (see Bau & Mo¨ller 1992). An oxidizing nature of the fluids is corroborated by the presence of the oxides magnetite and uraninite in the ore paragenesis. The zoned brecciated Fe– Mg carbonate is deduced to be the result of the interaction with the oxidizing ore fluid, because Fe –Mg carbonates form a perfect solid solution (Chang et al. 1998) and the meta-carbonate represents a virtually monomineralic rock. The Fe-rich rim of the Fe –Mg carbonate is explained by a change to a reducing, Fe-rich ore fluid composition. The REE pattern of the brecciated Fe –Mg carbonate, in particular the enrichment of HREE over LREE, and the presence of hydrothermal monazite and xenotime in the alteration paragenesis (Meyer et al. 2006) indicate that the REE were mobile and were added by the hydrothermal ore fluid from an external source. complexes in aqueous The (H)REE favour CO22 3 solution at hydrothermal conditions (Bau & Mo¨ller 1993). Therefore, we suggest that the ore fluid contained significant CO2 in solution, which possibly was derived from the breakdown of Fe –Mg carbonate during alteration. The Fe– Mg clinoamphibole–chlorite schists locally contain massive sulphide breccias with an ore assemblage similar to the one found in the brecciated meta-carbonate. Consequently, the Fe –Mg clinoamphibole–chlorite schists are also enriched in Ni, Co, As, Cu, Au and Bi (Fig. 7b, Table 6). The NASC-normalized data shows that K, Rb, Sr and Ba are variably depleted (Fig. 7b). It is, however, not clear whether this is related to the hydrothermal overprint or to the depleted nature of the source region of these meta-sediments. The brecciated Fe– Mg carbonate data indicate that Rb, Sr and Ba were removed and, thus, at least partly mobilized during the hydrothermal alteration and mineralization event. The REE signature remains relatively unaffected compared with PAAS and NASC (Figs 5b and 7b), which means that the Fe–Mg clinoamphibole–chlorite schists acted as a closed system for REE during hydrothermal alteration. The wall rocks are variably affected by the D2/ M2 overprint. The rocks of the Sainte Barbe


volcanic unit were deformed at peak metamorphic, upper greenschist-facies conditions and do not show significant geochemical changes. The amphibolite of the Akjoujt meta-basalt unit was retrogressed and as a result the biotite – actinolite schist formed with abundant grunerite in a 40 m halo surrounding the meta-carbonate (Fig. 2; Kolb et al. 2006). Alteration during deformation of this rock included enrichment of K, Rb and Cu compared with the amphibolites (Fig. 7a), which is explained by the formation of biotite and a weak disseminated sulphide mineralization. The abundance of other chemical components such as Nb, Zr, Ti and, to a lesser extent, Sc was not affected, because of their immobility during the alteration process. The IOCG class to which the Guelb Moghrein deposit belongs represents a family of related mineral deposits containing low-titanium iron ore together with variable amounts of Cu, U, Au and REE (Strickland & Martyn 2002; Kolb et al. 2006). Differences in Au/Cu ratios, the distinctive but inconsistently developed Fe, Cu, Au, Co, Ni, As, LREE and U element association, and the variation of alteration assemblages from deposit to deposit indicate a variable ore fluid chemistry (Williams et al. 1999). Alteration associated specifically with Cu –Au-dominated and T . 400 8C deposits such as those of Guelb Moghrein tends to be characterized by Fe– K metasomatism (Twyerould 1997). Such a metasomatism is clearly recorded in the biotite –actinolite schists, the Fe– Mg clinoamphibole chlorite schists and the meta-carbonate at Guelb Moghrein. In some IOCG deposits, two different fluids are required to explain the oxidized nature of the Fe mineralization (magnetite, hematite) and the reduced nature of the sulphide assemblage (pyrrhotite) (Mark et al. 2000; Marschik & Fontbote´ 2001; Skirrow & Walshe 2002). The geochemical composition of the alteration and mineralization at Guelb Moghrein suggests that the fluid composition changed from an initially relatively oxidizing Mg-rich fluid responsible for forming the early magnetite-uraninite assemblage, to a relatively reducing, Fe-rich fluid where pyrrhotite and magnetite represent the stable oxide– sulphide assemblage. Whether this can be explained by mixing of two contrasting fluids or by a progressive change of fluid characteristics on its migration path must, however, remain unresolved.

Geochemical composition of the retrograde greenschist-facies overprint at 1742 Ma During regional D3/M3, the ore body at Guelb Moghrein was displaced and the area was affected by a retrograde metamorphic overprint at temperatures of c. 300 8C (Kolb et al. 2006). Hydrothermal

73

monazite that formed during this event in the D3 shear zones was dated by the U –Pb concordia method at c. 1742 Ma (Meyer et al. 2006). The retrogression of the rocks of the Sainte Barbe volcanic unit involved the destruction of feldspar with depletion in Ba and Sr and the replacement of biotite by chlorite with concomitant K depletion (Tables 1 and 2). Similar metamorphic mineral reactions in the rocks of the Akjoujt metabasalt unit produced the chlorite schists, which dominate in the eastern open pit area. Mobile elements during the retrogression involved K, Rb, Sr and the LREE (Tables 1 and 2; Figs 4a and 7a). The presence of hydrothermal monazite and the LREE depletion in the chlorite schists compared with the amphibolites of the Akjoujt meta-basalt unit indicates REE redistribution by the hydrothermal fluid. Geochemical changes are not obvious in the meta-carbonate and the Fe– Mg clinoamphibole –chlorite schist, but discrete shear zones were formed especially in the rheologically weak Fe – Mg clinoamphibole–chlorite schist where the clinoamphibole was significantly reduced in grain size and in parts replaced by talc (Kolb et al. 2006; Meyer et al. 2006).

Conclusions The geology around the Guelb Moghrein IOCG deposit resembles a typical Archaean greenstone succession of pre-2492 Ma age. The bimodal metavolcanic suite together with typical local metapelite occurrences formed in an active volcanic arc setting at the western boundary of the West African craton. The immediate host to the mineralization is an assemblage of meta-carbonate rocks made up dominantly by Mg-rich Fe –Mg carbonate and Fe-rich meta-pelites represented by Fe–Mg clinoamphibole –chlorite schists. Our preferred interpretation is that the rocks formed as a carbonate-facies iron formation with intercalated shales. The succession was deposited on the continental shelf of the West African craton with the terrigeneous components probably being derived from the volcanic arc rocks. Subsequent deformation and metamorphism overprinted the lithologies at upper greenschist- to amphibolite- facies conditions. At 2492 Ma, North- and NW-directed D2 thrusting at c. 400 8C and 2–3 kbar caused brecciation of the Fe –Mg carbonate in the meta-carbonate and the formation of the IOCG ore breccia. This involved strong metasomatic overprinting with the hydrothermal transport of Fe, Mg, K, Rb, Sr, Ba, Ni, Co, Cu, Bi, Mo, As, Au and REE. The aqueous– carbonic hydrothermal ore fluid was initially Mg-rich and oxidizing and changed later to an Fe-rich composition and reducing conditions.

74

J. KOLB ET AL.

In a third tectonometamorphic event at c. 1742 Ma, thrusting to the east overprinted all lithologies in the lower greenschist facies. Hydrothermal fluid flow through the shear zones mobilized REE and precipitated monazite. This all testifies to the complex geological history of the West African craton. The craton underwent tectonothermal reactivation in the Late Archaean–Early Proterozoic and in the Mid-Proterozoic before collision between Laurentia and Gondwana during the Variscan orogeny finally consolidated the terrane. The authors would like to thank M. El Moctar O. M. El Hacen (Deputy General Manager, GEMAK, Nouakchott) and A. O. A. dit Ebaye (Director MORAK, Akjoujt) for their support in Mauritania and the permission to publish the results of this study. Thorough reviews from A. Wilde and T. De Putter and comments by J.-P. Lie´geois helped greatly to improve the manuscript. This study was made possible through grant Me 1425/ 6-1/2 of the Deutsche Forschungsgemeinschaft.

References ANONYMOUS 2006. Guelb Moghrein, Fact Sheet, 2006. First Quantum Minerals Ltd, Vancouver. B A G ATTA , A. 1982. Contribution a` l’e´tude geólogique et mine´ralogique du gisement d’Akjoujt, Mauritanie. PhD thesis, Universite´d’ Orleáns. B AU , M. & D ULSKI , P. 1996. Distribution of yttrium and rare-earth elements in the Penge and Kuruman IronFormations, Transvaal Supergroup, South Africa. Precambrian Research, 79, 37–55. B AU , M. & M O¨ LLER , P. 1992. Rare earth fractionation in metamorphic hydrothermal calcite, magnesite and siderite. Journal of Mineralogy and Petrology, 45, 231– 246. B AU , M. & M O¨ LLER , P. 1993. Rare earth element systematics of the chemically precipitated component in Early Precambrian iron formations and the evolution of the terrestrial atmosphere– hydrosphere– lithosphere system. Geochimica et Cosmochimica Acta, 57, 2239–2249. B AULUZ , B., M AYAYO , M. J., F ERNANDEZ -N IETO , C. & L OPEZ , J. M. G. 2000. Geochemistry of Precambrian and Paleozoic siliciclastic rocks from the Iberian Range (NE Spain): implications for source-area weathering, sorting, provenance, and tectonic setting. Chemical Geology, 168, 135– 150. C HANG , L. L. Y., H OWIE , R. A. & Z USSMAN , J. (eds) 1998. Non-silicates: Sulphates, Carbonates, Phosphates, Halides. Geological Society, London, Rockforming Minerals, 56. C LAUER , N., D ALLMEYER , R. D. & L E´ CORCHE´ , J. P. 1991. Age of the late Paleozoic tectonothermal activity in northcentral Mauritanide, West Africa. Precambrian Research, 49, 97– 105. D ALLMEYER , R. D. & L E´ CORCHE´ , J. P. 1989. 40Ar/39Ar polyorogenic mineral age record within the central Mauritanide orogen, West Africa. Geological Society of America Bulletin, 101, 55–70.

W IT , M. J. 1998. On Archean granites, greenstones, cratons and tectonics: does the evidence demand a verdict? Precambrian Research, 91, 181– 226. F AURE , K., H ARRIS , C. & W ILLIS , J. P. 1995. A profound meteoric water influence on carbonate genesis in the Permian Waterberg Coalfield, South Africa: evidence from stable isotopes. Journal of Sedimentary Petrology, 65, 605– 613. G NANESHWAR R AO , T. & N AQVI , S. M. 1995. Geochemistry, depositional environment and tectonic setting of the BIF of late Archaean Chitradurga schist belt, India. Chemical Geology, 121, 217 –243. H ECHT , L., F REIBERGER , R., G ILG , H. A., G RUNDMANN , G. & K OSTITSYN , Y. A. 1999. Rare earth element and isotope (C, O, Sr) characteristics of hydrothermal carbonates: genetic implications for dolomite-hosted talc mineralization at Go¨pfersgru¨n (Fichtelgebrirge, Germany). Chemical Geology, 155, 115–130. H ITZMAN , M. W. 2000. Iron oxide–Cu– Au deposits: What, where, when and why. In: P ORTER , T. M. (ed.) Hydrothermal Iron Oxide Copper– Gold and Related Deposits: A Global Perspective. Australian Mineral Foundation, Adelaide, 9– 25. I RVINE , T. N. & B ARAGAR , W. R. A. 1971. A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Sciences, 8, 523– 548. K HAN , R. M. K., D AS S HARMA , S., P ATIL , D. J. & N AQVI , S. M. 1996. Trace, rare-earth element, and oxygen isotopic systematics for the genesis of banded iron-formations: evidence from Kushtagi schist belt, Archaean Dharwar Craton, India. Geochimica et Cosmochimica Acta, 60, 3285– 3294. K OLB , J., S AKELLARIS , G. A. & M EYER , F. M. 2006. Controls on hydrothermal Fe oxide–Cu–Au–Co mineralization at the Guelb Moghrein deposit, Akjoujt area, Mauritania. Mineralium Deposita, 41, 68–81. L E´ CORCHE´ , J. P., D ALLMEYER , R. D. & V ILLENEUVE , M. 1989. Definition of tectonostratigraphic terranes in the Mauritanide, Bassaride, and Rokelide orogens, West Africa. In: DALLMEYER , R. D. (ed.) Terranes in the Circum-Atlantic Paleozoic Orogens. Geological Society of America, Special Papers, 230, 131–144. M ARK , G., O LIVER , N. H. S., W ILLIAMS , P. J., V ALENTA , R. K. & C ROOKES , R. A. 2000. The evolution of the Ernest Henry Fe-oxide– (Cu–Au) hydrothermal system. In: P ORTER , T. M. (ed.) Hydrothermal Iron Oxide Copper–Gold and Related Deposits: A Global Perspective, Australian Mineral Foundation, Adelaide, 123 –136. M ARSCHIK , R. & F ONTBOTE´ , L. 2001. The Candelaria – Punta del Cobre iron oxide Cu–Au(–Zn– Ag) deposits, Chile. Economic Geology, 96, 1799–1826. M ARTYN , J. E. & S TRICKLAND , C. D. 2004. Stratigraphy, structure and mineralization of the Akjoujt area, Mauritania. Journal of African Earth Sciences, 38, 489–503. M C L ENNAN , S. M. & T AYLOR , S. R. 1991. Sedimentary rocks and crustal evolution: tectonic setting and secular trends. Journal of Geology, 99, 1– 21. M C L ENNAN , S. M., M C C ULLOCH , M. T., T AYLOR , S. R. & M AYNARD , J. B. 1990. Geochemical and Nd–Sr isotopic composition of deep-sea turbidites: crustal DE

GUELB MOGHREIN MINERALIZATION evolution and plate tectonic associations. Geochimica et Cosmochimica Acta, 54, 2015–2050. M ESCHEDE , M. 1986. A method of discriminating between different types of mid-ocean ridge basalts and continental tholeiites with the Nb– Zr– Y diagram. Chemical Geology, 56, 207 –218. M EYER , F. M., K OLB , J., S AKELLARIS , G. A. & G ERDES , A. 2006. New ages from the Mauritanides Belt: recognition of Archean IOCG mineralization at Guelb Moghrein, Mauritania. Terra Nova, 18, 345–352. M OORE , S. E., F ERRELL , R. E. J. & A HARON , P. 1992. Diagenetic siderite and other ferroan carbonates in a modern subsiding marsh sequence. Journal of Sedimentary Petrology, 62, 357–366. M ORTIMER , R. J. & C OLEMAN , M. L. 1997. Microbial influence on the oxygen isotopic composition of diagenetic siderite. Geochimica et Cosmochimica Acta, 61, 1705–1711. M OZLEY , P. S. 1989. Relation between depositional environment and the elemental composition of early diagenetic siderite. Geology, 17, 704–706. M OZLEY , P. S. & B URNS , S. J. 1993. Oxygen and carbon isotopic composition of marine carbonate concretions: an overview. Journal of Sedimentary Petrology, 63, 73– 83. M OZLEY , P. S. & C AROTHERS , W. W. 1992. Elemental and isotopic composition of siderite in the Kuparuk Formation, Alaska: effect of microbial activity and water– sediment interaction on early pore-water chemistry. Journal of Sedimentary Petrology, 62, 681– 692. M OZLEY , P. S. & W ERSIN , P. 1992. Isotopic composition of siderite as an indicator of depositional environment. Geology, 20, 817–820. P ARTINGTON , G. A. & W ILLIAMS , P. J. 2000. Proterozoic lode gold and (iron)– copper–gold deposits: a comparison of Australian and global examples. In: H AGEMANN , S. G. & B ROWN , P. E. (eds) Gold in 2000. Society of Economic Geologists Reviews, 13, 69– 101. P ONSARD , J. F., R OUSSEL , J., V ILLENEUVE , M. & L ESQUER , A. 1988. The Pan-African orogenic belt of southern Mauritanides and northern Rokelides (southern Senegal and Guinea, West Africa): gravity evidence for a collisional suture. Journal of African Earth Sciences, 7, 463–472. P OUCLET , A., G UILLOT , P.-L. & B A G ATTA , A. 1987. Nouvelles donneés lithostructurales, pe´trographiques, mine´ralogiques et geochimiques sur le gisement de

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cuivre d’Akjoujt et son environnement geólogique (Re´publique Islamique de Mauritanie). Journal of African Earth Sciences, 6, 29–43. R ADVANEC , M., G RECULA , P. & Z A´ K , K. 2004. Siderite mineralization of the Germericum superunit (Western Carpathians, Slovakia): review and a revised genetic model. Ore Geology Reviews, 24, 267– 298. S ILLITOE , R. H. 2003. Iron oxide– copper-gold deposits: an Andean view. Mineralium Deposita, 38, 787– 812. S KIRROW , R. G. & W ALSHE , J. L. 2002. Reduced and oxidized Au– Cu–Bi iron oxide deposits of the Tennant Creek Inlier, Australia: an integrated geologic and chemical model. Economic Geology, 97, 1167– 1202. S PO¨ TL , C. & V ENNEMANN , T. W. 2003. Continuousflow IRMS analysis of carbonate minerals. Rapid Communications in Mass Spectrometry, 17, 1004– 1006. S TRICKLAND , C. D. & M ARTYN , J. E. 2002. The Guelb Moghrein Fe-oxide copper–gold–cobalt deposit and associated mineral occurrences, Mauritania: A geological introduction. In: P ORTER , T. M. (ed.) Hydrothermal Iron Oxide Copper– Gold and Related Deposits: A Global Perspective, Vol. 2. PGC Publishing, Adelaide, 275– 291. T WYEROULD , S. C. 1997. The geology and genesis of the Ernest Henry Fe– Cu– Au deposit, NW Queensland, Australia. PhD thesis, University of Oregon, Eugene. VAN K RANENDONK , M. J., W EBB , G. E. & K AMBER , B. S. 2003. Geological and trace element evidence for a marine sedimentary environment of deposition and biogenicity of 3.45 Ga stromatolitic carbonates in the Pilbara Craton, and support for a reducing Archaean ocean. Geobiology, 1, 91– 108. V ILLENEUVE , M. 2005. Paleozoic basins in West Africa and the Mauritanide thrust belt. Journal of African Earth Sciences, 43, 166 –195. W ILLIAMS , P. J., G UOYI , D., P OLLAND , P. J., P ERRING , C. S., R YAN , C. G. & M ERNAH , T. P. 1999. Fluid inclusion geochemistry of Cloncurry (Fe)–Cu– Au deposits. In: S TANLEY , C. J. ET AL . (eds) Mineral Deposits: Processes to Processing. Balkema, Rotterdam, 111–114. W INCHESTER , J. A. & F LOYD , P. A. 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chemical Geology, 20, 325 –343.

Possible primary sources of diamond in the North African diamondiferous province M. KAHOUI1, Y. MAHDJOUB1 & F. V. KAMINSKY2 1

Faculte´ des Sciences de la Terre, de la Geógraphie et de l’Ameńagement du Territoire, USTHB, BP 32, Algiers, 16111, Algeria (e-mail: [email protected])

2

KM Diamond Exploration Ltd., 2446 Shadbolt Lane, West Vancouver, B. C., V7S 3J1, Canada Abstract: The Eglab shield is the easternmost part of the Reguibat rise, which belongs to the West African craton (WAC). It corresponds to the amalgamation of the Yetti and Eglab Palaeoproterozoic domains. These domains are separated by a mega-shear zone called the ‘Yetti–Eglab Junction’ where fieldwork has led to the discovery of kimberlite indicator minerals but no diamond. In the southwestern part of this zone, an outcrop of Archaean basement and a komatiitic–picritic dyke had been recognized. Within the Eglab shield, deep-seated lithospheric faults control emplacement of alkaline complexes, and of small circular structures made up of mafic, ultramafic and silica-undersaturated rocks. These structural zones are characterized by widespread development of dyke swarms and repeated reactivations of earlier Eburnean trends from the Neoproterozoic to Mesozoic. Accordingly, they are sites of high magmatic permeability and crustal weakness. In this study, we summarize all known earlier and newly obtained structural, geophysical, geological and geochemical data on this area. They indicate that the ‘Yetti–Eglab Junction’ has good possibilities for the finding of kimberlite or/and other diamondiferous rocks. The features of the Eglab shield provide a possible explanation for the enigmatic sources of the diamond-bearing Reggane placer deposit located at the boundary of the WAC.

The first documented discovery of diamond from northern Africa was in 1953 by M. Ranoux. The diamond was found in a sample of sand collected probably from the In-Hihaou wadi in the western part of the Tuareg (Ahaggar) shield, Algerian Sahara (Fig. 1). This was a rounded crystal, about 40 mg in weight, associated with zircon, amphibole, martite and other minerals; no kimberlite indicator minerals (KIM) were recognized in the sample (The´bault 1959). In 1969, V. Izarov found a second diamond within the Tuareg shield, this time from within its eastern part, in the Tiririne area. This crystal was a small (0.36 mm) rhombic dodecahedral diamond that was recovered from a sample collected in eluvium of the red-stone conglomerate, part of the Neoproterozoic Tiririne Suite (Izarov & Biroutchev 1974). Along with this diamond, pyrope garnet and magnesian ilmenite (‘picroilmenite’) grains occurred in the same sample. In the 1970s–1980s, prospecting for diamond deposits in Algeria moved to the west and NW, and diamonds were found within a large area extending for almost 300 km from Tanezrouft, in the south, to El Kseibat in the north. In the central part of this area, a sub-economic Djebel Aberraz placer deposit was discovered in the Bled El Mass valley, some 30 km south of Reggane (Touahri et al. 1996). Here, under a few metres of aeolian sand, Lower–Upper Quaternary alluvial deposits

overlie Palaeozoic sedimentary rocks. In these alluvial sediments, which are up to 12–15 m in thickness, hundreds of diamonds were discovered. They are characterized mainly by dodecahedral and transitional crystal forms, and more rarely (c. 35%) by octahedra. Diamond crystals bear evidence of mechanical erosion in ancient coastal – marine and recent fluvial environments (Kaminsky et al. 1990). Along with diamond, in diamondiferous Quaternary sediments, numerous KIM, such as pyrope garnet, chrome-spinel and picroilmenite were discovered. These grains are well rounded; like the diamonds, they have undergone a long history of transportation. Grains of pyrope were found in Cretaceous sediments of the Tanezrouft Plateau, c. 10 km north of the Bled El Mass deposit (Kaminsky et al. 1992a; Sobolev et al. 1992). The geological position of the Bled El Mass diamond deposit and other diamond localities in the Algerian Sahara is ambiguous. This deposit borders the Sahara plate and West African craton (WAC). Diamond and KIM from it bear evidence of a long transportation history. They are not related to any known diamondiferous sources, and may form a new, North African diamondiferous province (Kaminsky et al. 1992b). In the 1990s, we suggested the possibility of finding primary sources to the Algerian diamonds in the Eglab shield, which is the easternmost part


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Fig. 1. Geological representation of the West African craton (WAC) and surrounding belts, with the location of the study area highlighted. The main structural features are the Archaean and Palaeoproterozoic domains surrounded by Neoproterozoic and Phanerozoic belts (WAC, West African craton; TS, Tuareg shield; NS, Nigerian shield). Northern and northeastern boundaries of the WAC are after Ennih & Lie´geois (2001), faults and discontinuities derive from the geological and geophysical data of and Roussel & Lesquer (1991) and Fabre (2005); E-Ch L, Erg Chech Line; El-Mah L, El Mahdi Line; AGL, Adrar–Guinea Line. Kimberlitic, ultramafic and ultrapotassic formations and diamondiferous occurrences are noted. Kimberlitic and diamond-rich areas for the Man shield are after Pouclet et al. (2004). Ad, Adrar; Tam, Tamanrasset; Ta, Taoudeni; Ti, Tindouf.

NORTH AFRICAN DIAMONDIFEROUS PROVINCE

of the Precambrian Reguibat rise of the WAC (Kahoui 1991; Kahoui & Benameur 1992; Kahoui & Mahdjoub 2001; Kahoui et al. 2004; Fig. 1) as also advocated by Gevin (1958) and The´bault (1959). Our hypothesis was based on: (1) the identification of deep-seated lithospheric faults (with associated magnetic anomalies) controlling the emplacement of mafic and ultramafic rocks, and (2) the presence of alkaline igneous ring complexes. In addition, many small circular structures (typically 250 m in diameter), a great number of lamprophyric and doleritic dykes, and our first pyrope garnet and ruby were found within the area (Figs 2–5). As the result of a new phase of work by geologists from the ORGM (Office National de la Recherche Geólogique et Minie`re) in 2000–2001 within the ‘Yetti–Eglab Junction’ (Sabate´ 1973) (Fig. 2), pyrope garnet, picroilmenite and chromediopside mineral dispersion haloes were identified (Labdi & Zeńia 2001; see Fig. 5). The choice of this region for diamond exploration was reinforced by our recent discovery of a mafic dyke of ‘komatiitic–picritic’ affinity that cross-cuts an Archaean relict. We note also that, in French Guiana, diamondiferous volcanoclastic komatiites have been recognized (Capdevila et al. 1999). The Eburnean shield is characterized by an average heat flow of 30 + 10 m Wm22. The thickness of the West African lithosphere is 150 –200 km (Roussel & Lesquer 1991). These features are typical for cratonic domains and favourable for the formation and conservation of diamond in their mantle lithospheric roots. On the Eglab shield, no heat flow values are available, but to the west, in Mauritania (Zouerate area), values range from 43 to 52 mW m22, with an average of 50 mW m22, which ‘confirm that the northern part of the WAC is characterized by heat flow density (HFD) values higher than the southern part, but that the difference is not very significant’ (Lesquer et al. 1991). To the north, in Morocco, on the Precambrian basement of the Anti-Atlas, which belongs to the WAC, the HFD is estimated at 40 mW m22 (Rimi 1999). Considering the West African Eburnean context and comparing it with other diamondiferous cratonic domains of the same age (e.g. Brazilian and Guianian cratons), the Eglab shield demonstrates opportunities for the discovery of kimberlite and/ or other diamondiferous rocks. The compilation of morphological structures, lineaments, geophysical data and existing prospecting results (identified KIM, alkaline and ultramafic rocks) is very useful in the selection of areas for diamond exploration (Kaminsky et al. 1995). The main goal of this work is to compile and analyse

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known and new geological, tectonic and geophysical data, to select the most prospective areas for a possible discovery of primary diamond source(s) in northern Africa.

Regional geology of the Eglab shield The Eglab shield is the easternmost part of the Reguibat rise, which is in turn the northern part of the WAC (Fig. 1). It is limited to the north by the Palaeozoic Tindouf basin, to the east by the dunes of Erg Chech and the Palaeozoic Reggane basin, and to the south by the Neoproterozoic Hank series and the Palaeozoic cover of the Taoudeni basin. This shield (Fig. 2) comprises Palaeoproterozoic terranes accreted during a major Palaeoproterozoic juvenile crust-forming event, which occurred between 2200 and 2070 Ma (Drareni et al. 1996; Peucat et al. 2005); the Palaeoproterozoic formations occur to the east of Archaean terranes located in the northern and southern WAC, within the Reguibat (Potrel et al. 1996, 1998; Chardon 1997) and Man shields (Fig. 1). The Eglab shield is defined as a proton, which is a part of the Earth’s crust that has attained stability and has experienced little deformation since the Early to Middle Proterozoic (Palaeoproterozoic) (Janse 1992). It is subdivided into two domains differing in their structural, lithological, stratigraphic and metamorphic characteristics: the Yetti domain to the west, and the Eglab domain to the east (Fig. 2). Both domains are intruded by granitoids of different ages and are separated by a mega-shear zone called the ‘Yetti– Eglab Junction’ (Sabate´ 1973); this zone corresponds to the amalgamation of the second domain with the first one (Lefort et al. 2004). The Yetti domain is a NNW–SSE-trending basin composed mainly of the Yetti series: volcanic (rhyolite, rhyodacite), volcano-sedimentary (tuffs) and sedimentary (quartzite, pelite, arkose, and conglomerate) units (Buffie`re et al. 1965a, b, 1966; Lameyre & Lasserre 1967; Lasserre et al. 1970). The Yetti series, which constitutes an envelope of a migmatitic dome, is cross-cut by post-orogenic Yetti granites, dated at 2073 Ma (Peucat et al. 2005). In the southwestern part of Eglab shield, 50 km NNW of Chegga (Fig. 2), an outcrop of Archaean basement has been recognized. This outcrop is formed by a series of amphibolites intercalated with garnet–hornblende banded grey gneisses, dated at 2.73 Ga (Peucat et al. 2005). This series, which may be considered as a relict of the Archaean core of the Eglab shield, is intruded by the Chegga granite, dated at 2.1 Ga (Peucat et al. 2005).

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Fig. 2. Schematic geological map of the Eglab shield compiled from Buffie`re et al. (1965a, b), the authors’ own geological fieldwork, interpretation of satellite images and aerial photographs. Ages are from Peucat et al. (2005).

The main part of Eglab domain is made up of the two following major units (Fig. 2). (1) A Lower Reguibat Complex (LRC) represented by (a) the Chegga series to the west and the Chenachane –Erg Chech series (granite– gneiss, migmatite, amphibolite) to the east of the Eglab domain, and (b) the Yetti series (rhyolite, greywacke, schist) in the eastern part of the Yetti domain (Gevin 1951, 1958; Sougy 1954, 1960; Buffie`re et al. 1965a, b; Buffie`re 1966). (2) An Upper Reguibat Complex (URC) represented by the Oued Souss series (Buffie`re et al. 1965a, b; Buffie`re 1966), the Akilet Deleil series (Sabate´ & Lameyre 1973), and the Guelb El Hadid series (Gevin 1951, 1958; Buffie`re et al. 1965a, b; Buffie`re 1966). The Oued Souss and Akilet Deleil series contain detrital (sandstone, arkose, conglomerate), calc-alkaline volcanic and volcano-sedimentary rocks. The Guelb El Hadid series is continental (arkose, sandstone, pink quartzite, arkosic sandstone, conglomerate) with some interbedded felsic volcanic rocks.

The Aftout magmatic suite, including felsic Aftout –Eglab volcanic rocks, mafic intrusions and large post-tectonic Aftout granitic plutons is associated with the URC. The Aftout –Eglab volcanic and plutonic rocks cross-cut or overlie the series of the LRC and URC, and cover nearly half of the present area of the URC. The alkaline –peralkaline Djebel Drissa ring complex belongs to these Aftout granitoids. The formations described above are overlain in the south by the marine and continental Neoproterozoic Hank series and intruded by doleritic and gabbro–doleritic dykes and/or sills; the stratigraphic position of these dykes and sills is discussed below. Three major Eburnean magmatic events are recognized in the Eglab shield, as follows: The first event corresponds mainly to 2.21 – 2.18 Ga magmatic activity that formed a metamorphosed batholith belonging to the LRC. Petrographical and geochemical features indicate two groups of magmatic rock suites, characterized


in the eastern part of Eglab (Erg Chech series) by the gabbroic Teggeur (u1 in Fig. 2) and the orthogneissic Tilemsi and Teggeur groups. These plutonic rocks form a juvenile calc-alkaline orogenic suite. They are compatible with active subduction in a continental active margin or a mature island arc setting. The lack of any significant Archaean Nd-isotopic signature argues for the recycling of young crustal components (Peucat et al. 2005). The second magmatic event at c. 2.09 Ga corresponds to the intrusion of a syntectonic trondhjemitic pluton (Chegga granite) into the Archaean relicts of the Chegga series and to a dacitic tuff of the Oued Souss series (Fig. 2). The Chegga syntectonic granitoids and the Oued Souss and Akilet Deleil volcanic series, mainly composed of rocks ranging from basaltic andesites to rhyolites, define a calc-alkaline suite with active-margin affinities. The third magmatic event at c. 2.07 Ga represents a large volume of high-K to peralkaline post-orogenic magmas (Aftout and Eglab magmatism), which are interpreted as resulting from an asthenospheric upwelling (Peucat et al. 2005). The mafic rocks related to this magmatic event occupy only a small area and consist of a suite ranging from olivine– hypersthene normative gabbros to oversaturated quartz gabbrodiorites; they crop out as small plutons, lenses and dykes (u2 in Fig. 2). In the Eglab shield, the age of the important doleritic and gabbro–doleritic sills and dyke swarms with various trends (Buffie`re et al. 1965a, b; Sabate´ & Lomax 1975) is not well constrained. These rocks intruded either the Eburnean basement or the sedimentary cover (Fig. 3). For dykes oriented north –south, N40, N130 and N160, it is not possible to precisely define the upper age limits, with the exception of those intruding the Lower Guelb El Hadid series and covered by the Upper series of this formation (Sabate´ & Lomax 1975). These dykes have for an upper limit the Lower Palaeozoic sandstones of the Tindouf basin, and could be associated with the magmatic event attributed to the pre-Pan-African continental margin extension. This magmatic activity is known in the Birimian formations in western Niger, the easternmost part of the Man shield (Ama Salah et al. 1996; Affaton et al. 2000) and more so in the north (El Ouali et al. 2001), in the Moroccan Anti-Atlas region; the latter corresponds to the northern boundary of the Eburnean WAC (Ennih & Lie´geois 2001). However, some dykes with the same trends can be traced throughout the Neoproterozoic Hank series and Palaeozoic series of the Tindouf basin and could suggest more probable recent ages (Palaeozoic or Mesozoic).

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Some dykes oriented N60 to N80 cross-cut the Neoproterozoic Hank series and the Palaeozoic sedimentary rocks of the Tindouf and Taoudeni basins (Villemur 1967; Sabate´ & Lomax 1975; Bertrand 1991; Sebai et al. 1991). They could be attributed to the tholeiitic magmatism that is estimated to have occurred at around 200 Ma (Sebai et al. 1991). This magmatism extends from Morocco through Algeria to the Ivory Coast and is injected along reactivated pre-existing fractures oriented NE–SW to ENE –WSW (Sebai et al. 1991). This reactivation of the pre-existing lithospheric structures (north–south and ENE –WSW) controlled, in Mali, the Tadhak alkaline magmatic event dated at 262, 215, 185 and 160 Ma (Lie´geois et al. 1991); this alkaline magmatism was synchronous with the tholeiitic event, and the major part of both is linked with the opening of the central Atlantic Ocean (Lie´geois et al. 1991). The reworking of these inherited structures is also indicated for the emplacement, in Mauritania, of the Cretaceous carbonatitic Richat structure, dated at 100 Ma (Poupeau et al. 1996). It seems clear in the Eglab shield, that without dating the important doleritic and gabbro–doleritic sill and dyke swarms, their emplacement could thus far be attributed to: the Eburnean magmatic events; the Pan-African continental margin extension magmatism; or the Mesozoic (Jurassic and/or Cretaceous?) magmatism.

Structural evolution of the Eglab shield The geology of the Eglab shield is dominated by NW–SE- to NNW–SSE-striking lithologies and regional foliation, as observed throughout the Reguibat rise and WAC (Fig. 2). The structural evolution of the Eglab shield is characterized by ductile and brittle deformations (Mahdjoub et al. 2002).

Ductile deformation Ductile deformation is characterized by north– south-striking foliations with a very strong vertical flattening and NNW–SSE to NW–SE-striking subvertical shear zones. The ductile deformation is coeval with major, early calc-alkaline stages of magmatic accretion (c. 2.2 and 2.09 Ga). The development of north–south-striking volcanosedimentary basins is synchronous with 2.09 Ga magmatic accretion during the second Eburnean stage, which is synchronous with the B2 volcanosedimentary basins of the Birimian part of the Man shield (Doumbia et al. 1998). The kilometre-scale NNW–SSE- to NW– SE-striking shear zone that separates the Yetti and Eglab domains (‘Yetti –Eglab Junction’) and

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Fig. 3. (a) Brittle fractures and their relationships with dykes, as defined from geological fieldwork and interpretation of satellite images and aerial photographs. D-shear is the direction of major shear corridors (AL-TC, Aouinet Legraa– Tilesmas Corridor; EDC, Eglab Dersa Corridor; KMC, Kahal Morrat Corridor). X-shear (Chenachane Corridor) is a conjugate shear of WSW–ENE trend (D-shear). Internal faults (R, R0 ) indicate a Riedel geometry. (b) Riedel model showing the geometric relationships between second-order Riedel faults (R, R0 ) within the sinistral major Eglab shear zone (D-shear zone); dykes are controlled by T-tensional fractures. The trend of the old sinistral transpressional Yetti–Eglab shear zone is reactivated during later transtensional stages. (c) Djebel Drissa emplacement model within the dextral X-shear zone or Chenachane Corridor.

discrete secondary shear bands show sinistral strikeslip and top-to-the-west thrusting components (Fig. 2). The relationships between foliations, strike-slip and associated compressional thrusts

indicate transpressive motions resulting from an east –west oblique convergence between the newly accreted Yetti and Eglab crustal blocks (Mahdjoub et al. 1994; Peucat et al. 2005).


Brittle deformation At map scale, the brittle structures (Fig. 3a and b) are represented by ENE –WSW (N70) and NW–SE (N120) conjugate sinistral and dextral faults, respectively. NW–SE-, NNW–SSE- and NNE-SSW-striking-faults represent discrete second-order faults. NNW–SSE, north–south and NNE–SSW felsic and mafic dykes form en echelon patterns. Minor strike-slip movements are associated with some of NNE –SSW-trending dykes (e.g. Djebel Drissa, Fig. 3c). North–South-striking compressional structures (reverse faults and folds) are documented from the Eglab shield and Taoudeni basin (Tokarski 1991; Bertrand-Sarfati et al. 1996). Here we describe only the extensional and transtensional deformations. Faults and fractures are located within prominent ENE– WSW-trending regional corridors (examples are the Aouinet Legraa– Tilesmas (Al– TC) and Kahal Morrat (KMC) corridors) (Fig. 3a), which correspond to major D-shears with sinistral displacement. The internal structures are characterized by discrete second-order faults with Riedel geometrical relationships to each other and to the major shears (Riedel 1929). NE– SW- and NNE– SSW-striking minor faults correspond to internal R-shear and R0 -shear, respectively (Fig. 3a and b). NW–SE-trending fault corridor X-shear (Chenachane corridor) corresponds to a later strikeslip reactivation of the early transpressional Eburnean structures. The brittle fracture corridors and associated mafic and felsic dykes are compatible with a transtensional bulk regime, following oblique Eburnean convergence. They controlled, at 2.07 Ga, the postcollisional emplacement of the Aftout granitoids, Aftout and Eglab volcanic rocks and continental sedimentary basins, filled by the detrital Guelb El Hadid formations (Mahdjoub et al. 1994; Peucat et al. 2005); this is an indication that final thermotectonic and sedimentary Eburnean events resulted from a reactivation of the ancient major ductile shear zones, during the post-collisional east– west-trending stage. The linear structures and faults appearing on the geological map of Buffie`re et al. (1965a, b) show that the earlier Eburnean trends have numerous reactivation histories in the following directions. (1) NW– SE and ENE –WSW conjugate faults with respectively sinistral and dextral displacements cutting the Neoproterozoic Hank sediments. Both are compatible with a compressional bulk regime during Pan-African collision between the WAC and Tuareg shield. (2) North – south to NNW–SSE faults (as the old deep structures) in the bordering Taoudeni and Tindouf basins, often extending out of the Eburnean

83

basement and reflecting Palaeozoic tectonic trends (Tokarski 1991). (3) ENE– WSW faults cutting the southern outcrops of Tindouf basin; this direction follows the northern Eglab shield border and major ENE –WSW trend of Aouinet Legraa –Tilesmas corridor. In the Taoudeni basin, the ENE –WSW trend (Fig. 1) corresponds to the direction of known surface and subsurface fractures and lineaments that controlled Cretaceous (100 Ma) emplacement of the annular carbonatitic structure of Richat in Mauritania (Netto et al. 1992; Poupeau et al. 1996). At the scale of the WAC, ENE – WSW and their associated NNW –SSE trends reflect Late Triassic – Early Jurassic rifting associated with the opening of the central Atlantic Ocean. An extensive episode of tholeiitic magmatism is related to this tectonic activity, and produced doleritic sills and dykes (Bertrand et al. 1982; Sebai et al. 1991; Deckart et al. 2005). The age of this magmatism (206 – 195 Ma) is well constrained from dykes, sills and associated lavas from Iberia, Morocco, Algeria and Mali (Sebai et al. 1991). The N80 direction is the main orientation of the Guinean –Nubian lineaments, which are deep structures followed from the Guinea margin to the Rea Sea. Sinistral transtensional movements and reworking of these pre-existing lithospheric zones of weakness during Early Cretaceous time appear to control the intra-plate magmatic activity (Guiraud et al. 1987; Netto et al. 1992; Maurin & Guiraud 1993; Poupeau et al. 1996).

Geophysical structures and relationships to structural features Within cratons, kimberlite and/or lamproite intrusions are usually located in zones of high magmatic permeability, corresponding to longlived deep-seated major faults. These zones of weakness provide channels for the ascent of mantle-derived magmas (Kaminsky et al. 1995; White et al. 1995). They are not always well reflected in the geological structures but can be traced by gravity and magnetic surveys. Geophysical survey methods, including gravity, magnetic, electromagnetic, resistivity and seismic techniques, are also very useful and can be successful in the location of primary diamond sources; they are largely used to delineate favourable areas for future diamond exploration (Macnae 1979; Morgan 1995). The following are interpretations of the structures of the WAC and the Eglab shield using geophysical data.

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Lithospheric deep faults and relationships with the major crustal features of the WAC NNW–SSE, NW –SE, ENE –WSW and NNE– SSW directions appear as the most striking features throughout the Reguibat rise and WAC as revealed by gravity, aeromagnetic and geological data (Roussel & Lesquer 1991; Fig. 1). NNW–SSE and NW– SE directions. The structure of the Eburnean basement of the Eglab shield is revealed by a succession of NNW–SSE shear zones (e.g. Yetti –Eglab shear zone) and by elongated high and low gravity anomalies (Roussel & Lesquer 1991). These major features are the result of transpression during oblique convergence between the Yetti and Eglab domains. The NNW–SSE and NW –SE faults divided the basement into crustal blocks and could be extrapolated under the post-Palaeoproterozoic sedimentary

cover. These deep structures controlled the Neoproterozoic –Palaeozoic and Palaeozoic –Mesozoic sedimentation in the Taoudeni and Tindouf basins, respectively (Bertrand-Sarfati et al. 1996; Moussine-Pouchkine & Bertrand-Sarfati 1997). ENE –WSW and NNE– SSW directions. On the Eburnean basement, ENE –WSW and NNE–SSW trends, controlling the distribution of the postcollisional Eglab volcanic units, were defined by aeromagnetic reinterpretation (Allek 2005). The southern parts of the Taoudeni and Tindouf basins show ENE– WSW gravity, aeromagnetic and structural trends, oblique to the early Palaeoproterozoic trends (Roussel & Lesquer 1991). The gravity trends appear coherent with WNW –ESE trans-extensional features of the Reguibat rise. Gravity and geological interpretations (Lesquer et al. 1984) indicate the ENE –WSW and

Fig. 4. Aeromagnetic lineaments and location of small-sized circular aeromagnetic anomalies within the Eglab shield. Limits of large recognized or supposed circular structures are from EREM (1983); small-sized anomalies and lineaments derive from EREM (1983) and Allek (2005).


NNE–SSW limits of sub-basins in the Taoudeni basin, which are filled with Neoproterozoic and Palaeozoic deposits (Villeneuve & Corneé 1994; Villeneuve 2005). These facts may support the assumption that the evolution of the Taoudeni and Tindouf basins was controlled by ENE –WSW and crustal zones of weakness inherited from Proterozoic deformational events (Villemur 1967). Roussel & Lesquer (1991) noted that the ENE – WSW transverse zone, well defined by geophysical and geological trends (Adrar Guinea line), may be regarded as evidence for the existence of a wide Proterozoic mobile belt reactivated throughout geological time.

Aeromagnetic anomalies and their relationships with the major structural features in the Eglab shield A qualitative interpretation (Entreprise Nationale de Recherche Minie`re, EREM 1983) of aeromagnetic data (Aeroservice Corporation 1974) based on DT anomalies and compared with geological mapping of the Eglab shield shows two groups of circular aeromagnetic structures (Fig. 4). The first group reflects large regional structures (50–100 km in diameter); the second group includes narrow to moderate local-scale structures (2–10 km in diameter). The first group of circular aeromagnetic structures corresponds to recognized or supposed large geological circular structures, with good examples located within the ‘Yetti –Eglab Junction’, to the west, and the Chenachane –Kahal Morrat area to the east of the Eglab domain. The recognized structures are well reflected in satellite imagery; the largest ones correspond to the limits of the granitic batholiths. The second group is located within the limits of the large recognized or supposed circular structures. This group is characterized by high (.3000 nT), moderate (300–3000 nT) and low (,300 nT) magnitudes. Anomalies with high intensities are rare and correspond probably to ultramafic rocks that do not crop out. Some moderate and low anomalies have annular forms, are superimposed on annular geological structures observed in the field, and are preferentially associated with the Yetti –Eglab Junction in the Bled M’Dena– Akilet Deleil area (Figs 4 and 5). Near Aouinet Legraa, in the limit of the northern Eburnean basement with the Palaeozoic Tindouf basin, two small anomalies are identified and are inferred to be associated with mafic dykes; further to the east, within this basin, one important anomaly is associated with a deposit of iron (Allek 2005).

85

In the Bled M’Dena–Akilet Deleil area two second-order trends control the aeromagnetic anomalies: the first, oriented NNE– SSW, is compatible with the direction of regional extension; the second is oriented SE–NW, with a transtensional dextral displacement (X fracture) (Fig. 5). In this region, four high anomalies are detected (Allek 2005) at the intersection of NNW– SSE and NE –SW faults; one anomaly (860 nT) is exactly superimposed on the ‘Anna’ annular structure (EREM 1983; Allek 2005) where KIM have been discovered (Labdi & Zeńia 2001; Fig. 5). The ground magnetic survey on the ‘Anna’ structure reveals the presence of two small, highly magnetic, anomalies, which are inferred to be associated with the presence of ultramafic bodies. In addition to the ‘Yetti– Eglab Junction’, four structural sites show isolated and/or cluster anomalies (Fig. 5). (1) The sinistral ENE –WSW Kahal Morat– Eglab Dersa corridor (KMC–EDC; D-shear). (2) The dextral NW –SE Chenachane corridor. This corridor, reactivated throughout geological time, controls the emplacement of the Eburnean alkaline peralkaline Djebel Drissa ring complex, the displacement of the Hank series and the sedimentation of Cenozoic Hamada limestones. In the Chenachane wadi, to the south of this corridor, seven grains of pyrope have been discovered, but no results have yet been published regarding these garnets. We note that at the intersection of this corridor with north –south and NNE –SSW faults, two high magnetic anomalies are detected, and two areas have been selected (not represented in Fig. 5) for diamondiferous exploration (Allek 2005). (3) At the intersection of ENE –WSW Kahal Morrat– Eglab Dersa –Aouinet Legraa –Tilesmas corridors with the north–south faults and the NNW –SSE, Chenachane corridor. (4) At the intersection of R and R0 (e.g. west of the Djebel Drissa). Within the Neoproterozoic Hank series and the Cenozoic Hamada limestones, in the south, moderate- and low-magnitude anomalies are located. They are situated on the same structural trend (NW–SE) to the ‘Anna’ structure or X fractures. The structural model proposed above shows that the location of circular aeromagnetic anomalies appears to be controlled by later extensional or strike-slip post-Eburnean or pre-Pan-African tectonics. The geometric relationships between the Riedel fractures and the regional distribution of the anomalies within Palaeoproterozoic basement, Neoproterozoic and Palaeozoic series and the Cenozoic Hamada series suggest a

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Fig. 5. Relationships between Riedel-fracture model, intrusions and small-sized circular aeromagnetic anomalies. EDC, Eglab Dersa Corridor; KMC, Kahal Morrat Corridor. D-shear, major shear-corridor; R and R0 , internal Riedel faults; X-shear, Chenachane Corridor. I, ‘Malignite’ circular structure; II, Merroucha circular structure; III, Chegga ‘komatiite– picrite’; IV, Anna circular structure; V, Bled M’Dena circular structure. Data relating to kimberlite indicator minerals are from Labdi & Zeńia (2001): (1) pyrope garnet þ ilmenite þ spinel; (2) pyrope garnet þ picroilmenite þ chrome-diopside; (3) pyrope garnet þ picroilmenite.

complex reactivation history for the older basement features. Good examples of Riedel fault patterns controlling the location of kimberlites, lamproites and carbonatites have been described in the Halls Creek Mobile Zone and Kimberley Block (Australia), the Lucapa corridor (Angola) and the Yengema area in Sierra Leone (White et al. 1995). In the WAC, extensional tectonics occurred several times during the Phanerozoic, mainly in the Early Jurassic and Early Cretaceous. In the Cretaceous period there appear to have been a good likelihood for the emplacement of alkaline, ultramafic and ultrapotassic magma derived from deep mantle sources (Netto et al. 1992; Poupeau et al. 1996).

Alkaline, mafic and ultramafic intrusions and relationships with the major structural features of the Eglab shield Compilation of geological, structural, and geophysical features and distribution of KIM (see below) show that the most favourable areas for diamond exploration within the Eglab shield are (1) the long-lived Chenachane shear zone, and (2) the Yetti– Eglab Junction and its neighbouring Yetti domain. In the first area, the major deep-seated Chenachane shear zone controlled the emplacement of the alkaline–peralkaline Djebel Drissa ring complex and was reactivated after the deposition

Table 1. Chemical composition of circular intrusions and dykes Rock:

‘komatiite’ GH28

Sample: 44.4 0.3 8.6 8.14 0.13 16.1 7.37 0.21 0.54 0.19 12.94 98.92

na, not analysed.

Doleritic and gabbroic dykes

Lamprophyric dykes

4-3/4

GH19

GH20a

GH20b

GH22

30B

35

38H

38G

25C

25B

25A

14

48.73 0.40 8.19 7.06 0.12 11.38 17.31 3.80 0.75 0.14 0.76 98.65

48.91 0.36 8.52 6.50 0.12 11.43 17.43 3.81 0.91 0.11 1.01 99.10

49.32 0.40 8.52 7.02 0.13 11.41 17.54 3.90 0.80 0.14 0.83 100.01

55.99 0.68 16.49 8.2 0.12 4.07 6.12 4.09 1.53 0.26 2.38 99.93

55.32 0.59 13.7 8.42 0.12 7.21 7.74 3.18 1.07 0.23 2.16 99.74

55.71 0.6 13.42 8.42 0.14 7.31 7.62 2.97 1.1 0.21 1.94 99.44

56.69 0.75 14.39 11.74 0.19 4.5 7.54 2.96 0.27 0.09 0.64 99.76

52.61 0.76 14.99 9.64 0.17 6.11 10.48 2.31 1.01 0.11 1.67 99.86

50.34 1.67 15.22 10.44 0.14 5.51 6.21 3.85 1.05 0.43 4.88 99.74

55.46 0.75 16.52 8.11 0.15 4.62 5.05 4.25 3.1 0.35 1.58 99.94

54.56 0.55 13.66 9.58 0.17 7.7 8.37 3.39 1.11 0.19 1.21 100.49

53.95 0.46 12.59 7.81 0.12 7.96 5.38 2.03 1.35 0.23 8.03 99.91

55.99 0.64 15.06 8.81 0.1 5.72 5.25 3.34 1.09 0.2 7.74 99.9

48.84 0.54 13.03 9.16 0.14 5.97 9.18 1.8 1.09 0.17 10.03 99.95

55.08 0.89 14.42 10.92 0.17 4.89 8.12 2.27 2.05 0.13 0.99 99.93

153.1 392.8 54.6 9.6 21.5 508.8 4.6 1.1 150.0 17.3 41.6 25.3 4.5 1.2 3.1 2.0 0.9 0.1 0.8 0.1 0.2 0.4 9.7 43.5

139.0 56 13.1 19.8 39.1 506 4.2 0.96 380 15.9 36.94 18.37 4.09 1.16 3.18 2.67 1.63 0.223 1.37 0.26 0.28 1.77 15.3 122

159.0 323.3 145.2 9.7 20.4 542.6 4.5 0.9 180.7 16.3 40.1 25.9 4.6 1.2 3.1 2.0 0.9 0.1 0.8 0.1 0.2 0.4 9.5 44.3

147.4 301.8 89.3 8.7 26.5 497.9 2.5 1.6 180.7 12.9 31.7 21.1 4.0 1.0 2.7 1.7 0.8 0.1 0.7 0.1 0.1 0.3 8.0 36.4

169.0 304.0 51.3 16.9 24.1 542.0 2.9 0.7 354.0 11.9 24.9 14.7 2.9 1.0 2.6 2.2 1.1 0.2 1.1 0.2 0.2 3.1 12.3 79.3

182.0 381.0 55.0 17.0 28.0 600.0 4.0 na 422.0 na na na na na na na na na na na na 0.0 14.0 85.0

268.0 334.0 34.0 14.0 4.0 215.0 2.1 na 238.0 na na na na na na na na na na na na 1.0 20.0 59

174.0 207.0 62.3 17.6 36.3 254.0 4.2 1.4 363.0 14.2 30.5 16.9 3.7 1.1 3.5 3.1 1.8 0.3 1.7 0.3 0.4 2.6 17.7 93.8

145.0 216.0 73.1 19.4 28.6 824.0 13.8 1.0 447.0 25.8 65.2 36.3 7.3 2.2 6.0 4.6 2.2 0.3 2.0 0.3 0.9 2.0 23.3 222

89.0 262.0 55.6 18.6 81.0 515.0 4.3 1.3 943.0 20.3 42.8 23.1 4.6 1.6 3.7 3.1 1.6 0.3 1.7 0.3 0.4 2.9 17.2 141

133.0 423.0 55.3 15.3 20.5 564.0 2.2 0.7 809.0 9.3 20.7 13.5 3.1 0.9 2.7 2.3 1.4 0.2 1.4 0.2 0.2 0.8 13.4 62.6

135.0 650.0 130.0 14.3 45.4 246.0 3.1 1.7 267.0 12.7 26.2 14.1 3.0 0.7 2.4 2.0 1.1 0.2 1.2 0.2 0.3 2.5 12.0 88.2

154.0 311.0 41.5 17.6 30.0 356.0 3.1 0.8 602.0 12.6 26.6 15.8 3.3 0.9 2.8 2.2 1.3 0.2 1.3 0.2 0.3 2.3 12.7 76.6

169.0 429.0 55.5 14.9 31.2 442.0 2.2 0.7 657.0 9.7 20.8 12.6 2.9 0.9 2.6 2.3 1.3 0.2 1.3 0.2 0.2 1.5 13.1 52.9

220.0 74.8 52.1 18.8 88.3 311.0 4.9 2.1 932.0 18.8 39.1 20.5 4.4 1.1 4.3 4.0 2.3 0.3 2.4 0.3 0.4 4.1 22.9 132

87

118.0 2168.0 513.0 9.1 18.9 275.0 1.9 1.6 142.0 7.0 16.0 7.6 1.8 0.5 1.3 1.3 0.7 0.1 0.7 0.1 0.1 1.0 7.2 48.3

Merroucha annular structure

4-2/4


wt% SiO2 TiO2 Al2O3 Fe2OT3 MnO MgO CaO Na2O K2O P2O5 H2Oþ Total ppm V Cr Ni Ga Rb Sr Nb Cs Ba La Ce Nd Sm Eu Gd Dy Er Tm Yb Lu Ta Th Y Zr

Yetti ‘malignite’ 4-1/4

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M. KAHOUI ET AL.

of the Hank series; the latter series are cross-cut by doleritic and gabbro– doleritic sills and dykes. The magnetic anomalies detected at the intersection of this zone with the north–south and NNE –SSW faults may be due to mafic –ultramafic alkaline rocks (including lamproite, kimberlite and/or related intrusions) that do not crop out. Alkaline magmatism is preferentially associated with major linear zones, may be repeated over a long period of time in the same region (Black et al. 1985), and is located at the intersections of the major and subsidiary structures (White et al. 1995). In the second area, the principal feature is the presence of small dioritic stocks and plutons, gabbroic and mafic– ultramafic alkaline intrusions and numerous basic dykes (Buffie`re et al. 1965a, b; Buffie`re 1966; Azzouni-Sekkal 1976; AzzouniSekkal et al. 2003). Within this area we located new small mafic circular structures (diameter of 100– 250 m) at the intersection of NNW–SSE and NNE–SSW conjugate faults, ultramafic and

basic dykes. The mafic rocks are dolerite, gabbro–dolerite, gabbro, basalt and lamprophyre (Table 1); however, dolerite and gabbro–dolerite are the most abundant. The rocks are undeformed. In the areas selected, we describe two intrusive groups that are thought to be of mantle origin: (1) a group including annular or circular intrusions (Djebel Drissa complex, Yetti ‘malignite’ intrusion, Merroucha structure); (2) a group of ultramafic and mafic dykes (Figs 3 and 5). New geochemical data are presented for rocks and minerals. The samples were analysed for major and trace elements, respectively, by inductively coupled plasma mass spectrometry (ICP-MS) and ICP atomic emission spectrometry (CP-AES) in the laboratories of CRPG–CNRS (Nancy). The analytical procedures are described at http://crpg.cnrsnancy.fr/SARM. The chemical composition of the minerals was determined using a Cameca Camebax electron probe microanalyser at the University of Paris VI.

Fig. 6. Photomicrographs of the Eglab ‘malignite’: (a) clinopyroxene inclusions in nepheline (Ne) surrounded by prismatic clinopyroxene; overgrowths of pale green CPX2 with CPX1; (b) CPX1– CPX2 inclusions in the poikilitic brown–green amphibole (Mg-hastingsite) with partial replacement by phlogopite (Phl); (c) large laths of green amphibole (Amp) including CPX1– CPX2; (d) CPX1– CPX2 inclusions in nepheline surrounded by amphibole.


The alkaline – peralkaline Djebel Drissa ring complex The Djebel Drissa complex belongs to the Aftout granitoids and is dated at 2081 + 13 Ma (zircon evaporation method; Kahoui et al. 1996). It displays characteristics of alkaline–peralkaline post-collisional granitoids (A-type granitoids), has a sub-circular form (16 12 km) and is located on the NW–SE-striking mega-shear zone of Chenachane (Kahoui 1988; Fig. 2); the latter constitutes an internal NW–SE-trending fault and fracture corridor. The complex is cross-cut by NNE– SSW and north–south syenitic, felsic and mafic dykes (Fig. 3c), which exhibit discontinuous en echelon

89

patterns indicating dextral displacement. This dextral displacement is consistent with a sinistral, prominent, kilometre-scale ENE –WSW-striking corridor that indicates a trans-tensional regime. For the Djebel Drissa ring complex, partial melting of an enriched lithospheric mantle evolving by fractional crystallization is advocated for the generation of the peralkaline granites (Kahoui & Mahdjoub 2004).

The Yetti ‘malignite’ circular complex In the Yetti domain, a melanocratic coarse-grained circular intrusion (2–2.5 km in diameter) is located within a releasing bend of en echelon dykes (Location I in Fig. 5). This intrusion is composed

Table 2. Chemical composition of pyroxene from Yetti ‘malignite’ Mineral: Analysis:

CPX1 94

CPX2 95

CPX2 96

CPX1 98

CPX2 99

CPX1 100

CPX2 101

CPX1 102

CPX1 106

CPX2 109

CPX1 110

wt% SiO2 TiO2 Al2O3 Cr2O3 FeOT FeO Fe2O3 MnO MgO NiO CaO Na2O K2O Total

53.43 0.24 1.81 0.06 3.46 1.60 2.07 0.09 15.74 0.17 24.12 0.71 0.00 100.05

52.46 0.46 1.70 0.00 9.31 4.72 5.10 0.21 11.59 0.00 21.25 2.23 0.00 99.71

52.88 0.43 1.67 0.05 8.80 5.10 4.11 0.05 11.68 0.06 21.27 2.22 0.01 99.54

53.18 0.30 1.85 0.34 3.64 1.95 1.88 0.00 15.36 0.01 23.99 0.82 0.00 99.67

52.61 0.38 1.29 0.10 10.40 7.90 2.78 0.04 11.31 0.04 21.05 1.76 0.00 99.26

52.65 0.19 2.08 0.06 3.98 2.25 1.93 0.01 15.23 0.06 24.43 0.51 0.00 99.40

52.07 0.35 1.62 0.04 8.98 4.15 5.36 0.20 12.47 0.04 22.95 1.41 0.02 100.67

53.06 0.30 2.24 0.08 3.72 2.56 1.29 0.04 15.52 0.05 23.86 0.61 0.00 99.61

53.16 0.24 2.07 0.10 3.57 3.10 0.52 0.05 15.66 0.01 23.63 0.49 0.06 99.07

52.47 0.20 2.34 0.00 5.75 4.51 1.38 0.25 13.62 0.06 22.25 1.15 0.00 98.23

52.87 0.18 2.21 0.12 3.71 1.52 2.44 0.01 15.55 0.09 23.95 0.70 0.04 99.66

Si AlIV AlVI Alt Ti Cr Fe3þ Fe2þ Mg Ni Mn Ca Na K Total

1.95 0.05 0.03 0.08 0.01 0.00 0.06 0.05 0.86 0.01 0.00 0.94 0.05 0.00 4.00

1.96 0.04 0.03 0.07 0.01 0.00 0.14 0.15 0.64 0.00 0.01 0.85 0.16 0.00 4.00

1.97 0.03 0.05 0.07 0.01 0.00 0.12 0.16 0.65 0.00 0.00 0.85 0.16 0.00 4.00

1.95 0.05 0.03 0.08 0.01 0.01 0.05 0.06 0.84 0.00 0.00 0.94 0.06 0.00 4.00

1.98 0.02 0.04 0.06 0.01 0.00 0.08 0.25 0.64 0.00 0.00 0.85 0.13 0.00 4.00

1.94 0.06 0.03 0.09 0.01 0.00 0.05 0.07 0.84 0.00 0.00 0.96 0.04 0.00 4.00

1.93 0.07 0.00 0.07 0.01 0.00 0.15 0.13 0.69 0.00 0.01 0.91 0.10 0.00 4.00

1.95 0.05 0.04 0.10 0.01 0.00 0.04 0.08 0.85 0.00 0.00 0.94 0.04 0.00 4.00

1.96 0.04 0.05 0.09 0.01 0.00 0.01 0.10 0.86 0.00 0.00 0.93 0.03 0.00 4.00

1.96 0.04 0.07 0.10 0.01 0.00 0.04 0.14 0.76 0.00 0.01 0.89 0.08 0.00 4.00

1.94 0.06 0.03 0.10 0.00 0.00 0.07 0.05 0.85 0.00 0.00 0.94 0.05 0.00 4.00

XMg XDi XHd XJd XCaTs XAc

0.95 0.84 0.05 0.03 0.02 0.06

0.81 0.68 0.16 0.03 0.01 0.14

0.80 0.67 0.17 0.05 0.00 0.13

0.93 0.82 0.06 0.03 0.02 0.07

0.72 0.60 0.24 0.04 0.00 0.10

0.92 0.84 0.07 0.03 0.02 0.04

0.84 0.75 0.14 0.00 0.03 0.11

0.92 0.82 0.08 0.04 0.02 0.03

0.90 0.81 0.09 0.04 0.02 0.02

0.84 0.72 0.13 0.07 0.01 0.04

0.95 0.83 0.05 0.03 0.03 0.06

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M. KAHOUI ET AL.

Fig. 7. Pyroxene composition from the Eglab ‘malignite’ and ‘komatiite–picrite’ compared with pyroxene composition of Guinea komatiites (Tegyey & Johan 1989) and kimberlites (Kaminsky et al. 2004; Masun et al. 2004).

of a ‘malignite’ rock and is devoid of any deformation or metamorphism. It intrudes the Yetti granitoids and is associated with a mesocratic syenite, showing cumulate textures (Azzouni-Sekkal et al. 2003). The lithological relationships between the two intrusions are unclear because of the superficial cover. Mineralogically, the ‘malignite’ is composed predominantly of clinopyroxene, which is the cumulus mineral; clinopyroxene occurs as green, zoned, euhedral prisms (CPX1) and subhedral to euhedral laths (CPX2) with different sizes (Fig. 6a and b). The intercumulus minerals are anhedral nepheline, subhedral to euhedral magnesiohastingsite and pargasite, rare euhedral to subhedral biotite, and oxides. Secondary minerals are blue–green amphibole after clinopyroxene, biotite replacing pyroxene, and amphibole occurring as small interstitial, acicular crystals. The ‘malignite’ has low SiO2 (48 –49 wt%), and high MgO (11.4 wt%), CaO (17.5 wt%) and Na2O (3.9 wt%) contents (Table 1). The presence of normative nepheline (Ne ¼ 18%) is in accordance with the undersaturated character of the rock, and

indicative of malignite (Mitchell & Platt 1979; Mitchell 1996), although K2O is lower and MgO higher than typical for rocks of malignite affinity. The microprobe analyses highlight the two groups of clinopyroxene: a high-Mg diopside (CPX 1: 15.75 wt% MgO) and a lower –Mg diopside-salite one (CPX 2: 11– 13 wt% MgO) (Table 2; Fig. 7). The zoned pyroxenes (from Di84Hd5Ac6 to Di60Hd24Ac10 and Jd3 –Jd4) are comparable with those found in other malignite occurrences and are similar to the least evolved pyroxenes of some alkaline rocks (Mitchell & Platt 1979). The large green laths of amphibole (Fig. 6b and 6c) have Mg-hastingsitic compositions and the pale blue secondary acicular crystals pargasitic compositions (Na2O 3.5–4 wt%). Azzouni-Sekkal et al. (2003) indicated that one analysis of amphibole gave a value of 500 ppm Cr. Analyses of the large plates of nepheline enclosing earlier pyroxene (Fig. 6a and 6d) and of the large poikilitic phlogopite biotite replacing clinopyroxene (Figs 6b, c and 8) are presented in Tables 3 and 4.


91

(a) 3

Eastonite

Siderophyllite

2.8

AlIV

2.6

2.4

2.2

Eglab “komatiite” Eglab “malignite” Phlogopite

2 0

Annite 0.2

0.4

0.6

0.8

1

Fe/Fe + Mg

(b) 20

K1

Al2O3 wt%

16

UML

12

K2 8

4 Eglab “komatiite” Eglab “malignite” 0 0

1

2

3

4 TiO2 wt%

5

6

7

8

Fig. 8. Phlogopite composition (a) in AlIV v. Fe/(Fe þ Mg) diagram and (b) in Al2O3 v. TiO2 diagram, with fields for kimberlites (K1, K2) and ultramafic lamprophyre (UML) after Mitchell (1995).

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Table 3. Chemical composition of nepheline from Yetti ‘malignite’ Analysis: wt% SiO2 TiO2 Al2O3 Cr2O3 FeOT MnO MgO NiO CaO Na2O K2O BaO F Cl Total

111

112

113

114

43.73 0.04 33.59 0.02 0.01 0.08 0.01 0.00 1.46 15.50 5.14 0.12 0.00 0.00 99.69

42.94 0.00 33.36 0.02 0.08 0.00 0.02 0.00 1.47 15.08 5.42 0.00 0.05 0.01 98.44

43.64 0.02 33.55 0.00 0.00 0.00 0.01 0.00 1.41 15.56 5.39 0.27 0.14 0.00 99.97

43.67 0.00 33.64 0.06 0.00 0.00 0.03 0.04 1.32 15.47 4.85 0.00 0.00 0.02 99.09

The mesocratic syenite contains small crystals of diopside, alkali feldspar (perthite þ perthitic microcline), albite, brown –green hornblende, rare biotite and interstitial quartz; accessories are apatite and titanite. The alteration minerals are kaolinite, sericite, actinolite and chlorite.

The Merroucha circular structure The Merroucha circular structure is a small body (about 250 m in diameter), representative of the mafic rocks discovered within the Yetti –Eglab Junction. This structure, which is cross-cut by olivine basalt dykes, is located at the intersection of NNW–SSE and NNE–SSW conjugate faults (Location II in Fig. 5; Fig. 9). It is intrusive into the Chegga series and is composed of dark to redcoloured rocks with medium- to coarse-grained cumulate textures.

Table 4. Chemical composition of phlogopite from ‘malignite’ and Eglab ‘komatiite – picrite’ Rock:

‘malignite’

‘komatiite– picrite’

Analysis:

107

108

47

48

49

58

59

60

61

(wt%) SiO2 TiO2 Al2O3 Cr2O3 FeOT MnO MgO NiO CaO Na2O K2O BaO F Cl O2 ¼ F O2 ¼ Cl Total

38.57 1.47 14.18 0.09 11.86 0.14 17.70 0.00 0.05 0.28 9.27 0.56 0.67 0.02 0.28 0.00 94.86

37.64 1.88 14.41 0.05 12.33 0.21 17.30 0.03 0.18 0.24 8.71 0.15 0.29 0.00 0.12 0.00 93.42

38.40 2.95 14.53 0.04 13.14 0.03 16.62 0.04 0.04 0.18 7.52 0.15 0.00 0.03 0.00 0.01 93.67

37.89 2.90 14.31 0.03 12.85 0.05 16.94 0.00 0.05 0.13 7.41 0.26 0.03 0.05 0.01 0.01 92.89

39.51 2.86 14.39 0.00 13.70 0.18 14.55 0.02 0.10 0.23 7.86 0.11 0.00 0.05 0.00 0.01 93.56

36.71 1.94 14.41 0.04 12.32 0.04 19.76 0.09 0.00 0.04 6.35 0.00 0.05 0.06 0.02 0.01 91.81

37.57 0.37 14.11 0.07 12.14 0.09 20.98 0.11 0.04 0.05 4.70 0.04 0.05 0.06 0.02 0.01 90.37

38.18 2.76 14.24 0.05 12.75 0.01 16.53 0.06 0.15 0.26 7.95 0.33 0.00 0.03 0.00 0.01 93.32

37.80 3.11 14.46 0.00 14.20 0.11 15.09 0.15 0.10 0.28 8.58 0.11 0.00 0.02 0.00 0.00 93.99

Si AlIV AlVI Ti Cr Ni Mg Fe2þ Mn Ca Na K Ba F Cl

5.77 2.23 0.26 0.17 0.01 0.00 3.94 1.48 0.02 0.01 0.08 1.77 0.03 0.32 0.01

5.67 2.33 0.24 0.21 0.01 0.00 3.89 1.55 0.03 0.03 0.07 1.68 0.01 0.14 0.00

5.72 2.28 0.27 0.33 0.00 0.01 3.69 1.64 0.00 0.01 0.05 1.43 0.01 0.00 0.01

5.69 2.31 0.23 0.33 0.00 0.00 3.79 1.62 0.01 0.01 0.04 1.42 0.02 0.01 0.01

5.90 2.10 0.43 0.32 0.00 0.00 3.24 1.71 0.02 0.02 0.07 1.50 0.01 0.00 0.01

5.54 2.46 0.10 0.22 0.01 0.01 4.44 1.55 0.00 0.00 0.01 1.22 0.00 0.02 0.02

5.68 2.32 0.20 0.04 0.01 0.01 4.73 1.54 0.01 0.01 0.02 0.91 0.00 0.02 0.02

5.73 2.27 0.25 0.31 0.01 0.01 3.70 1.60 0.00 0.02 0.08 1.52 0.02 0.00 0.01

5.69 2.31 0.25 0.35 0.00 0.02 3.38 1.79 0.01 0.02 0.08 1.65 0.01 0.00 0.01

XFe

0.27

0.29

0.31

0.30

0.35

0.26

0.25

0.30

0.35


93

Fig. 9. Mafic dykes (dolerite), Merroucha (II in Fig. 5) and other circular structures on aerial photo (scale: 1:50 000).

The rocks are quartz-diorites and mainly composed of large laths of subhedral to euhedral zoned andesine. Augite forms clusters and is altered to brown or green hornblende at the rims; brown hornblende occurs as small crystals whereas quartz and orthoclase are interstitial and have micrographic textures. Biotite is rare. Accessories include apatite, zircon, allanite, magnetite and ilmenite, sometimes surrounded by coronas of titanite. Hydrothermal alteration is observed with the development of sericite, epidote, actinolite and chlorite assemblages from plagioclase, amphibole and biotite. The rocks have a calc-alkaline composition with some tholeiitic affinities. They are magnesian, enriched in large ion lithophile elements (LILE) þ Th, compared with light REE (LREE) and HFSE high field strength elements (HFSE), and display negative anomalies in Nb–Ta, and Ti (Peucat et al. 2005; Table 1). The high LILE and negative Nb–Ta and Ti anomalies are characteristics of the Guinea tholeiites (low-Ti), which were emplaced around 200 Ma ago and are related

to the break-up of the Pangaea supercontinent (Deckart et al. 2005). To the north of the Merroucha structure, Azzouni-Sekkal (1976) described other mafic massifs with calc-alkaline composition and tholeiitic affinities.

The ultramafic dyke with ‘komatiitic – picritic’ affinities Until now, only one mafic dyke of ‘komatiitic – picritic’ affinities has been reported. This was discovered in 1992 during our fieldwork in the Chegga area (Location III in Fig. 5); it cross-cuts metagabbro and Chegga Archaean relicts, which are themselves intruded by the 2.1 Ga Chegga granite. The dyke crops out for only a few metres and is present also as metre-sized boulders that are irregularly scattered. The rock is much altered and exhibits radiating carbonate –serpentine aggregates (Fig. 10a). It is composed of abundant euhedral and globular

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Fig. 10. Photomicrographs of the Eglab ‘komatiite–picrite’: (a) texture of ‘komatiite– picrite’ underlined by radiating carbonate– serpentine aggregates; (b) euhedral olivine phenocrysts (Ol) completely replaced by serpentine and carbonate; (c) minor corroded globular olivine including spinel; (d) spinel inclusions preserved in clinopyroxene (Cpx) and phlogopite (Phl); clinopyroxene is partially replaced by amphibole (Amp); (e) spinel atolls showing preserved uniform core (relicts) and discontinuous rim corona; (f) euhedral, hexahedral poikilitic spinel (Sp2) with limited resorption of its crystal margins.

olivine phenocrysts completely replaced by serpentine and carbonate (Fig. 10b and c), acicular pyroxene, chrome spinel, phlogopite and opaque minerals. The fine-grained matrix of the rock contains essentially alteration minerals, such as

carbonate, serpentine and amphibole with euhedral, isolated or grouped, oxides. Clinopyroxene (Table 5, Fig. 10d) occurs as prisms and is partially replaced by secondary amphiboles (Tschermakite and Mg-hastingsite). In


95

Table 5. Chemical composition of pyroxene from Eglab ‘komatiite – picrite’ Analysis:

17

19

27

32

40

41

42

45

72

84

85

wt% SiO2 TiO2 Al2O3 Cr2O3 FeOT FeO Fe2O3 MnO MgO NiO CaO Na2O K2O Total

51.05 0.57 4.26 0.90 11.43 11.43 0.00 0.33 16.24 0.09 11.05 0.88 0.31 97.11

51.17 0.42 3.15 0.05 14.43 14.43 0.00 0.44 14.97 0.00 10.05 0.79 0.20 95.67

51.00 0.54 4.45 0.16 11.50 11.50 0.00 0.20 16.66 0.14 10.73 0.88 0.29 96.55

51.98 0.49 3.68 0.27 11.20 11.20 0.00 0.20 16.67 0.00 11.11 0.90 0.22 96.73

58.66 0.45 4.60 0.22 10.18 10.18 0.00 0.30 14.21 0.14 6.65 0.59 0.12 96.10

52.38 0.59 4.06 0.04 11.20 11.20 0.00 0.22 15.87 0.00 10.20 0.77 0.28 95.62

51.75 0.51 4.02 0.14 11.28 11.28 0.00 0.29 16.36 0.06 10.81 0.72 0.31 96.25

52.04 0.62 3.58 0.21 11.52 11.52 0.00 0.17 16.37 0.14 11.58 0.80 0.18 97.19

51.18 0.76 3.82 0.00 12.20 12.20 0.00 0.32 15.55 0.09 11.14 0.87 0.35 96.27

53.20 0.28 2.70 0.21 10.00 10.00 0.00 0.24 17.65 0.06 11.16 0.73 0.18 96.40

53.35 0.33 2.91 0.43 10.84 10.84 0.00 0.05 17.21 0.00 10.84 0.66 0.17 96.79

Si AlIV AlVI AlT Ti Cr Fe3þ Fe2þ Mg Ni Mn Ca Na K

1.59 0.41 0.10 0.50 0.07 0.00 0.36 0.03 0.78 0.00 0.01 0.47 0.14 0.05

1.94 0.06 0.13 0.19 0.02 0.03 0.00 0.36 0.92 0.00 0.01 0.45 0.06 0.01

2.00 0.00 0.14 0.14 0.01 0.00 0.00 0.47 0.87 0.00 0.01 0.42 0.06 0.01

1.94 0.06 0.14 0.20 0.02 0.00 0.00 0.37 0.95 0.00 0.01 0.44 0.07 0.01

1.98 0.02 0.14 0.17 0.01 0.01 0.00 0.36 0.94 0.00 0.01 0.45 0.07 0.01

2.28 0.00 0.21 0.21 0.01 0.01 0.00 0.33 0.82 0.00 0.01 0.28 0.04 0.01

2.02 0.00 0.18 0.18 0.02 0.00 0.00 0.36 0.91 0.00 0.01 0.42 0.06 0.01

1.98 0.02 0.16 0.18 0.01 0.00 0.00 0.36 0.93 0.00 0.01 0.44 0.05 0.02

1.98 0.02 0.14 0.16 0.02 0.01 0.00 0.37 0.93 0.00 0.01 0.47 0.06 0.01

1.97 0.03 0.14 0.17 0.02 0.00 0.00 0.39 0.89 0.00 0.01 0.46 0.06 0.02

2.02 0.00 0.12 0.12 0.01 0.01 0.00 0.32 1.00 0.00 0.01 0.45 0.05 0.01

XMg XDi XHd XJd XOpx XCaTs XAc

0.96 0.08 0.00 0.09 0.36 0.13 0.34

0.72 0.28 0.11 0.08 0.46 0.04 0.03

0.65 0.26 0.14 0.07 0.49 0.03 0.01

0.72 0.27 0.11 0.08 0.48 0.05 0.02

0.73 0.30 0.11 0.08 0.46 0.03 0.02

0.71 0.17 0.07 0.06 0.59 0.08 0.02

0.72 0.28 0.11 0.08 0.48 0.04 0.02

0.72 0.29 0.11 0.07 0.47 0.04 0.02

0.72 0.31 0.12 0.07 0.45 0.03 0.02

0.69 0.29 0.13 0.08 0.45 0.02 0.02

0.76 0.33 0.11 0.06 0.46 0.02 0.01

the Mg –Ca– Fe diagram (Fig. 7), clinopyroxene plots within the field of augite, as is also the case for the Eburnean Guinean komatiites (Tegyey & Johan 1989); we note, however, a difference in the Al and Ti values (2.7– 4.6 wt% Al2O3 and 0.28– 0.76 wt% TiO2), which are lower than for the Guinean pyroxenes (5.71–9.37 wt% Al2O3 and 0.45 –0.92 wt% TiO2). Spinel appears predominantly as (Fig. 10b, c, e and f) (1) small flat-faced euhedral groundmass minerals (0.01 mm, rarely up to 0.30 mm), (2) inclusions in olivine and clinopyroxene, or (3) replacement products formed during the serpentinization of olivine. Homogeneous and discrete continuous zoned spinels crystallized with an euhedral habit; some large crystals are poikilitic and

most of them show various degrees of resorption, limited to a minor corrosion of their rims (Fig. 10e); others show dissolution, which gives rise to atoll-textured grains (Fig. 10f). Atoll spinels are common in many kimberlites (Mitchell 1995). The compositions of spinel (Table 6) show 34.6 –65 wt% Cr2O3, 2 –11 wt% MgO and 4.5– 16 wt% Al2O3. TiO2 contents are low (0.01– 0.72 wt%), and MnO is generally not detected. Compared with Guinean basaltic komatiites and tholeiitic basalts spinels (Tegyey & Johan 1989), Eglab spinels have much higher MgO contents. The former have 0.3–0.5 wt% MgO in the basaltic komatiites and 1.2 wt% MgO in the tholeiitic basalts.

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Table 6. Chemical composition of spinel from Eglab ‘komatiite– picrite’ Mineral: Analysis:

Titano-magnetite 1

Sp2 Poikilitic 3

Sp 4

Sp2 in olivine 5

6

Sp2 grain 7

(wt%) SiO2 TiO2 Al2O3 Cr2O3 FeOT MnO MgO CaO NiO ZnO Na2O K2O Total

2.29 5.86 0.66 0.41 82.63 0.09 0.80 0.04 0.17 0.00 0.00 0.04 92.98

0.08 0.65 14.26 37.68 40.73 0.00 2.16 0.02 0.08 0.61 0.06 0.02 96.35

0.03 0.72 14.30 37.66 41.25 0.00 2.11 0.00 0.10 0.56 0.00 0.00 96.72

0.07 0.31 10.52 44.87 29.99 0.00 9.69 0.03 0.14 0.13 0.00 0.03 95.77

0.12 0.30 10.45 46.21 29.82 0.00 9.81 0.02 0.14 0.17 0.01 0.02 97.05

0.17 0.41 13.09 41.18 32.43 0.00 9.68 0.01 0.17 0.07 0.01 0.01 97.23

Si AlIV AlVI Ti Cr Fe3þ Fe2þ Mn Mg Mn Ca Na K Ni Zn

0.70 0.24 0.00 1.35 0.10 11.58 9.58 0.00 0.37 0.02 0.01 0.00 0.01 0.04 0.00

0.02 1.98 2.76 0.14 8.40 2.58 7.02 0.00 0.91 0.00 0.01 0.03 0.01 0.02 0.13

0.01 1.99 2.75 0.15 8.37 2.57 7.14 0.00 0.88 0.00 0.00 0.00 0.00 0.02 0.12

0.02 1.98 1.40 0.06 9.68 2.79 4.06 0.00 3.94 0.00 0.01 0.00 0.01 0.03 0.03

0.03 1.97 1.35 0.06 9.85 2.66 4.07 0.00 3.94 0.00 0.01 0.00 0.01 0.03 0.03

0.04 1.96 2.15 0.08 8.66 2.99 4.22 0.00 3.84 0.00 0.00 0.01 0.00 0.04 0.01

%Magnetite %Hercynite %Chromite XFe

97.18 2.00 0.82 0.56

16.43 30.14 53.44 0.33

16.36 30.23 53.41 0.33

17.58 21.34 61.08 0.62

16.79 20.97 62.23 0.62

19.00 26.04 54.96 0.62

Primary groundmass spinel crystals (Sp1) show higher contents of Cr2O3 (54–65 wt%) and MgO (9–11 wt%), and lower FeO (20– 26 wt%), TiO2 (,0.2 wt%) and Al2O3 (4.5–9.5 wt%) than the poikilitic and atoll spinel crystals (Sp2); the latter have 38–50 wt% Cr2O3, 2–5 wt% MgO, 30– 40 wt% FeOT, ,0.8 wt% TiO2 and 10–14 wt% Al2O3. Rims of grains are usually enriched in Fe. The plots of spinel compositions in TiO2 v. Cr2O3 (Fig. 11a) and Al2O3 v. Cr2O3 (Fig. 11b) diagrams show general trends with increasing TiO2 and Al2O3 while Cr2O3 decreases. These features are well expressed by the contents of chromite, hercynite and magnetite (Table 6). The spinel fields overlap those of Guinea (Tegyey & Johan 1989), Newton Township (Cattell & Arndt 1987) and Munro Township (Ontario, Canada; Arndt et al. 1977). They closely match the peridotitic

trend as defined by Kharkiv et al. (1989), but are distinctive from those of kimberlites (Mitchell 1995; Fig. 11b). Phlogopite, present as laths enclosing smaller opaque minerals, is aluminous (13.1 –14.5 wt% Al2O3) with varying TiO2 (0.37–3.11 wt%), MgO (14.54–20.97 wt%) and FeOT (13.1–14.5 wt%) contents (Table 4, Fig. 8a and 8b); lower Al2O3 and TiO2 and higher FeOT may be due to alteration. Al2O3 v. TiO2 plots (Fig. 8b) show that most of the phlogopite falls within the compositional field for Group 1 kimberlite (K1) and slightly overlapping the fields for lamproite and ultramafic lamprophyre (UML) (Mitchell 1995, 1996, 1997). Chemically, the Eglab ‘komatiitic –picritic’ rock is characterized by 44.4 wt% SiO2, 16 wt% MgO, 0.75 wt% Na2O þ K2O, 0.3 wt% TiO2, 2168 ppm Cr and 513 ppm Ni (Table 1). In Jensen’s diagram


rim 8

97

rim 9

rim 10

rim 11

rim 12

Sp1 grain 13

Sp1 grain 14

Sp2 atoll 51

0.08 0.38 10.29 48.94 31.37 0.00 6.40 0.00 0.13 0.12 0.00 0.03 97.74

0.03 0.30 7.76 55.52 24.61 0.00 10.11 0.00 0.17 0.06 0.01 0.00 98.56

0.03 0.53 12.19 44.98 34.46 0.00 6.14 0.02 0.04 0.02 0.00 0.01 98.41

0.04 0.35 9.48 49.99 35.39 0.00 4.38 0.03 0.13 0.15 0.05 0.01 99.99

0.03 0.65 11.17 44.71 36.13 0.00 4.99 0.06 0.12 0.22 0.00 0.00 98.07

0.12 0.10 4.80 64.98 20.86 0.00 9.58 0.05 0.00 0.15 0.01 0.00 100.64

0.02 0.01 4.57 64.97 20.17 0.00 9.52 0.09 0.20 0.22 0.05 0.00 99.80

0.07 0.12 13.38 41.76 38.44 0.00 4.04 0.12 0.00 1.79 0.05 0.03 99.78

03.02 1.98 1.36 0.08 10.65 1.82 5.40 0.00 2.62 0.00 0.00 0.00 0.01 0.03 0.03

0.01 1.99 0.47 0.06 11.82 1.58 3.96 0.00 4.06 0.00 0.00 0.00 0.00 0.04 0.01

0.01 1.99 1.91 0.11 9.65 2.22 5.61 0.00 2.48 0.00 0.01 0.00 0.00 0.01 0.00

0.01 1.99 1.08 0.07 10.84 1.95 6.17 0.00 1.79 0.00 0.01 0.03 0.00 0.03 0.03

0.01 1.99 1.64 0.13 9.75 2.33 6.00 0.00 2.05 0.00 0.02 0.00 0.00 0.03 0.04

0.03 1.52 0.00 0.02 13.83 0.54 4.16 0.00 3.84 0.00 0.01 0.00 0.00 0.00 0.03

0.00 1.46 0.00 0.00 13.96 0.59 3.99 0.00 3.85 0.00 0.03 0.02 0.00 0.04 0.04

0.02 1.98 2.29 0.02 8.93 2.75 5.95 0.00 1.63 0.00 0.03 0.03 0.01 0.00 0.36

11.51 21.11 67.38 0.45

9.98 15.51 74.51 0.59

14.05 24.73 61.22 0.46

12.32 19.32 68.36 0.38

14.85 23.10 62.05 0.42

3.37 9.59 87.04 0.51

3.71 9.13 87.16 0.53

17.24 26.74 56.02 0.42

(Continued)

(Fig. 12), the representative point of the Eglab ‘komatiite –picrite’ occupies a position close to French Guiana volcanoclastic komatiites (Capdevila et al. 1999) and the Guinea komatiites and komatiitic basalts (Tegyey & Johan 1989). Arndt & Nesbitt (1982) defined komatiites as lavas or volcanoclastic rocks with more than 18 wt% MgO (anhydrous basis) but the problem of the large variation in the composition of magnesian volcanic rocks and their classification as picritic or komatiitic is still debateable (Le Bas 2000; Hanski et al. 2001; Kerr & Arndt 2001); so the Eglab rock could be classified as either komatiitic or picritic according to the various researchers. With respect to the value of the CaO/Al2O3 ratio (0.86), which is close to unity, the high Al2O3/TiO2 ratio (28) and the flat heavy REE (HREE) pattern, the Eglab ‘komatiitic-picritic’ rock could be classified as an Al-undepleted

komatiite (Nesbitt et al. 1979) or Munro-type komatiite (Arndt 1994). REE patterns (Fig. 13a) show that the Eglab ‘komatiitic –picritic’ rock has a parallel profile to those of Guiana komatiites; however, the latter are more enriched in these elements. In the spidergram (Fig. 13b), the range for the Eglab ‘komatiite – picrite’ is also comparable with the phlogopitebearing metakomatiites of French Guiana with the negative anomalies for Th, Nb and Ti; nevertheless, it differs from typical komatiite. The concentration of immobile elements is very low and distinct from those of kimberlite.

The lamprophyric, doleritic and gabbro dykes In the Eglab shield, the most mafic dykes are oriented north– south (exceptionally NW –SE and

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Table 6. (Continued) Mineral: Analysis:

52

53

54

55

56

57

(wt%) SiO2 TiO2 Al2O3 Cr2O3 FeOT MnO MgO CaO NiO ZnO Na2O K2O Total

0.11 0.03 15.22 44.13 29.22 0.00 9.13 0.03 0.00 1.78 0.03 0.01 99.67

0.07 0.16 14.00 42.67 35.78 0.00 5.16 0.16 0.11 1.58 0.00 0.01 99.69

0.09 0.11 14.46 44.42 35.95 0.00 4.92 0.11 0.08 1.45 0.01 0.00 101.61

0.14 0.12 13.85 43.08 34.64 0.00 5.85 0.12 0.27 1.98 0.00 0.01 100.06

0.12 0.15 13.78 43.13 35.24 0.00 5.71 0.15 0.18 2.24 0.02 0.01 100.73

0.13 0.43 13.91 43.76 34.90 0.00 5.38 0.15 0.01 0.09 0.07 0.02 98.85

Si AlIV AlVI Ti Cr Fe3þ Fe2þ Mn Mg Mn Ca Na K Ni Zn

0.03 1.97 2.68 0.00 9.06 2.24 4.11 0.00 3.53 0.00 0.01 0.01 0.00 0.00 0.35

0.02 1.98 2.44 0.03 9.05 2.43 5.60 0.00 2.06 0.00 0.05 0.00 0.00 0.02 0.32

0.02 1.98 2.52 0.02 9.26 2.16 5.76 0.00 1.93 0.00 0.03 0.01 0.00 0.02 0.29

0.04 1.96 2.38 0.02 9.07 2.46 5.25 0.00 2.32 0.00 0.03 0.00 0.00 0.06 0.40

0.03 1.97 2.33 0.03 9.03 2.56 5.25 0.00 2.25 0.00 0.04 0.01 0.00 0.04 0.44

0.04 1.96 2.45 0.09 9.33 2.05 5.81 0.00 2.16 0.00 0.04 0.03 0.01 0.00 0.02

14.01 29.19 56.79 0.58

15.27 27.83 56.89 0.45

13.58 28.23 58.19 0.42

15.52 27.37 57.11 0.48

16.11 27.06 56.84 0.48

13.00 27.98 59.03 0.42

%Magnetite %Hercynite %Chromite XFe

NE–SW), up to several kilometres in length and from 2–5 cm to 2–3 m in width; their colour is variable from dark to dark green or dark red. The rocks correspond to lamprophyre, basalt, dolerite and gabbro–dolerite. The textures and compositions of the rocks are very variable. They could be microlitic, coarse- to fine-grained, sub-ophitic to ophitic, intersertal, intergranular and sometimes porphyritic. Mineralogically, the primary minerals are plagioclase (andesine, labradorite), diopside–augite, hornblende, biotite, some relicts of olivine, and oxides. Secondary minerals include uralite, serpentine, chlorite, actinolite, sericite, epidote, apatite, calcite and clay minerals; locally intergrowths of quartz and K-feldspar are observed, showing a micrographic texture. Chemically, the dykes have magnesian and medium-K compositions (Table 1). The majority plot on the line that separates high-Mg tholeiite

from calc-alkaline fields (Fig. 12) with two samples plotting above this line and one within the komatiitic basalt field. In the REE diagram (Fig. 13a) the samples show parallel patterns with weak or no Eu anomalies. Sample 35, the most enriched in REE, is characterized mineralogically by numerous isolated or clustered opaque mineral grains (titanomagnetite?). The spidergram (Fig. 13b) shows negative anomalies for Th, Nb and Ti, and a positive K anomaly.

Kimberlite indicator minerals within the Eglab shield The Eglab shield is characterized by large areas of monotonous terrain with superficial soil and sand cover and little or no active drainage. For


Sp1 64

99

Sp1 65

Sp1 67

Sp1 76

Sp1 77

Sp2 86

Sp2 87

Sp2 88

0.08 0.19 8.77 56.71 24.39 0.00 8.47 0.14 0.06 0.05 0.00 0.00 98.85

0.25 0.16 6.68 55.28 32.94 0.00 3.32 0.03 0.07 0.24 0.00 0.00 98.96

0.09 0.18 7.72 60.23 21.40 0.00 10.67 0.10 0.25 0.06 0.10 0.00 100.80

0.06 0.18 8.07 55.01 24.46 0.00 10.65 0.09 0.17 0.10 0.04 0.00 98.82

0.05 0.34 9.56 53.80 26.27 0.00 9.91 0.21 0.16 0.05 0.05 0.00 100.39

0.00 0.38 12.75 44.06 31.43 0.00 8.97 0.04 0.10 0.33 0.05 0.00 98.10

0.07 0.79 16.04 34.64 40.37 0.00 5.39 0.00 0.13 0.02 0.04 0.00 97.50

0.06 0.63 11.13 44.32 36.19 0.00 4.41 0.06 0.25 0.08 0.00 0.03 97.17

0.02 1.98 0.82 0.04 12.14 0.94 4.58 0.00 3.42 0.00 0.04 0.00 0.00 0.01 0.01

0.07 1.93 0.30 0.03 12.38 1.18 6.62 0.00 1.40 0.00 0.01 0.00 0.00 0.02 0.05

0.02 1.98 0.42 0.04 12.52 1.02 3.68 0.00 4.18 0.00 0.03 0.05 0.00 0.05 0.01

0.02 1.98 0.56 0.04 11.61 1.77 3.69 0.00 4.24 0.00 0.02 0.02 0.00 0.04 0.02

0.01 1.99 0.97 0.07 11.18 1.73 4.05 0.00 3.88 0.00 0.06 0.02 0.00 0.03 0.01

0.00 2.00 2.00 0.07 9.26 2.61 4.37 0.00 3.56 0.00 0.01 0.02 0.00 0.02 0.07

0.02 1.98 3.12 0.16 7.39 3.17 5.94 0.00 2.17 0.00 0.00 0.02 0.00 0.03 0.00

0.02 1.98 1.68 0.13 9.79 2.25 6.20 0.00 1.84 0.00 0.02 0.00 0.01 0.06 0.02

5.94 17.63 76.43 0.49

7.48 14.12 78.40 0.28

6.43 15.02 78.56 0.59

11.10 15.95 72.95 0.62

10.88 18.66 70.46 0.58

16.47 25.18 58.36 0.59

20.24 32.57 47.19 0.47

14.35 23.33 62.32 0.40

prospecting of KIM, the investigations concerned some areas selected in the Yetti–Eglab shear zone and in the nearby Tindouf basin, near Aouinet Legraa (Labdi & Zeńia 2001). The results obtained are considered especially encouraging, in that 40 grains of KIM (pyrope, picroilmenite and chrome-diopside) were recovered, and that these appear to form haloes (Fig. 5). On the basement, the haloes of these newly discovered diamond indicator minerals are located: (1) in or near the ‘Anna’ circular structure (Site 1 in Fig. 5), and (2) in the Areigat Lemha (Akilet Deleil area), south of Aouinet Legraa (Site 2 in Fig. 5). In the Tindouf basin, they occur within the bed of the Talha wadi (Site 3 in Fig. 5). Although no analyses were presented for these KIM, Labdi & Zeńia (2001) indicated that two pyrope garnet grains and one chrome-diopside grain were analysed by De Beers Group, and that they revealed a kimberlitic origin.

For the present study, analyses of nine of the garnet grains from the ‘Anna’ area (Table 7) were transmitted to us by the De Beers Group (South Africa). The grains were analysed in their Analytical Services Department, at Johannesburg. Electron microprobe analysis (EMPA) was performed with a wavelength-dispersive spectrometer-equipped Cameca SX-50 operated at an acceleration potential of 20 kV and at a probe current of 30 nA. MgO (Mg), Cr2O3 (Cr), Fe3O4 (Fe), TiO2 (Ti), Al2O3 (Al), ZnS (Zn), wollastonite (Ca,Si), rhodonite (Mn) and jadeite (Na) were used as standards. Counting times were 10 s for all elements. Apparent concentrations were corrected for matrix effects with the on-line PAP program. Detection limits for all elements are of the order of 0.03–0.06 wt% (2s). Of the nine garnet grains, three are most probably from crustal or metamorphic rocks: one grain (B1 054) corresponds in composition to almandine,

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(a)

45

Sp2 (3-4) poikilitic Sp Sp2 in olivine 5-6 Sp2 (7) grain Sp1 (9) grain Sp1 (13-14) Sp2 (8-10-11-12) Sp2 (51 to 55) atoll Sp2 (56-57) in phlogopite Sp1 (64-65-67) Sp1 (76-77) in olivine Sp2 (86-87-88) in olivine Arndt et al.1977

40

Cattell & Arndt 1987 Tegyey & Johan 1989

70 65

Cr2O3 wt%

60 55 50

35 30 0

1

2

3

TiO2 wt%

(b)

60 Sp2 (3-4) poikilitic Sp Sp2 in olivine 5-6 Sp2 (7) grain Sp1 (9) grain Sp1 (13-14) Sp2 (8-10-11-12) Sp2 (51 to 55) atoll Sp2 (56-57) in phlogopite Sp1 (64-65-67) Sp1 (76-77) in olivine Sp2 (86-87-88) in olivine Arndt et al. 1977

Al2O3 wt%

40

Peridotitic trend

Cattell & Arndt 1987 Kimberlite, Mitchell 1995

20

Picritic trend 0 0

10

20

30 40 Cr2O3 wt%

50

60

70

Fig. 11. Composition of spinel from the Eglab ‘komatiite–picrite’: (a ) Cr2O3 v. TiO2; (b) Al2O3 v. Cr2O3 (trends from Kharkiv et al. 1989), compared with spinel compositions of Guinea komatiites (Tegyey & Johan 1989); 7-8-9, Newton Township komatiites (Cattell & Arndt 1987); 8-9-10, Munro Township (Ontario, Canada, Arndt et al. 1977) and kimberlite (Mitchell 1995).

and two grains (B1 118a, b) are spessartite. The garnet grains Bl 129a –f correspond in composition to low-Cr, low-Ti pyrope of the lherzolite suite (G9 group, according to the classification scheme of Dawson & Stephens (1975)). Their data points form a very compact cluster in both the Cr– Ca and Ti –Cr diagrams (Fig. 14a and b). Pyrope with

this composition does occur in kimberlite pipes, both diamondiferous and non-diamondiferous. However, in diamondiferous pipes, pyrope of this type accounts for only a minor proportion of the pyropes occurring therein, whereas in nondiamondiferous pipes (for example, the Obnazhennaya pipe in Yakutia, the Middle Timan pipes in


101

FeOT + Ti

High-Fe Tholeiite Basalt

Komatiitic basalt

High-Mg Tholeiite Basalt

Komatiite

Al

Mg Eglab “komatiite-picrite” French Guiana komatiite (Capdevila et al. 1999) Guinea komatiite (Tegyey & Johan 1989) lamprophyres (25A, B, C, 14) doleritic and gabbroic dykes (30 B, 35, 38G, 38H)

Fig. 12. AFM diagram for ‘komatiite– picrite’, and doleritic, gabbroic and lamprophyric dykes (Jensen 1976).

the Russian platform, and others) pyrope of this sort strongly predominates. Sometimes, rare grains of this pyrope variety occur in alkali basaltic rocks (Mongolia, Minusinsk Kettle, etc.). On the Cr –Ca diagram (Fig. 14a), pyropes plot close to the horizontal line drawn at 2 wt% Cr2O3, used as an arbitrary division between the eclogitic garnet ‘E’ field (,2 wt% Cr2O3) and peridotitic ‘P’ field (.2 wt% Cr2O3) (Gurney 1984). Table 7 shows that some Eglab pyropes (samples 6, 7, 8 and 10) have trace amounts of Na2O (0.07–0.17 wt%). The presence of Na2O in garnet (Na2O 0.07 wt%) is a distinctive feature of eclogite pyropes associated with diamonds, in eclogitic sources of diamond (Gurney et al. 1993).

Discussion The Eglab shield belongs to the WAC Palaeoproterozoic domain, which is known for its diamondbearing field related to kimberlite and lamproite dykes (Knopf 1970; Bardet 1974; Rombouts 2003; Pouclet et al. 2004). The discovery of KIM (G9 type) in this area indicates proximity to the sources, which most probably are kimberlite, as for the diamonds and KIM of the Cretaceous – Quaternary deposits of the Reggane area (Kaminsky et al. 1992a) and those within the Mauritanian Reguibat shield (Rombouts 2003). The geological position of the Reggane area, close to the boundary of WAC, and the long transportation history of diamond and KIM therein

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Sample/Chondrite

(a) 1000

GH28 30B, 35, 38G, 38H 25A, B, C, 14 K6-41 Fr Guiana K3-29* Fr Guiana Komatiite Ontario

10

0 La

Ce

Nd

Sm

Eu

Gd

Dy

Er

Sample/ Primitive Mantle

(b) 1000

Yb

Lu

GH28 30B, 35, 38G, 38H 25A, B, C, 14 K6-41 Fr Guiana K3-29* Fr Guiana Kimberlite

10

0.1 Cs

Rb

Ba

Th

K

Nb

La

Ce

Sr

Zr

Sm

Ti

Tb

Y

Fig. 13. (a) Chondrite-normalized REE patterns for Eglab ‘komatiite– picrite’, and doleritic, gabbroic and lamprophyric dykes compared with (K6-41) metakomatiite, K3-29* metasomatized komatiite with phlogopite of French Guiana (Capdevila et al. 1999), peridotitic komatiite and basaltic komatiite of the Munro Township (Ontario, Canada, Arndt et al. 1977). (b) Mantle-normalized spidergram for Eglab ‘komatiite– picrite’, and doleritic, gabbroic and lamprophyric dykes. Normalization values after Sun & McDonough (1989). Patterns for komatiites of French Guiana are from Capdevila et al. (1999); (K6-41, metakomatiite; K3-29*, metasomatized komatiite with phlogopite) and average kimberlite (Scott Smith 1996) are shown for comparison.

(Kaminsky et al. 1992a) provide a possible explanation for the enigmatic sources of this diamondbearing placer. In this region, the lherzolitic type garnet (G9) is the dominant pyrope population,

but harzburgitic (G10) and eclogitic (E) groups are also present (Kahoui et al. 1998). In the western Reguibat shield (Mauritania), two kimberlite provinces were discovered in the

Table 7. Chemical composition of garnet grains from Eglab shield B1054 2

B1 118 3

B1 118b 4

B1 129a 5

B1 129b 6

B1 129c 7

B1 129d 8

B1 129e 9

B1 129f 10

wt% SiO2 TiO2 Al2O3 Cr2O3 FeOT MnO MgO CaO Na2O K2O Total

36.27 0.04 21.90 0.02 31.25 2.08 6.04 1.65 0.07 0.00 99.31

36.11 0.10 21.73 0.10 12.79 25.18 3.55 0.79 0.13 0.00 100.49

37.41 0.15 21.72 0.15 12.78 24.07 3.25 0.78 0.07 0.00 100.36

42.41 0.06 23.42 1.86 8.17 0.45 19.82 4.76 0.00 0.00 100.96

41.45 0.00 23.02 2.08 8.18 0.35 20.62 4.83 0.12 0.00 100.65

41.64 0.07 23.44 1.87 7.89 0.33 21.03 4.91 0.10 0.00 101.30

39.10 0.05 23.75 1.86 7.82 0.38 20.33 4.88 0.17 0.00 98.26

41.65 0.02 22.57 1.89 7.73 0.26 20.65 4.91 0.00 0.00 99.69

41.14 0.06 23.63 1.90 7.85 0.37 20.62 4.99 0.07 0.00 100.62

Si Ti AlIV AlVI Cr Fe3þ Fe2þ Mg Mn Ca Total

2.90 0.00 0.10 1.96 0.00 0.15 1.94 0.72 0.14 0.14 7.99

2.90 0.01 0.10 1.96 0.01 0.14 0.72 0.43 1.71 0.07 7.98

3.01 0.01 0.00 2.06 0.01 0.00 0.86 0.39 1.64 0.07 7.99

3.01 0.00 0.00 1.96 0.10 0.00 0.49 2.10 0.03 0.36 8.00

2.94 0.00 0.06 1.87 0.12 0.09 0.39 2.18 0.02 0.37 7.98

2.93 0.00 0.07 1.87 0.10 0.10 0.36 2.21 0.02 0.37 7.99

2.83 0.00 0.17 1.85 0.11 0.23 0.24 2.19 0.02 0.38 7.98

2.98 0.00 0.02 1.89 0.11 0.02 0.44 2.20 0.02 0.38 8.00

2.92 0.00 0.08 1.89 0.11 0.09 0.37 2.18 0.02 0.38 7.99

% Almandine % Pyrope % Spessartite % Grossular XFe

65.93 24.48 4.79 4.81 0.73

24.61 14.52 58.54 2.32 0.63

29.08 13.18 55.47 2.27 0.69

16.32 70.58 0.91 12.18 0.19

13.25 73.64 0.71 12.40 0.15

12.29 74.53 0.66 12.51 0.14

8.58 77.27 0.82 13.33 0.10

14.51 72.56 0.52 12.40 0.17

12.65 73.77 0.75 12.83 0.15

Classification Schulze (2003)

crustal

crustal

crustal

lherzolite

lherzolite

lherzolite

lherzolite

lherzolite

lherzolite


Sample: Analysis:

103

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(a) 16 G10

14

G9

12 Cr2O3 wt%

harzburgitic garnet 10 8

lherzolitic garnet

6 4

pyropes (5 to 10) garnets (2 to 4)

2 0 0

2

4

6

8

10

CaO wt% (b) 1.6 pyropes (5 to 10) garnets (2 to 4)

1.4 1.2 pyropes-ilmenites ultrabasites

TiO2 wt%

1 0.8 0.6 0.4

pyropes ultrabasites 0.2 0 0

2

4

6

8

10

12

14

16

Cr2O3 wt% Fig. 14. (a) CaO v. Cr2O3 diagram (Sobolev 1974; Gurney 1984) and (b) TiO2 v. Cr2O3 diagram (Schulze 2003) for Eglab garnets.

Taoudeni basin, along a major NE –SW lineament (Rombouts 2003) and the pyropes associated with diamonds there are harzburgitic (G10). A distinct feature in this region is the location of one province on or near the carbonatitic Richat circular structure, dated at 100 Ma (Fig. 1; Netto et al. 1992; Poupeau et al. 1996). Poupeau et al. (1996) indicated that the Richat structure is situated at the intersection of north–south and NE– SW faults and its position is

controlled by two lithospheric lineaments; on one of these lineaments (ENE –WSW to east –west) doleritic sills and the Djebel Drissa ring complex crop out. The lithospheric structures oriented north–south and ENE –WSW occur in the eastern part of the WAC, where the late phases of the alkaline magmatism of the Tadhak province is dated at 185– 160 Ma (Lie´geois et al. 1991) and in the southern


part where kimberlitic intrusions are emplaced (Haggerty 1992). Diamond-bearing fields in Ivory Coast are also associated with magmatic structures that have a north –south-trend (Pouclet et al. 2004). Reactivation along these lithospheric trends and the association with alkaline magmatism are favourable for the possibility of finding primary diamond sources within the important north– south mafic dyke swarms in the Eglab shield. The discovery of unusual diamondiferous calc-alkaline lamprophyres (Armstrong & Barnett 2003; Lefevbre et al. 2005; De Stephano et al. 2006) reinforces this possibility. Another, alternative primary source of diamonds may not be kimberlitic but actually komatiitic in origin. For example, the Chegga ‘komatiitic– picritic’ rock, discovered in the ‘Yetti –Eglab Junction’, could have the same implications for diamond exploration as the diamondiferous komatiites in French Guiana. In the latter province, the pyrope garnets recognized are predominantly G9 types with subordinate G10 and E types; diamonds are of eclogitic sources (Capdevila et al. 1999). We note that the Guiana Birimian komatiite is totally devoid of diamonds, as are the Birimian ultramafic rocks of the Ivory Coast, Burkina Faso and Niger. Besides the above, the existence of diamondiferous lamproites should not be excluded. In the Man shield, diamonds have two origins: (1) Cretaceous (and Jurassic?) kimberlitic intrusions in Guinea, Sierra Leone and Ivory Coast; (2) Birimian ‘conglomerates’ in Ivory Coast, Ghana, and, perhaps, Burkina Faso. In the ‘conglomerate’ occurrences, garnets and pyroxenes are lacking; diamonds are octahedral and their primary origin is, as yet, unknown (Pouclet et al. 2004).

Conclusion The results of the tectonic analysis and prospecting for diamonds of the Eglab shield demonstrate that this area may be considered as a possible location of primary sources for the North African diamondiferous province, at least within SW Algeria. In relation to its Palaeoproterozoic age (2.2– 2.0 Ga) and its structure, the Eglab shield is defined as a proton and is favourable for the occurrence of diamondiferous primary sources. The tectonic model proposed using aeromagnetic and structural interpretations indicates a control by lithospheric structures on both moderate- to highmagnitude anomalies and the emplacement of alkaline ring complexes (Djebel Drissa), small-sized circular intrusions (Yetti ‘malignite’ and Merroucha intrusions) and widespread mafic, ultramafic and lamprophyric dyke swarms.

105

Three major directions (north –south, NW–SE and WSW–ENE) may be related to reactivation of inherited structures by successive stress fields (post-Eburnean, pre-Pan-African and Early Cretaceous). The north–south direction seems to be the most favourable for emplacement of deep magmatic sources during east–west extension; the associated transtensional NW– SE and WSW– ENE directions are also sites of intrusions. The diamond-bearing fields in the Ivory Coast and the kimberlitic province of the Early Cretaceous Richat carbonatite structure are associated with these structures. The presence of a ‘komatiitic–picritic’ rock and the occurrence of KIM near geological and aeromagnetic structures (e.g. the ‘Anna’ structure) could be considered as especially encouraging. The ‘Yetti –Eglab Junction’ would appear to be the most favourable area for diamonds on the basis of known results to date, and both the Chenachane and Kahal Morrat corridors are other zones where fieldwork could lead to the identification of additional sources. Hence, further exploration, with sampling for diamonds, could lead to the discovery of the primary source(s) of diamonds in northern Africa. The authors are very grateful to D. Dyck, M. Lehtonen and A. Pouclet for their careful reviews and constructive advice, to J. P. Lie´geois for his valuable recommendations, and to I. Coulson for assistance with editing the text. We are very grateful to M. De Wit and the De Beers Group for providing the analyses of pyrope, and also D. Lakrache for her assistance.

References AEROSERVICE CORPORATION 1974. Etude ae´romagne´tique et radiome´trique de l’Alge´rie: Interpre´tation de la re´gion de l’Eglab. Rapport final ine´dit, Vol. IV. SONAREM, Alger. A FFATON , P., G AVIGLIO , P. & P HARISAT , A. 2000. Reáctivation du craton ouest-africain au Panafricain: paleócontraintes de´duites de la fracturation des gre`s neoprote´rozoı¨ques du Karey Gorou (Niger, Afrique de l’Ouest). Comptes Rendus de l’Acade´mie des Sciences, 331, 609–614. A LLEK , K. 2005. Traitement et interpre´tation des donneés ae´romagne´tiques acquises au dessus des blocs de Tindouf et Eglab (Sud Ouest alge´rien): impact sur l’exploration du diamant. The`se de magister, USTHB, Alger. A MA S ALAH , I., L IE´ GEOIS , J. P. & P OUCLET , A. 1996. E´volution d’un arc insulaire oceánique birimien prećoce au Liptako nige´rien (Sirba): geólogie, geóchronologie et geóchimie. Journal of African Earth Sciences, 22, 235– 254. A RMSTRONG , J. & B ARNETT , R. 2003. The association of Zn-chromite with diamondiferous lamprophyres and diamonds: unique compositions as guide of the

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M. KAHOUI ET AL.

diamond potential of non-traditional diamond host rocks. Extended Abstracts, 8th International Kimberlite Conference, FLA 230, Victoria, BC, Canada, 22–27 June. A RNDT , N. T. 1994. Archaean komatiites. In: C ONDIE , K. C. (ed.) Archaean Crustal Evolution. Elsevier, Amsterdam, 11– 44. A RNDT , N. T. & N ESBITT , R. W. 1982. Geochemistry of Munro Township basalts. In: A RNDT , N. T. & N ISBET , E. G. (eds) Komatiites. George Allen & Unwin, London, 309–330. A RNDT , N. T., N ALDRETT , A. J. & P YKE , D. R. 1977. Komatiite and iron-rich tholeiitic lavas of Munro Township, northeast Ontario. Journal of Petrology, 18, 319 –369. A ZZOUNI -S EKKAL , A. 1976. Les stocks plutoniques basiques de la jointure ‘Yetti–Eglab’ (Sahara Occidental Alge´rien). The`se de 3e`me cycle, Universite´ d’Alger, Alge´rie. A ZZOUNI -S EKKAL , A., D EBABHA , F. & I KHLEF , F. 2003. Malignites et syeńites me´socrates associeés, stock plutonique sud Tinguicht, zone de jointure Yetti– Eglab (Dorsale Re´guibat, Alge´rie). Bulletin du Service Geólogique de l’Alge´rie, 14, 79–95. B ARDET , M. G. 1974. Les gisements kimberlitiques de l’Ouest africain. In: Geólogie du diamant. Deuxieme partie: Gisements de diamant d’Afrique. Me´moires du Bureau de Recherche Geólogique et Minie`re, Paris, 83, 178– 212. B ERTRAND , H. 1991. The Mesozoic tholeiitic province of northwest Africa: a volcano-tectonic record of the early opening of Central Atlantic. In: K AMPUZU , A. B. & L UBALA , R. T. (eds) Magmatism in Extensional Structural Setting. The Phanerozoic African Plate. Springer, Berlin, 147–188. B ERTRAND , H., D OSTAL , J. & D UPUY , C. 1982. Geochemistry of early Mesozoic tholeiites from Morocco. Earth and Planetary Science Letters, 58, 225– 239. B ERTRAND -S ARFATI , J., M OUSSINE -P OUCHKINE , A & A IT K ACI , A. 1996. Subdivisions stratigraphiques nouvelles dans la couverture Neóprote´rozoique au nord-est du bassin de Taoudenni (Sahara, Alge´rie). In: BITAM , L. & FABRE , J. (eds) Geodynamique du craton ouest africain central et oriental: he´ritage et e´volution post-panafricains. Me´moires du Service Geólogique de l’Alge´rie, 8, 63– 111. B LACK , R., L AMEYRE , J. & B ONIN , B. 1985. The structural setting of alkaline complexes. Journal of African Earth Sciences, 3, 5 –16. B UFFIE` RE , J.-M. 1966. Sur l’ensemble prećambrien Yetti– Eglab et sur sa couverture infra-tillitique en territoire alge´rien. Comptes Rendus de l’Acade´mie des Sciences, 26, 1513–1516. B UFFIE` RE , J.-M., F AHY , J. C. & P ETEY , J. 1965a. Etude geólogique de la partie orientale de la dorsale Re´guibat. Re´gion des Eglab et secteur nord du Yetti. Rapport ine´dit, SERMI, Paris, ALG 63–09-IV. B UFFIE` RE , J.-M., F AHY , J. C. & P ETEY , J. 1965b. Notice explicative de la carte geólogique au 1/500 0008 de la re´gion de l’Eglab et de la bordure nord du Yetti. SERMI, Paris. C APDEVILA , R., A RNDT , N. T., L ETENDRE , J. & S AUVAGE , J. F. 1999. Diamonds in the volcanoclastic komatiite from French Guiana. Nature, 399, 456–458.

C ATTELL , A. & A RNDT , N. T. 1987. Low- and highalumina komatiites from a late Archaean sequence, Newton Township, Ontario. Contributions to Mineralogy and Petrology, 97, 218–227. C HARDON , D. 1997. Les de´formations continentales archeénnes, exemples naturels et mode´lisation thermomećanique. Me´moires de Geósciences Rennes, 76. D AWSON , J. B. & S TEPHENS , W. E. 1975. Statistical analysis of garnets from kimberlites and associated xenoliths. Journal of Geology, 83, 589–607. D ECKART , K., B ERTRAND , H. & L IE´ GEOIS , J. P. 2005. Geochemistry and Sr, Nd, Pb isotopic composition of the Central Atlantic Magmatic Province (CAMP) in Guyana and Guinea. Lithos, 82, 289 –314. D E S TEPHANO , A., L EFEBVRE , N. & K OPYLOVA , M. 2006. Enigmatic diamonds in Archean calc-alkaline lamprophyres of Wawa, southern Ontario, Canada. Contributions to Mineralogy and Petrology, 151, 158–173. D OUMBIA , S., P OUCLET , A., K OUAMELAN , A. N., P EUCAT , J. J. & V IDAL , M. 1998. Petrogenesis of juvenile-type Birimian (Paleoproterozoic) granitoids in central Coˆte-d’Ivoire, West Africa: geochemistry and geochronology. Precambrian Research, 87, 33–63. D RARENI , A., P EUCAT , J.-J. & F ABRE , J. 1996. Isotopic data (Sr, Nd, Pb) from the West African Craton: the ‘Dorsale Reguibat’, The Eglab Massif (Algeria). Terra Nova, 7, 102. E L O UALI , E. H, G ASQUET , D. & I KENNE , M. 2001. Le magmatisme de la boutonnie`re d’Igherm (Anti-Atlas occidental, Maroc): jalon de distensions neóprote´rozoı¨ques sur la bordure nord du craton ouest africain. Bulletin de la Socie´te´ Geólogique de France, 172, 309–317. E NNIH , N. & L IE´ GEOIS , J. P. 2001. The Moroccan AntiAtlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton. Precambrian Research, 112, 289– 302. EREM 1983. Travaux de Reconnaissance et de Prospection dans les Eglab. Rapport final, ine´dit. EREM, Boumerdes. F ABRE , J. 2005. Geólogie du Sahara occidental et central. Tervuren African Geoscience Collection, 108. G EVIN , P. 1951. Sur la structure du massif cristallin Eglab–Yetti (Sahara Occidental). Comptes Rendus de l’Acade´mie des Sciences, 233, 1129–1130. G EVIN , P. 1958. Notice explicative des cartes geólogiques au 1/500 0008 Tindouf–Eglab. Service de la Carte Geólogique de l’Alge´rie, Alger. G UIRAUD , R., B ELLION , Y., B ENKHELIL , J. & M OREAU , C. 1987. Post-Hercynian tectonics in North and West Africa. Geological Journal, Thematic Issue, 22, 433–466. G URNEY , J. J. 1984. A correlation between garnets and diamond. In: G LOVER , J. E. & H ARRIS , P. G. (eds) Kimberlite Occurrences and Origin: A Basis for Conceptual Models in Exploration. Geology Department and University Extension, University of Western Australia, Publication, 8, 143– 166. G URNEY , J. J., H ELMSTAEDT , H. & M OORE , R. O. 1993. A review of the use and application of mantle mineral

NORTH AFRICAN DIAMONDIFEROUS PROVINCE geochemistry in diamond exploration. Pure and Applied Chemistry, 65, 2423– 2442. H AGGERTY , S. E. 1992. Diamonds in West Africa: tectonic setting and diamond productivity. Russian Geology and Geophysics, 33, 35– 49 [in Russian]. H ANSKI , E., H UHMA , H., R ASTAS , P. & K AMENETSKY , V. V. 2001. The Palaeoproterozoic komatiite– picrite association of Finnish Lapland. Journal of Petrology, 12, 855–876. I ZAROV , V. & B IROUTCHEV , S. 1974. Rapport sur les re´sultats de recherches geólogiques pour le diamant exećuteés au Hoggar de 1969 a` 1973. SONAREM, Alger. J ANSE , A. J. A. 1992. New ideas in subdividing cratonic areas. Geology and Geophysic, 33, 9– 25 [in Russian]. J ENSEN , L. S. 1976. A new cation plot for classifiying volcanic rocks. Ontario Division of Mines, Miscellaneous Papers, 66. K AHOUI , M. 1988. Etude d’un complexe granitique diffe´rencie´ et de sa couverture volcanique: incidence me´tallogeńique (Cas du Djebel Drissa, massif des Eglab, Alge´rie). The`se Doctorat, Universite´ de Nancy I. K AHOUI , M. 1991. Projet de prospection et ve´rification syste´matique des indices dans le massif cristallin des Eglab. Rapport provisoire ine´dit. EREM, Boumerdes. K AHOUI , M. & B ENAMEUR , O. 1992. Recherche et prospection dans les Eglab. EREM, CRD, Boumerdes. K AHOUI , M. & M AHDJOUB , Y. 2001. Crite`res pour la recherche de sources primaires de diamant dans la zone de ‘jointure’ Yetti–Eglab (Dorsale Re´guibat, Craton Ouest Africain). In: 11e`me Se´minaire National des Sciences de la Terre, Universite´ de Tlemcen, Alge´rie, 8. K AHOUI , M. & M AHDJOUB , Y. 2004. An Eburnean alkaline–peralkaline magmatism in the Reguibat Rise: the Djebel Drissa ring complex (Eglab Shield, Algeria). Journal of African Earth Sciences, 39, 115–122. K AHOUI , M., D RARENI , A., F ABRE , J., P EUCAT , J. J. & K ADDOUR , M. 1996. Age e´burneén du complexe annulaire du Dje´bel Drissa (Est de la Dorsale Re´guibat, Alge´rie). In: BITAM , L. & FABRE , J. (eds) Geodynamique du craton ouest africain central et oriental: he´ritage et e´volution post-panafricains. Me´moires du Service Geólogique de l’Alge´rie, 8, 15– 22. K AHOUI , M., B OUZIDI , O. & R AZIBAOUENE , A. 1998. La recherche du diamant dans le Saharien Alge´rien: synthe`se et mise au point. ORGM, Boumerdes. K AHOUI , M., M AHDJOUB , Y. & K AMINSKY , F. V. 2004. Possible kimberlites in the ‘Yetti– Eglab Junction’ (Reguibat Rise, West African Craton, Algeria). In: A SHWAL , L. D. (ed.) Geosciences Africa 2004, Abstract Volume 1, University of Witwatersrand, Johannesburg, 323. K AMINSKY , F. V., K ONYUKHOV , Yu. I., V ERZHAK , V. V., H AMANI , M. & H ENNI , A. 1990. Diamonds of the Algerian Sahara. Mineralogicheskii Zhurrnal, 12, 76– 80 [in Russian]. K AMINSKY , F. V., K OLESNIKOV , S. K., P ETELINA , N. A. ET AL . 1992a. Minerals-indicators of diamond in Algerian Sahara. Mineralogicheskii Zhurrnal, 14, 15– 24 [in Russian]. K AMINSKY , F. V., V ERZHAK , V. V., D AUEV , Y U . M. ET AL . 1992b. The North-African diamondiferous

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province. Russian Geology and Geophysics, 33, 91–95 [in Russian]. K AMINSKY , F. V., F ELDMAN , A. A., V ARLAMOV , V. A., B OYKO , A. N., O LOFINSKY , L. N., S HOFMAN , I. L. & V AGANOV , V. I. 1995. Prognostication of primary diamond deposits. Journal of Geochemical Exploration, 53, 167 –182. K AMINSKY , F.V., S ABLUKOV , S. M., S ABLUKOVA , L. I. & C HANNER , D. M. DeR. 2004. Neoproterozoic ‘anomalous’ kimberlites of Guianiamo, Venezuela: mica kimberlites of ‘isotopic transitional’ type. Lithos, 76, 565– 590. K ERR , A. C. & A RNDT , N. T. 2001. A note on the IUGS reclassification of the high-Mg and picritic volcanic rocks. Journal of Petrology, 42, 2169–2171. K HARKIV , A. D., K VASNITSA , V. N., S AFRONOV , A. F. & Z INCHUK , N. N. 1989. Diamond and diamond indicator minerals of the kimberlites. Naukova, Kiev [in Russian]. K NOPF , D. 1970. Les kimberlites et les roches apparenteés de Coˆte-d’Ivoire. Direction des Mines et de la Geólogie de Coˆte-d’Ivoire, Bulletin, 3. L ABDI , A. & Z E´ NIA , M. S. 2001. Recherche des sources primaires potentielles de diamant dans le massif des Eglab. ORGM, Bećhar. L AMEYRE , J. & L ASSERRE , M. 1967. Etude geóchronologique des syeńites alcalines et ne´phe´liniques du massif annulaire de Hassi-El-Fogra, Mauritanie du Nord. Comptes Rendus de l’Acade´mie des Sciences, 265, 733– 736. L ASSERRE , M., L AMEYRE , J. & B UFFIE` RE , J. M. 1970. Donneés geóchronologiques sur l’axe prećambrien Yetti–Eglab en Alge´rie et en Mauritanie du Nord. Bulletin du Bureau de Recherche Geólogique et Minie`re, IV, 5 –13. L E B AS , M. 2000. IUGS reclassification of the high-Mg and picritic rocks. Journal of Petrology, 41, 1467– 1470. L EFEBVRE , N., K OPYLOVA , M. & K IVI , K. 2005. Archaean calc-alkaline lamprophyres of Wawa, Ontario, Canada: Unconventional diamondiferous volcaniclastic rocks. Precambrian Research, 138, 57–87. L EFORT , J. P., A IFA , T. & B OURROUILH , R. 2004. Evidences paleómagne´tiques et paleóntologiques en faveur d’une position antipodale du craton ouest africain et de la Chine du nord dans le super-continent Rodinia: conse´quences paleógeógraphiques. Comptes Rendus Geóscience, 336, 159–165. L ESQUER , A., B ELTRAO , J. F. & D E A BREU , F. A. M. 1984. Proterozoic links between northeastern Brazil and West Africa; a plate tectonic model based on gravity data. Tectonophysics, 110, 9– 26. L ESQUER , A., V ILLENEUVE , J. C. & B RONNER , G. 1991. Heat flow data from the western margin of the West African Craton (Mauritania). Physics of the Earth and Planetary Interiors, 66, 320 –329. L IE´ GEOIS , J. P., S AUVAGE , J. F. & B LACK , R. 1991. The Permo-Jurassic alkaline province of Tadhak, Mali: Geology, geochronology and tectonic significance. Lithos, 27, 95– 105. M ACNAE , J. C. 1979. Kimberlites and exploration geophysics. Geophysics, 44, 1395– 1416. M AHDJOUB , Y., D RARENI , A. & G ANI , R. 1994. Accre´tion crustale et tectonique verticale a` l’E´burneén dans

108

M. KAHOUI ET AL.

les Massifs des E´glab et du Yetti (Dorsale Re´guibat, Alge´rie). Bulletin du Service Geólogique de l’Alge´rie, 5, 97–107. M AHDJOUB , Y., K AHOUI , M., D RARENI , A. & G ANI , R. 2002. Magmatic evolution during convergence in Palaeoproterozoic Eglab domain, Reghibat Rise (Algeria). 19th Colloquium of African Geology, University of El Djadida, Morocco, 19–22 March, 129. M ASUN , M. K., D OYLE , B. J., B ALL , S. & W ALKER , S. 2004. The geology and mineralogy of the Anuri kimberlite, Nunavut, Canada. Lithos, 76, 75–97. M AURIN , J.-C. & G UIRAUD , R. 1993. Basement control in the development of the early Cretaceous West and Central African rift system. Tectonophysics, 228, 81–95. M ITCHELL , R. H. 1995. Kimberlites, Orangeites and Related Rocks. Plenum, New York. M ITCHELL , R. H. 1996. Classification of undersaturated and related alkaline rocks. In: M ITCHELL , R. H. (ed.) Undersaturated Alkaline Rocks: Mineralogy, Petrogenesis, and Economic Potential. Mineralogical Association of Canada, Short Course Series, 24, 1– 22. M ITCHELL , R. H. 1997. Kimberlite, Orangeite, Lamproite, Melilite, and Minette: A Petrographic Atlas. Almaz Press, Thunder Bay, Ont. M ITCHELL , R. H. & P LATT , R. G. 1979. Nephelinebearing rocks from the Poohbah Lake complex, Ontario: Malignites and Malignites. Contributions to Mineralogy and Petrology, 69, 255– 264. M ORGAN , P. 1995. Diamond exploration from the bottom up: regional geophysical signatures of lithosphere conditions favorable for diamond exploration. Journal of Geochemical Exploration, 53, 145–165. M OUSSINE -P OUCHKINE , A. & B ERTRAND -S ARFATI , J. 1997. Tectonosedimentary subdivisions in the Neoproterozoic to Early Cambrian cover of the Taoudenni basin (Algeria– Mauritania– Mali). Journal of African Earth Sciences, 24, 425 –443. N ESBITT , R. W., S UN , S. S. & P URVIS , A. C. 1979. Komatiites: Geochemistry and genesis. Canadian Mineralogist, 17, 165– 186. N ETTO , A. M., F ABRE , J., P OUPEAU , G. & C HAMPENOIS , M. 1992. Datation par traces de fission de la structure circulaire des Richat (Mauritanie). Comptes Rendus de l’Acade´mie des Sciences, 314, 1179– 1186. P EUCAT , J.-J., C APDEVILA , R., D RARENI , A., M AHDJOUB , M. & K AHOUI , M. 2005. The Eglab massif in the West African Craton (Algeria), an original segment of the Eburnean orogenic belt: petrology, geochemistry and geochronology. Precambrian Research, 136, 309–352. P OTREL , A., P EUCAT , J.-J., F ANNING , C. M., A UVRAY , B., B URG , J. P. & C ARUBA , C. 1996. 3.5 Ga old terranes in the West Africa Craton, Mauritania. Journal of the Geological Society, London, 153, 507–510. P OTREL , A., P EUCAT , J.-J. & F ANNING , C. M. 1998. Archaean crustal evolution of the West African Craton: example of the Amsaga Area (Reguibat Rise). U–Pb and Sm–Nd evidence for crustal growth and recycling. Precambrian Research, 90, 107–117. P OUCLET , A., A LLYVALY , M., D AOUDA -Y AO , B. & E SSO , B. 2004. Dećouverte d’un diatre`me de kimberlite diamantife`re a` Se´gue´la en Coˆte-d’Ivoire. Comptes Rendus de Geóscience, 336, 9–17.

P OUPEAU , G., F ABRE , J., L ABRIN , E., A ZDIMOUZA , A., N ETTO , A.-M. & M ONOD , Th. 1996. Nouvelles datations par traces de fission de la structure circulaire des Richat (Mauritanie). In: BITAM , L. & FABRE , J. (eds) Geodynamique du craton ouest africain central et oriental: he´ritage et e´volution post-panafricains. Me´moires du Service Geólogique de l’Alge´rie, 8, 231–236. R IEDEL , W. 1929. Zur Mechanik geologischer Brucherscheinungen. Zentrablatt fu¨r Mineralogie, Geologie und Pa¨laeontologie, 3, 354– 368. R IMI , A. 1999. Mantle heat flow and geotherms for the main geologic domains in Morocco. International Journal of Earth Sciences, 88, 458–466. R OMBOUTS , L. 2003. Distribution of diamond and kimberlites on the Reguibat craton, Mauritania. In: Extended Abstracts, 8th International Kimberlite Conference, FLA 0034, Victoria, BC, Canada, 22– 27 June. R OUSSEL , J. & L ESQUER , A. 1991. Geophysics and the crustal structure of West Africa. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 9 –28. S ABATE´ , P. 1973. La jointure Yetti–Eglab dans la dorsale prećambienne du pays Reguibat (Sahara Occidental Alge´rien). Comptes Rendus de l’Acade´mie des Sciences, 276, 2237–2239. S ABATE´ , P. & L AMEYRE , J. 1973. Magmatism and metamorphism in the Yetti–Eglab Precambrian formations of the Reguibat Dorsale (Occidental Algerian Sahara). Travaux du Laboratoire des Sciences de la Terre, Marseille, B1, 131–133. S ABATE´ , P. & L OMAX , K. 1975. Donneés stratigraphiques et paleómagne´tiques de la re´gion Yetti– Eglab (Sahara occidental alge´rien). Bulletin du Bureau de la Recherche Geólogique et Minie`re, II, 293–311. S CHULZE , D. J. 2003. A classification scheme for mantlederived garnets in kimberlites: a tool for investigating the mantle and exploring for diamonds. Lithos, 7, 195–213. S COTT S MITH , B. H. 1996. Kimberlites. In: M ITCHELL , R. H. (ed.) Undersaturated Alkaline Rocks: Mineralogy, Petrogenesis, and Economic Potential. Mineralogical Association of Canada, Short Course Series, 24, 217–242. S EBAI , A., F ERAUD , G., B ERTRAND , H. & H ANES , J. 1991. 40Ar/39Ar dating and geochemistry of tholeiitic magmatism related to early opening of the Central Atlantic rift. Earth and Planetary Science Letters, 104, 455– 472. S OBOLEV , N. V. 1974. Deep-seated inclusions in kimberlites and the problem of the composition of the upper mantle. Nauka, Novosibirsk [in Russian]. S OBOLEV , N. V., A FANASYEV , V. A., P OKHILENKO , N. P., K AMINSKY , F. V., T ARASYUK , O. N. & H ENNI , A. 1992. Pyropes and diamonds from Algerian Sahara. Doklady Akademii Nauk SSSR, 325, 367–372. S OUGY , J. 1954. Rapport de fin de campagne 1953– 1954 (Feuilles El Mzereb, Chegga, Tindouf). Rapport ine´dit de la Direction Fe´de´rale de Mines et de la Geólogie. AOF, Dakar.

NORTH AFRICAN DIAMONDIFEROUS PROVINCE S OUGY , J. 1960. Les se´ries prećambriennes de la Mauritanie nord-orientale (A.O.F.). Det Berlingske Bogtrykkeri, Copenhagen, 59– 68. S UN , S. S. & M C D ONOUGH , W. F. 1989. Chemical and isotopic systematics of oceanic basalts: implication for mantle composition and processes. In: S AUNDERS , A. D. & N ORRY , M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications, 42, 313–345. T EGYEY , M. & J OHAN , V. 1989. Une se´quence komatiitique dans le Prote´rozoique inferieur de Guineé (Afrique de l’Ouest): caracte`res pe´trographiques, mine´ralogiques et geóchimiques. Comptes Rendus de l’Acade´mie des Sciences, 308, 193 –200. T HE´ BAULT , J. Y. 1959. Proble`me de la recherche du diamant en pays Saharien et plus prećise´ment dans le Hoggar. Bulletin des Sciences Economiques, Bureau de Recherche Minie`re de l’Alge´rie, 6, 65– 81. T OKARSKI , A. 1991. Tectonics of Hank sequence (Upper Proterozoic) in the eastern part of Eglab Massif, Reguibat shield (West African Craton). Journal of African Earth Sciences, 12, 555–560.

109

T OUAHRI , B., F ABRE , J., P IBOULE , M. & K ADDOUR , M. 1996. Les diamants du Bled El Mass (Touat): Contexte geólogique. In: BITAM , L. & FABRE , J. (eds) Geodynamique du craton ouest africain central et oriental: he´ritage et e´volution post-panafricains. Me´moires du Service Geólogique de l’Alge´rie, 8, 259– 272. V ILLEMUR , J. R. 1967. Reconnaissance geólogique et structurale du Nord du bassin de Taoudeni. Me´moires du Bureau de Recherche Geólogique et Minie`re, 51. V ILLENEUVE , M. 2005. Paleozoic basins in West African and the Mauritanide thrust belt. Journal of African Earth Sciences, 43, 166 –195. V ILLENEUVE , M. & C ORNE´ E , J. J. 1994. Structure and palaeogeography of the West African and bordering belts during the Neoproterozoic. Precambrian Research, 69, 307– 326. W HITE , S. H., B OORDER , H. & S CHMIDT , C. B. 1995. Structural controls of kimberlite and lamproite emplacement. Journal of Geochemical Exploration, 53, 245– 264.

Geochronology and metamorphic P – T – X evolution of the Eburnean granulite-facies metapelites of Tidjenouine (Central Hoggar, Algeria): witness of the LATEA metacratonic evolution ABDERRAHMANE BENDAOUD1, KHADIDJA OUZEGANE1, GASTON GODARD2, JEAN-PAUL LIE´GEOIS3, JEAN-ROBERT KIENAST4, OLIVIER BRUGUIER5 & AMAR DRARENI1 1

Faculte´ des Sciences de la Terre, de la Geógraphie et de l’Ameńagement du Territoire, USTHB, BP 32, Dar el Beida 16111, Alger, Alge´rie (e-mail: [email protected]) 2

Equipe Geóbiosphe`re actuelle et primitive, CNRS IPGP, Universite´ Paris 7-Denis Diderot, 4 place Jussieu, case 89, Paris Cedex 05, France 3

Isotope Geology, Africa Museum, B-3080 Tervuren, Belgium

4

Laboratoire de Geósciences Marines, UFR des Sciences Physiques de la Terre, Universite´ Paris 7-Denis Diderot, UMR 7097, 4 place Jussieu, Tour 14, 5ie`me Etage Paris Cedex 05, France 5

ISTEEM-CNRS, cc 056, Universite´ de Montpellier II, Place Euge`ne Bataillon, F-34095 Montpellier, France Abstract: Central Hoggar, within the Tuareg shield to the east of the West African craton, is known for its complexity owing to the interplay of the Eburnean and Pan-African orogenies. The Tidjenouine area in the Laouni terrane belongs to the LATEA metacraton and displays spectacular examples of granulite-facies migmatitic metapelites. Here, we present a detailed petrological study coupled with in situ U –Pb zircon dating by laser-ablation inductively coupled plasma mass spectrometry (ICP-MS) that allows us to constrain the relative role of the Eburnean and Pan-African orogenies and hence to constrain how the LATEA Eburnean microcontinent has been partly destabilized during the Pan-African orogeny; that is, its metacratonic evolution. These metapelites have recorded different metamorphic stages. A clockwise P–T evolution is demonstrated on the basis of textural relationships, modelling in KFMASH and FMASH systems and thermobarometry. The prograde evolution implies several melting reactions involving the breakdown of biotite and gedrite. Peak metamorphic P –T conditions of 860 + 50 8C and 7 –8 kbar (M1) were followed by a decrease of pressure to 4.3 + 1 kbar and of temperature to around 700 8C, associated with the development of migmatites (M2). After cooling, a third thermal phase at c. 650 8C and 3– 4 kbar (M3) occurred. U– Pb zircon laser ablation ICP-MS analysis allows us to date the protolith of the migmatites at 2151 + 8 Ma, the granulite-facies and migmatitic metamorphisms (M1 –M2) at 2062 + 39 Ma and the medium-grade metamorphic assemblage (M3) at 614 + 11 Ma. This last event is coeval with the emplacement of large Pan-African granitic batholiths. These data show that the main metamorphic events are Eburnean in age. The PanAfrican orogeny, in contrast, is associated mainly with medium-grade metamorphism but also mega-shear zones and granitic batholiths, characterized by a high temperature gradient. This can be considered as typical of a metacratonic evolution.

The Tidjenouine metapelites (Central Hoggar, Fig. 1) show a great diversity of minerals (garnet, biotite, quartz, sillimanite, gedrite, corundum, orthopyroxene, cordierite, spinel, feldspar, plagioclase, ilmenite, rutile) forming different assemblages depending on whole-rock composition and extent of metamorphic transformation. The rocks were involved in a prograde metamorphic evolution

followed by decompression. Granulite-facies metamorphism was accompanied by melting favoured by biotite or gedrite dehydration. The successive stages of melting, with a progressively increasing amount of melt escape, produced metapelites with a restitic composition. In these rocks, corundum, spinel and sillimanite crystallized in the most All rich microdomains and orthopyroxene in the


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Fig. 1. Geological sketch maps of the Hoggar (a, Bertrand et al. 1986), of the Tuareg shield (b, Black et al. 1994) and geological map of the study area (c, Lie´geois et al. 2003). Eg-Al, Ege´re´ –Aleksod; Te, Tefedest; Az, Azrou-n-Fad; Se, Serouenout; Is, Issalane; La, Laouni; Isk, Iskel; It, In Teidini; Tz, Tazat; As-Is, Assode´-Issalane.

most Mg-rich zones. In Central Hoggar, this prograde metamorphism in granulite facies has never been described and the large variability of the metapelite compositions allows us to constrain the P– T–aH2O evolution. On the other hand, the Tuareg shield is characterized by the interplay of the Eburnean (c. 2 Ga) and the Pan-African (c. 0.6 Ga) orogenies. Several terranes of this shield were mostly generated during the Pan-African orogeny (Black et al. 1994) whereas others have been only slightly affected, such as the In Ouzzal terrane (Ouzegane et al. 2003, and references therein), perfectly preserving ultrahigh-temperature parageneses (Ouzegane & Boumaza 1996; Adjerid et al. 2008). The situation of Central Hoggar is much more debated: for some researchers (e.g. Caby 2003), its granulite-facies metamorphism is Pan-African in

age (protoliths being mostly Palaeoproterozoic or Archaean); for others, this metamorphism is Eburnean in age, the Pan-African orogeny having generated only high-T greenschist- or amphibolite-facies metamorphism, with high-pressure metamorphism being present only in Neoproterozoic oceanic material thrust on the granulitic basement constituting the LATEA metacraton (Lie´geois et al. 2003; Peucat et al. 2003). This debate sharply emphasizes the question of how a cratonic basement behaves during an orogeny and how it can be remobilized and what are the consequences of such behaviour. This questions also the nature of the LATEA microcontinent: craton, metacraton or mobile belt? To tackle this problem, this paper focuses on the well-preserved granulites of the Tidjenouine area. It aims at (1) reconstructing the thermotectonic evolution of

TIDJENOUINE METAPELITES EVOLUTION

these granulites on the basis of detailed mineralogical and paragenetic study of diverse reaction textures preserved in the metapelites; (2) dating the metamorphic assemblages deciphered. For this purpose, a large number of samples of metapelites have been collected and studied. The P–T conditions and P– T path were constrained by using textural relationships, thermobarometry and appropriate petrogenetic grids and P –T pseudosections. The resulting constrained P–T paths, coupled with additional field relationships, allow us to interpret properly the different U – Pb zircon ages provided by laser ablation inductively coupled plasma mass spectrometry (ICP-MS) and by the older conventional U –Pb bulk zircon method. Finally, this allows us to propose a geodynamical evolution of the LATEA microcontinent, highlighting a metacratonic evolution.

Regional geology and lithology The Tidjenouine area (Central Hoggar, Algeria; Fig. 1) is located in the NW part of the Laouni terrane (Fig. 1b), one of the 23 terranes of the Tuareg shield that were amalgamated during the Pan-African orogeny (Black et al. 1994). The Laouni terrane is composed of a granulite- to amphibolite-facies basement separated from Pan-African lithologies by mega-thrusts, such as the Tessalit ophiolitic remnant in the south and the eclogite lenses and associated oceanic material in the Tin Begane area (Lie´geois et al. 2003). The Laouni terrane is one of the four terranes constituting the LATEA micro-continent (LATEA is an acronym of Laouni, Azrou-n-fad, Tefedest and Ege´re´-Aleksod terranes; Fig. 1b). According to Lie´geois et al. (2003), the Archaean and Eburnean LATEA microcontinent was dismembered by mega-shear zones and intruded by granitic batholiths during the main episode of the Pan-African orogeny (640 –580 Ma). The granulite-facies rocks of the Tidjenouine area are composed of two units: (1) migmatitic gneisses with locally recognizable metapelitic and metabasic lenses; (2) migmatitic biotite –garnet – sillimanite metapelites interbanded with olivine– spinel marbles, sillimanite-bearing quartzites and metabasic layers. The quartzites form 100 m thick folded ridges, whereas the marbles occur as boudin alignments, a few metres in thickness. All these rocks are crosscut by Pan-African granites. At contacts between marbles and granites, skarns can be observed. The granulite-facies metamorphism is accompanied by subhorizontal foliations and tangential tectonics. Few geochronological data are available in the Laouni terrane: these include the following: (1)

113

the Pan-African Anfeg granitic batholith has been dated at 608 + 7 Ma (U– Pb zircon, Bertrand et al. 1986; recalculated by Lie´geois et al. 2003); (2) the Pan-African amphibolite-facies metamorphism of the thrust oceanic material at Tin Begane has been dated at 685 + 20 Ma (Sm –Nd mineral isochron; Lie´geois et al. 2003); (3) a granulite and a migmatitic granite in the Tidjenouine area have been dated to Eburnean ages of 1979 + 33 Ma and 2038 + 15 Ma (U –Pb zircon, Bertrand et al. 1986; recalculated by Lie´geois et al. 2003). A migmatite from the neighbouring Azrou n’Fad terrane gave strongly discordant zircons with an upper intercept of 2131 + 12 Ma and a lower intercept of 609 + 17 Ma (Barbey et al. 1989), thus the age of the migmatitization is ambiguous. The c. 2 Ga ages are interpreted either as the age of the protoliths and the granulitefacies metamorphism (Bertrand & Jardim de Sa´ 1990; Ouzegane et al. 2001; Lie´geois et al. 2003) or as the age of the protoliths, the granulitefacies metamorphism being Pan-African in age (Barbey et al. 1989; Caby 2003). Other workers have indicated that they cannot choose between the two hypotheses (Bendaoud et al. 2004; Benyahia et al. 2005). Three arguments sustain an Eburnean age for the granulite-facies metamorphism: (1) the zircons dated by Bertrand et al. (1986) in the Tidjenouine area have not recorded the Pan-African orogeny; (2) in the Gour Oumelalen region (NE LATEA), a series of granulitic rocks have been dated both by the conventional and ion microprobe U –Pb on zircon methods, and an age of c. 1.9 Ga has been inferred for the metamorphism without any record of the Pan-African orogeny (Peucat et al. 2003); (3) the c. 685 Ma old eclogite- and amphibolite-facies oceanic material has not been affected by the granulitic metamorphism. However, this controversial issue must be resolved by a detailed study of the metamorphic phases and by in situ zircon dating of key lithologies.

Main characteristics of the Tidjenouine migmatitic granulites The main Tidjenouine rock type is a medium- to coarse-grained migmatitic orthogneiss made of quartz þ K-feldspar þ plagioclase þ biotite with minor amounts of garnet. The metapelites that will be described in this study are less abundant. In the central part of the area, the orthogneisses are mainly leucomigmatites surrounded by darker migmatitic gneiss. Their silica values range from 66.4 to 76.1 wt% and the Mg/(Mg þ Fe) ratio varies between 0.35 and 0.52. Aluminium saturation index (ASI, A/CNK) values between

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1.1 and 1.3 indicate strongly peraluminous compositions (Table 1). Most REE patterns (Fig. 2) of these migmatites show pronounced depletion in heavy REE (HREE), which is a characteristic of

magmatic suites that have garnet in their source (Hanson 1989). Some samples, however, have flatter HREE patterns. Ba occurs in the 856–1825 ppm range. Sr (334–533 ppm) and Rb

Table 1. Representative geochemical data for migmatitic gneiss and metapelites from Tidjenouine area Rock type:

Type A

Type C

Type D

Migmatitic gneiss

Sample:

TD 39

TD 60

Tj 58

TD 67

Tj 5

Tj 80

Tj 139

Tj 120

SiO2 TiO2 Al2O3 FeO* FeO Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Sum

69.88 0.78 13.68 6.7

63.62 1.13 16.81 6.2

43.29 0.98 25.76 12.1

60.4 1.35 14.68 13.18

69.11 0.87 13.69 6.15

66.42 0.69 16.15 3.63

68.96 0.58 15.62 5.13

76.78 0.06 13.1 1.55

7.45 0.08 2.05 0.45 0.92 2.56 0.04 1.8 99.69

6.89 0.08 2.06 3.68 3.04 2.18 0.51

13.45 0.09 10.2 0.33 0.3 2.52 0.09 2.88 99.89

14.65 0.15 6.49 1.23 0.55 0.16 0.33

4.03 0.03 1.83 2.45 3.72 3.43 0.26 0.92 99.93

5.7 0.04 3.84 0.64 0.89 1.6 0.1 1.97 99.94

1.73 0.05 0.47 0.56 2.29 4.84 0.12

99.73

6.84 0.16 0.84 3.2 2.41 2.09 0.25 0.28 99.74

1.77 168 618 29.6 1.82 1.21 18.9 73 149 11.3 17.1 49.2 64.6 12.9 336 8.2 0.729 9.37 1.02 4.82 24.8 0.816 41.8 2.25 0.322 2.18 0.349 9.9

0.143 5.67 114 0.71 0.39 0.557 11.15 28.87 63.28 2.15 7.94 40 32.65 7.5 184 4.25 1.26 8.13 1.29 8.72 62.3 2.32 7.13 6.37 1.42 10.1 1.74 16.1 0.25 102 28.1 20.2 159 8.98

1.56 86.34 1121 13.53 1.12 1.01 13.04 58.63 119.5 14.3 13.63 188 51.24 10.27 362 9.46 1.92 8.29 1.3 6.91 42.4 1.46 20.7 4.06 0.55 4.17 0.66 29.5 0.23 50 69.2 13.4 50.5 17.3

2.06 116.85 855.76 50.87 2.44 0.39 7.01 142.05 261.17 31.53 27.76 334.47 97.06 14.42 693.38 16.57 2.08 8.11 0.88 3.21 12.67 0.41 26.9 1.35 0.14 1.11 0.13 8.66 0.35 77.18 85.32 11.3 56.3 29.17

0.38 75.5 262.64 13.31 1.07 0.4 9.01 54.67 104.18 5.11 11 81.74 38.66 6.81 232.33 5.28 0.82 5.33 0.73 3.71 18.68 0.58 22.61 1.5 0.19 1.24 0.14 12.64 0.09 71.84 66.8 14.89 87.6 34.68

Cs Rb Ba Th U Ta Nb La Ce Pb Pr Sr Nd Sm Zr Hf Eu Gd Tb Dy Y Ho Ga Er Tm Yb Lu Cu Cd V Zn Co Cr Ni

0.123 91.72 670 12.35 0.71 0.392 9.44 55.35 108.8 5.66 11.34 102 42.25 7.19 254 6.05 1.46 5.22 0.79 4.17 22.4 0.86 19.5 2.2 0.39 2.31 0.36 39.6 0.08 118 50.6 18.9 108 53.4

LOI, loss on ignition.

100 92 912

15

344 278

43 24

28 103 105 25 52 30

162 231 17.3 137 66.1

100 0.22 92.53 1825.35 8.74 0.63 0.1 1.36 33.45 64.27 28.72 6.77 533.49 23.84 4.48 130.21 3.6 1.75 3.19 0.39 2.1 16.57 0.43 13 1.09 0.16 1.15 0.12 5.35 0.1 6.19 17.52 2.97 9.85 2.73


(a) 1000

Opx Free Metapelite Opx Bearing Metapelite Gedrite Bearing Granulite

115

(b) 1000

100

Rock/Primitive Mantle

Rock/Chondrites

100

10

10

1

0.1

Opx Free Metapelite Opx Bearing Metapelite Gedrite Bearing granulite

1

0.01 La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

(c) 1000

(d) Tj 5 Tj 80 Tj 120 Tj 139

100

10

P Nd Zr Hf Sm Eu Ti Gd Dy Y Er Yb Lu V Fe Co Mg Cr Ni

1000

100

Rock/Primitive Mantle

Migmatitic gneiss

Rock/Chondrites

Rb Ba Th K Nb Ta La Ce Sr

10

1

0.1

Tj 5 Tj 80 Tj 120 Tj 139

Migmatitic gneiss

0.01

1 La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Rb Ba Th K Nb Ta La Ce Sr P Nd Zr Hf Sm Eu Ti Gd Dy Y Er Yb Lu V Fe Co Mg Cr Ni

Fig. 2. REE bulk-rock compositions normalized to chondrite (a, c) and spider diagram normalized to primitive mantle (b, d) for metapelites (a and b) and migmatitic gneisses (c and d), respectively. Chondrite and primitive mantle normalization values are from Taylor & McLennan (1985) and Sun & McDonough (1989), respectively.

(93–117 ppm) give Rb/Sr values between 0.17 and 1.26. The composition of these rocks suggests that their protoliths resulted from the partial melting of the continental crust, which left a garnet-bearing residue. The transition from the orthogneiss to the metapelite corresponds to a decrease in the size and abundance of the migmatitic leucosomes, which have a mineralogical composition identical to that of the orthogneiss, until their total disappearance. The orthogneisses can then be considered as sharing the same origin as the leucosomes, but being slightly more allochthonous. This indicates that the felsic intrusions, the granulitic metamorphism and the migmatitization occurred within the same event. Garnet-bearing mafic rocks occur as centimetre- to metre-sized boudins along the granulite-facies foliation within the orthogneiss. Larger bodies (hundreds of metres in size) do not bear garnet. These mafic rocks are not studied here. The garnet-bearing mafic rocks are composed of the Grt–Cpx –Pl–Qtz primary assemblage, which broke down to Opx –Pl during the

decompressional stage. On the other hand, the garnet-free mafic rocks with Opx –Cpx–Am – Pl + Qtz assemblage are characterized by later destabilization of the amphibole to Opx– Pl, after the decrease of pressure. The metapelites are migmatitic and dominantly restitic, felsic minerals being commonly less abundant than the mafic ones. A strong layering is observed: Grt–Bt –Sil– Cd-rich restitic layers alternate with Qtz– Pl + K-feldspar-rich leucosomes. Migmatitization and dehydration are generally thought to be caused by melting of hydrous phases such as biotite and, less commonly, gedrite. A major feature of the Tidjenouine migmatitic granulites is the presence of well-preserved petrological textures that have developed during prograde metamorphism (e.g. inclusions in garnet or sillimanite), as well as during retrogression (i.e. spectacular symplectites and coronas). This allows an accurate determination of the P– T path evolution. Different assemblages (Table 2) have been distinguished in the metapelites on the basis of the

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Table 2. Representative mineral assemblages of granulite-facies metapelites from Tidjenouine Rock type: Samples:

Quartz Biotite Garnet Sillimanite Plagioclase K-feldspar Gedrite Orthopyroxene Cordierite Spinel Corundum Ilmenite Rutile Apatite Magnetite Graphite Pyrite

Type A

Type B

Type C

Type D

Type E

TD 80–128 TD 63 TD 65 TD 29 TD 39 TD 159

TD 130 TD 134 TD 56

TD 57 TD 57b TD 59

TD 67 TD 67C TD 67A

TD 38

X X X X X x

X

X X X

X X X X X X X X X x

X X X x X X X X X

X

X X X

X X X x X X X X

X x x

X X x

X X

X

X x X X X X x x

X, abundant; x, scarce.

presence or absence, depending on the protolith composition, of phases such as orthopyroxene, gedrite, biotite, sillimanite, corundum or quartz. For example, the peak paragenesis of the most Fe-rich metapelites is garnet þ sillimanite þ quartz þ biotite þ cordierite þ plagioclase + Kfeldspar, whereas the most Mg-rich metapelites have orthopyroxene þ garnet þ biotite þ quartz þ cordierite þ plagioclase + K-feldspar. As these rocks occur intimately associated in the field, the variations in the mineral assemblage are controlled by the bulk composition of the rocks rather than by P–T conditions. Five main assemblages (Table 2), ranging from Fe-rich to Mg-rich compositions, have been distinguished: orthopyroxene-free quartz-bearing metapelites (type A, Table 2); orthopyroxene-free corundum-bearing metapelites (type B, Table 2); secondary orthopyroxene-bearing metapelites (type C, Table 2); gedrite-bearing granulites (type D, Table 2); sillimanite-free orthopyroxene-bearing metapelites (type E, Table 2). Migmatitic gneisses without sillimanite and orthopyroxene with flat HREE patterns are similar in composition to orthopyroxene-free quartz-bearing metapelites. The leucosome-free metapelites have compositions typical of residual rocks: for instance, the secondary orthopyroxene-bearing rocks show high contents of FeO, MgO and Al2O3 and low contents of SiO2, K2O and Na2O, leading to high normative corundum (up to 12 wt%). These rocks are enriched in

light REE (LREE) and display negative Eu anomalies (Eu/Eu* ¼ 0.50, Fig. 2).

The orthopyroxene-free quartz-bearing metapelites (type A) The orthopyroxene-free quartz-bearing metapelites display medium to coarse granoblastic texture with a layered structure. They consist mainly of garnet, biotite, cordierite, sillimanite, quartz and K-feldspar porphyroblasts, with subordinate plagioclase, spinel, ilmenite and graphite. Rutile, zircon and apatite are accessory phases. All the primary minerals (garnet, sillimanite, biotite and quartz) are deformed. A large variability in proportions exists from leucocratic varieties rich in quartz and K-feldspar to melanocratic varieties rich in garnet, biotite and sillimanite where quartz and K-feldspar are absent. The modal percentage of garnet varies between 2 and 25 vol%. The core of the garnet porphyroblasts frequently contains inclusions of biotite, sillimanite, quartz and plagioclase. This feature suggests the prograde reaction Bt þ Sil þ Qtz + Pl ! Grt þ Melt + Crd + Kfs + Ilm:

ð1Þ

In all samples, biotite and sillimanite are never in contact, because there are always separated by


symplectites, as a result of later reactions between them. In the presence of quartz, we observe the growth of cordierite with sometimes K-feldspar from the assemblage biotite þ sillimanite þ quartz, where sillimanite occurs both as porphyroblasts and as fine needles included in cordierite cores. In the absence of quartz, symplectites of cordierite þ spinel and K-feldspar developed on the interfaces between primary biotite and sillimanite. This corresponds to the reactions Bt þ Sil þ Qtz ! Crd + Kfs þ Melt

ð2Þ

and Bt þ Sil ! Crd þ Spl + Kfs þ Melt:

ð3Þ

In a similar way, garnet porphyroblasts, in the presence of sillimanite and quartz, have been partly resorbed, being surrounded by cordierite; in quartz-free microdomains containing garnet and sillimanite, we observe the growth of cordierite toward garnet and of cordierite þ spinel symplectites around sillimanite grains. These textures suggest the reactions Grt þ Sil þ Qtz ! Crd

ð4Þ

Grt þ Sil ! Crd þ Spl.

ð5Þ

and

Garnet is also occasionally observed in another mineral assemblage where it occurs as euhedral grains with cordierite as result of the reaction (4), which becomes Grt1 þ Sil þ Qtz ! Crd þ Grt2 :

ð40 Þ:

In some samples, a late sillimanite replaced primary sillimanite, crosscutting the foliation defined by the other phases composing the rock.

reaction Grt þ Cor þ Melt þ Ksp ! Sill þ Spl þ Bt: ð6Þ The sillimanite is separated from garnet and biotite by cordierite þ spinel þ K-feldspar symplectites. This texture may be explained by the KFMASH univariant reaction (Fig. 3a) Grt þ Bt þ Sil ! Spl þ Crd + Ksp þ Melt: ð7Þ These symplectites also occur on contacts of garnet or biotite with sillimanite corresponding to multivariant KFMASH equilibria (3) and (6).

Secondary orthopyroxene-bearing metapelites (type C) Metapelites with secondary orthopyroxene are melanocratic, aluminous and consist of alternating quartz-rich and silica-undersatured sillimanite-rich layers. They are coarse-grained heterogeneous rocks with granoblastic texture and are mainly composed of sillimanite, cordierite, garnet, biotite, spinel, orthopyroxene, quartz, plagioclase and smaller amounts of ilmenite, rutile, graphite and pyrite. They are characterized by largest abundance of plagioclase with respect to K-feldspar and by spectacular crystals of sillimanite up to 10 cm in length. The garnet porphyroblasts have the same inclusions (biotite, sillimanite quartz and plagioclase) as those of the orthopyroxene-free quartz-bearing metapelites and reaction (1) should have also operated. The breakdown of biotite in the presence of garnet with sillimanite or quartz produced symplectites of spinel þ cordierite and cordierite þ orthopyroxene, respectively. These textures may be explained by the reactions Grt þ Bt þ Sil ! Spl þ Crd þ Ksp þ Melt

Orthopyroxene-free corundum-bearing metapelites (type B) Corundum-bearing (quartz-free) metapelites are also the rocks richest in garnet and sillimanite. They form centimetre-sized layers. Sillimanite associated with spinel and biotite is rich in inclusions of garnet and corundum representing relics of the earlier paragenesis (Fig. 3b). This textural relationship suggests the existence of a very early melt and the corundum-consuming prograde

117

ð7Þ

(Fig. 3c) and Grt þ Bt þ Qtz ! Crd þ Opx þ Ksp þ Melt: ð8Þ One sample (Tj57b) displays the breakdown of garnet to orthopyroxene, cordierite, spinel and plagioclase according to the reaction (Fig. 3d) Grt ! Opx þ Spl þ Crd þ PlðAn96 Þ:

ð9Þ

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A. BENDAOUD ET AL.

Fig. 3. Representative reaction textures of orthopyroxene-free metapelites (a –d) and orthopyroxene-bearing metapelites (e–h). (a) Primary, elongated garnet, sillimanite and biotite reacting out to cordierite–spinel in orthopyroxene-free, corundum-bearing metapelites (backscattered electron (BSE) image). (b) The same rock with sillimanite enclosing garnet, corundum and ilmenite (plane-polarized light). (c) Well-developed spinel– cordierite symplectite close to sillimanite and cordierite corona between garnet and biotite, suggesting prograde reaction


In the same sample, garnet reacted with quartz and sometimes with rutile inclusions to produce orthopyroxene and cordierite symplectites associated with ilmenite: Grt þ Qtz + Rut !Opx þ Crd + Ilm + Pl:

119

symplectites associated with plagioclase or melt. These features should correspond to the reactions Sil þ Ged þ Qtz ¼ Grt þ Crd þ Melt

ð13Þ

and ð10Þ

In the absence of biotite, garnet and sillimanite reacted to produce cordierite, spinel and quartz symplectites following the univariant FMASH reaction

Sil þ Ged ¼ Crd þ Spl þ Melt.

ð14Þ

In some microdomains, a corona of later cordierite separates spinel from quartz, suggesting the reaction

Other reaction textures in these rocks are similar to those of the orthopyroxene-bearing metapelites. Garnet, quartz and sillimanite are never observed in contact and are always separated either by symplectitic or coronitic textures corresponding to reactions (4) and (5) with the implication of plagioclase (Fig. 3g). At the contact between garnet and quartz, quartz is rimmed by a corona of orthopyroxene, whereas garnet is mantled by a cordierite þ orthopyroxene symplectite (Fig. 3g and h):

Spl þ Qtz ! Crd:

Grt þ Qtz ¼ Opx þ Crd + Pl2 :

Grt þ Sil + Melt !Crd þ Spl þ Qtz + Ksp:

ð11Þ

ð12Þ

Gedrite-bearing granulites (type D) The gedrite-bearing rocks contain a quartz þ garnet þ sillimanite þ cordierite þ orthopyroxene þ plagioclase þ spinel þ gedrite þ ilmenite þ rutile assemblage with very minor K-feldspar and biotite. They display heterogranular coarse-grained texture with spectacular coronitic and symplectitic associations. Porphyroblasts of garnet, sillimanite, quartz, gedrite, rutile and ilmenite are systematically separated by fine symplectites of orthopyroxene þ cordierite + plagioclase + orthoamphibole or of cordierite þ spinel. The orthopyroxene occurs also as coronas entirely surrounding quartz or ilmenite. Plagioclase associated with quartz is antiperthitic. Quartz occurs as discontinuous ribbons that form lenses with asymmetric tails. The texture suggests two successive reactions: sillimanite, gedrite and quartz are separated by cordierite, plagioclase and garnet corona structures (Fig. 3e); sillimanite reacted with gedrite giving cordierite–spinel

ð10Þ

Locally, the orthopyroxene–cordierite symplectites are accompanied by secondary orthoamphibole and this reaction becomes (Fig. 3f) Grt þ Qtz + Pl1 ¼ Opx þ Crd þ Oam + Pl2 :

ð15Þ

Opx-bearing sillimanite-free metapelites (type E) Orthopyroxene-bearing, sillimanite-free metapelites are distinctly marked by the absence of sillimanite and the presence of orthopyroxene as primary phase. They show quartz–plagioclase–K-feldspar microdomains corresponding to leucosome. These rocks are coarse-grained with a polygonal granoblastic texture associated with a undulose extinction of quartz and kink-bands of biotite. This suggests deformation at a high temperature, contemporaneous with the granulite-facies metamorphism. The

Fig. 3. (Continued) Grt þ Sil þ Bt ! Crd þ Spl þ Ksp (plane-polarized light). (d) Development of complex Opx þ Spl þ Crd þ Pl intergrowths in cracks of garnet (plane-polarized light). (e) Gedrite originally in contact with primary quartz (included in garnet) and sillimanite, now enclosed by multiple coronae of phase products: plagioclase and quartz after melt, cordierite þ spinel replacing sillimanite, and orthopyroxene þ cordierite symplectite close to garnet (BSE image). This complex textural relationships suggests the following successive reactions: (1) Ged þ Sil þ Qtz ! Grt þ Crd þ Melt; (2) Ged þ Sil ! Spl þ Crd þ Melt; (3) Grt þ Qtz ! Opx þ Crd. (f) Close-up view of garnet and quartz breakdown to cordierite þ orthopyroxene þ orthoamphibole (BSE image). (g) Fine intergrowth of cordierite þ spinel þ calcic plagioclase close to sillimanite suggesting the reaction Grt þ Sil ! Crd þ Spl þ Pl2, and breakdown of garnet at quartz contact giving orthopyroxene þ cordierite. Layers of plagioclase and drops of quartz could represent melt phases (RGB image: red, Fe; green, Ca; blue, Al). (h) Garnet reacting out with quartz to orthopyroxene þ cordierite þ plagioclase (it should be noted zoning in garnet (RGB image: red, Fe; green, Ca; blue, Si).

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observed minerals are orthopyroxene–garnet– biotite–plagioclase–K-feldspar–cordierite–spinel– quartz–ilmenite–rutile–zircon and apatite. Primary orthopyroxene occurs commonly as subhedral porphyroclasts up to 1 cm in size coexisting with biotite and garnet. The presence of inclusions of biotite, quartz and garnet in the orthopyroxene suggests the prograde reaction Bt þ Qtz + Grt ! Opx þ Kfs þ Melt:

ð16Þ

The spinel is also primary and occurs both as inclusions in garnet and in textural equilibrium with the association garnet –orthopyroxene– quartz –biotite –ilmenite (Fig. 4a).

The orthopyroxene porphyroclasts have exsolved garnet and small amounts of plagioclase and ilmenite lamellae mainly along (100) and (010) crystallographic planes (Fig. 4c). This feature corresponds to the reaction High-Al Opx !Low-Al Opx þGrtðþPlþIlmÞ:

ð17Þ

This reaction is generally interpreted as being indicative of isobaric cooling (Harley 1989). Locally, orthopyroxene is destabilized in Opx – Crd symplectites according to the reaction High-Al Opx ! Low-Al Opx þ Crd:

ð18Þ

Fig. 4. Representative reaction textures of sillimanite free-metapelites. (a) Photomicrograph showing two successive parageneses (plane-polarized light). The primary assemblage is composed of spinel in equilibrium with quartz, garnet, biotite and orthopyroxene surrounded by secondary symplectites of cordierite þ orthopyroxene2 þ spinel2. (b) Late reaction observed between an inclusion of garnet and primary orthopyroxene giving very fine orthopyroxene þ cordierite symplectites. (c) Close-up view of exsolved garnet in orthopyroxene showing two preferential directions. The presence of plagioclase and ilmenite as exsolutions in primary orthopyroxene should noted (BSE image). (d) Biotite and garnet breakdown to complex intergrowths of spinel þ orthopyroxene þ cordierite þ plagioclase and secondary biotite (BSE image).


Symplectites of cordierite –orthopyroxene– spinel – plagioclase and minor biotite, ilmenite, and magnetite occur between garnet and biotite, suggesting the reaction (Fig. 4d) Grt þ Bt1 ! Crd þ Opx þ Spl þ Pl þ Mt þ Ilm þ Bt2 :

ð19Þ

The later reactions are marked by very fine-grained symplectites of orthopyroxene and cordierite surrounding garnet and primary orthopyroxene (Fig. 4b): Grt þ Opx1 ! Opx2 þ Crd:

ð20Þ

Mineral chemistry Representative analyses are listed in Table 3. The analyses have been performed with a CAMEBAX microprobe at the CAMPARIS centre (CNRS, Paris). The operating conditions were 15 kV accelerating voltage and 10 nA sample current. Natural silicates and synthetic oxides were used as standards for all elements, except for fluorine, which has been calibrated on fluorite. Some volumetric proportions of various phases have been determined (e.g. orthopyroxene and exsolved phases). To reconstruct the original composition of orthopyroxene before exsolution, we adopted the following procedure: (1) processing of the images made by the X-ray maps (22 500 mm2) generated by the scanning electron microscope; (2) conversion of the obtained volumetric proportions in molar proportions by weighting molar volumes (data from Holland & Powell 1990); (3) using the phase compositions measured by the microprobe, calculation of the cation proportions and of the oxide weight per cent. For the microprobe scanning, during each analysis, the electron beam scanned a surface of 180 mm2 (12 mm 15 mm); 250 analyses were carried out on adjacent areas and averaged. During calibration, standards were analysed with the same beam conditions (scanning of a 180 mm2 area). Garnet, in orthopyroxene-free quartz-bearing metapelites (type A, Table 3), is an almandine (64–82 mol%) rich in pyrope (12–30%) and poor in grossular and spessartine (both at 3–4 mol%). In the cores, the XFe value ranges from 0.68 in the melanosome to 0.78 in the rare grains present in the leucosome; there is a progressive increase of XFe towards the rims (to 0.86). Small euhedral garnet grains within cordierite are unzoned and have the same composition as the coarse-grained garnet rims.

121

In orthopyroxene-free corundum-bearing metapelites (type B, Table 3), garnet porphyroblasts have an XFe between 0.71 (core, Alm67Py27Gr2Sps4) and 0.84 (rim, Alm76Py15Gr2Sps7), whereas garnet inclusions in sillimanite have an XFe of 0.75 (Alm70Py23Grs2Sps5). In secondary-orthopyroxene-bearing metapelites and gedrite-bearing granulites (types C and D, Table 3), garnet is an almandine–pyrope solid solution and shows significant XFe zoning with Fe-rich rims (Fig. 5). The largest core–rim difference (from 0.49 to 0.72) is observed in the garnet found in quartz-rich microdomains; in spinel-bearing domains, XFe ranges only from 0.60 to 0.72. Grossular and spessartine contents are always , 3 mol%. In sillimanite-free orthopyroxene-bearing metapelites, the garnet from the matrix and that included in the orthopyroxene show an increase in XFe from 0.57 to 0.73 from core (Alm51Py39Gros7Sps4) to rim (Alm58Py28Gros7Sps6). Garnet exsolved in orthopyroxene has a homogeneous composition (Alm53Py37Gros7Sps5) with a XFe of 0.59. Biotite compositions are highly variable (Table 3). In orthopyroxene-free quartz-bearing metapelites (type A, Table 3), biotite inclusions in garnet have XFe in the range 0.39–0.55, with TiO2 contents between 1 and 5.72 wt% and F content , 0.2%; biotite in the matrix is richer in Fe and Ti (XFe ¼ 0.50–0.65 and TiO2 ¼ 3.65 – 7.47 wt%) and smaller biotite grains from the symplectites have lower contents of Fe and Ti. In orthopyroxene-free corundum-bearing metapelites (type B, Table 3), biotite has an XFe of 0.60–0.63 and contains generally up to 5 wt% TiO2. In the secondary orthopyroxene-bearing metapelites (type C, Table 3), biotite has XFe in the range of 0.27 –0.56, TiO2 contents between 1.27 and 6.15 wt% and F in the range 0.21–0.31 wt%; larger biotite grains in the matrix are consistently richer in Fe, Ti and F. In the sillimanite-free orthopyroxene-bearing metapelites (type E, Table 3), biotite is poorer in Fe (XFe around 0.33 in contact with orthopyroxene and around 0.26 in contact with garnet) with TiO2 and F contents around 4 wt% and 1 wt%, respectively. Three types of substitution have taken place in all assemblages: a substitution of Tschermakitic vi type, Si21(Mg,Fe)21Aliv þ1Alþ1, a substitution of titano-tschermakitic type in reverse Ti21Aliv 21Mgþ1Siþ2 and a subtitution Ti21V21(Fe,Mn,Mg)þ2 (where V is an ¼ octahedral vacancy). Cordierite shows a varying XFe that depends on the lithologies: 0.39–0.51 (orthopyroxene-free quartzbearing metapelites, type A, Table 3), 0.41–0.43 (corundum-bearing metapelites, type B, Table 3), 0.22–0.39 (secondary orthopyroxene-bearing metapelites and gedrite-bearing granulites, types C and D, Table 3) and 0.19–0.26 (sillimanite-free orthopyroxene-bearing metapelites, type E, Table 3).

122

Table 3. Chemical compositions of garnet, biotite, cordierite, orthopyroxene, orthoamphibole, spinel and plagioclase of metapelites from Tidjenouine area Biotite Rock type:

Type E

Type C

Type A

Type B

TD 38 49 i/opx

TD 38 23

TD 38 53

Tj 57b 23 c

Tj 59 61 c

Tj 59 55 s

TD 63 50 s

TD 39 109 c

TD 159 1 i/grt

Tj 130 17 c

Tj 56 86 c

Tj 56 60 s

SiO2 TiO2 Al2O3 Cr2O3 FeOt MnO MgO CaO Na2O K2O F Cl Sum

37.44 3.55 15.5 0.64 11.2 0.15 16.07 0.03 0.51 7.79 0.62 0.00 93.5

38.32 3.47 15.18 0.72 10.81 0.12 17.03 0.00 0.5 8.26 1.09 0.02 95.52

36.26 3.84 15.31 0.79 13.79 0.00 14.47 0.00 0.46 8.34 0.55 0.02 93.83

39.42 1.59 16.65 0.05 12.77 0.01 18 0.02 0.51 6.41 1.86 0.07 97.36

34.5 6.15 16.17 0.18 19.39 0.16 8.72 0.00 0.24 8.19 0.21 0.14 94.05

35.59 3.61 15.02 0.04 21.08 0.00 11.49 0.16 0.1 7.45 0.43 0.1 95.07

35.04 4.19 17.28 0.25 20.79 0.06 8.25 0.06 0.2 9.39 0.31 0.35 96.17

33.4 7.47 16.08 0.06 22.35 0.12 6.57 0.00 0.1 9.52 0.00 0.18 95.85

37.23 0.97 17.03 0.05 16.11 0.07 14.12 0.00 0.22 9.44 0.00 0.00 95.24

34.68 5.7 17.37 0.19 22.02 0.17 7.42 0.00 0.2 9.09 0.07 0.31 97.22

32.8 4.7 16.9 0.05 21.18 0.03 7.94 0.08 0.28 8.51 0.15 0.29 93.18

34.67 3.56 16.55 0.29 19.91 0.03 9.27 0.04 0.32 8.98 0.4 0.25 94.54

Si AlIV

5.554 2.446

5.562 2.438

5.462 2.538

5.565 2.435

5.334 2.666

5.447 2.553

5.354 2.646

5.195 2.805

5.576 2.424

5.267 2.733

5.208 2.792

5.378 2.622

AlVI Ti Cr Mg Fe2þ Mn Ca Na K F Cl S

0.264 0.396 0.075 3.553 1.389 0.019 0.005 0.147 1.474 0.291 0.000 15.322

0.159 0.379 0.083 3.684 1.312 0.015 0.000 0.141 1.529 0.5 0.005 15.302

0.18 0.435 0.094 3.249 1.737 0.000 0.000 0.134 1.603 0.262 0.005 15.432

0.335 0.169 0.006 3.787 1.508 0.001 0.003 0.14 1.154 0.83 0.017 15.102

0.28 0.715 0.022 2.009 2.507 0.021 0.000 0.072 1.615 0.103 0.037 15.241

0.156 0.415 0.005 2.621 2.698 0.000 0.026 0.03 1.454 0.208 0.026 15.406

0.465 0.481 0.03 1.879 2.656 0.008 0.01 0.059 1.83 0.15 0.091 15.419

0.143 0.874 0.007 1.523 2.907 0.016 0.000 0.03 1.889 0.000 0.047 15.389

0.582 0.109 0.006 3.152 2.018 0.009 0.000 0.064 1.803 0.000 0.000 15.743

0.377 0.651 0.023 1.68 2.797 0.022 0.000 0.059 1.761 0.034 0.08 15.369

0.371 0.561 0.006 1.879 2.813 0.004 0.014 0.086 1.724 0.075 0.078 15.458

0.404 0.415 0.036 2.143 2.583 0.004 0.007 0.096 1.777 0.196 0.066 15.465

XFe

0.28

0.26

0.35

0.28

0.56

0.51

0.59

0.66

0.39

0.62

0.6

0.55 (Continued)

A. BENDAOUD ET AL.

Sample: Analysis: Position:

Table 3. Continued Cordierite Rock type:

Type E

Type C

Type D

Type A

Type B

TD 38 16 /bi

TD 38 47 /opx

Tj 57 8 /opx

Tj 57b 32 /opx

Tj 57b 35 /sp

TD 67C 3 /spl

TD 67C 5 /opx

TD 39 99

TD 39 102

Tj 56 51 /sill

Tj 56 72 /grt

Tj 130 13

SiO2 TiO2 Al2O3 MgO FeOt MnO CaO Na2O K2O F Cl Sum

49.94 0.00 32.66 9.75 5.67 0.12 0.04 0.18 0.01 0.00 0.01 98.38

49.95 0.06 32.34 10.94 4.77 0.18 0.01 0.19 0.08 0.05 0.02 98.59

48.73 0.00 32.31 8.6 7.96 0.22 0.04 0.2 0.00 0.00 0.00 98.06

49.97 0.03 34.12 10.78 5.39 0.00 0.04 0.12 0.02 0.08 0.00 100.55

49.22 0.02 33.78 9.62 6.7 0.04 0.05 0.12 0.01 0.09 0.01 99.66

49.79 0.01 34.37 10.1 6.13 0.05 0.05 0.14 0.04 0.00 0.00 100.68

49.29 0.05 33.5 9.02 7.92 0.08 0.00 0.13 0.01 0.00 0.01 100.01

48.41 0.00 33.07 7.57 10.39 0.04 0.03 0.08 0.00 0.00 0.00 99.59

47.86 0.00 33.02 6.38 11.77 0.13 0.07 0.09 0.02 0.00 0.00 99.34

48 0.12 33.52 7.21 10.52 0.19 0.04 0.13 0.00 0.09 0.00 99.82

48.27 0.03 33.39 6.88 10.96 0.18 0.03 0.12 0.00 0.07 0.03 99.96

49.11 0.01 33.19 7.63 9.31 0.18 0.06 0.08 0.00 0.00 0.03 99.6

Si Ti Alt Mg Fe2þ Mn Ca Na K F Cl S

5.072 0.000 3.91 1.476 0.482 0.01 0.004 0.036 0.001 0.000 0.002 10.993

5.055 0.004 3.858 1.65 0.404 0.015 0.001 0.038 0.01 0.015 0.003 11.053

5.027 0.000 3.928 1.323 0.686 0.02 0.004 0.04 0.000 0.000 0.000 11.029

4.967 0.002 3.998 1.598 0.448 0.000 0.004 0.023 0.002 0.026 0.000 11.07

4.967 0.001 4.018 1.448 0.566 0.003 0.005 0.024 0.001 0.03 0.001 11.066

4.951 0.001 4.029 1.498 0.51 0.004 0.005 0.028 0.005 0.000 0.000 11.043

4.98 0.004 3.989 1.358 0.67 0.007 0.000 0.025 0.001 0.000 0.002 11.036

4.966 0.000 3.999 1.158 0.891 0.003 0.003 0.016 0.000 0.000 0.000 11.041

4.958 0.000 4.032 0.985 1.02 0.011 0.008 0.018 0.003 0.000 0.000 11.036

4.925 0.009 4.054 1.103 0.903 0.016 0.004 0.025 0.000 0.029 0.000 11.077

4.956 0.002 4.04 1.053 0.941 0.016 0.004 0.023 0.000 0.021 0.005 11.061

5.012 0.001 3.992 1.16 0.794 0.016 0.007 0.015 0.000 0.000 0.005 11.003

XFe

0.25

0.2

0.34

0.22

0.28

0.25

0.33

0.43

0.51

0.45

0.47

0.41



(Continued) 123

Table 3. Continued 124

Garnet Rock type: Sample: Analysis: Position:

Type B

TD 38 TD 38 TD 38 TD 38 43 14 12 6 in opx ex c r 39.75 0.07 21.67 0.41 23.65 23.31 0.38 1.73 10.16 2.54 0.02 0.02 100.06

38.59 0.00 21.41 0.68 24.68 23.92 0.85 1.96 9.87 2.4 0.03 0.00 99.71

39.53 0.02 20.99 0.52 23.95 22.88 1.19 1.81 10.07 2.74 0.03 0.00 99.78

Tj 56 67 c

Tj 56 71 r

Type C

Tj 130 Tj 130 Tj 130 Tj 57 2 9 24 8 in sill c r c

37.73 36.98 37.36 37.25 37.9 37.93 0.03 0.08 0.00 0.04 0.08 0.00 20.62 22.12 21.75 22.18 21.3 21.35 0.28 0.1 0.00 0.02 0.00 0.12 27.81 32.12 35.85 33.56 32.81 34.43 26.65 31.36 35.85 32.81 32.06 34.4 1.29 0.84 0.00 0.84 0.84 0.03 2.87 0.66 1.8 2.02 1.67 2.63 7.14 7.09 2.97 5.84 6.3 4.19 2.42 0.83 1.00 0.53 0.53 0.59 0.00 0.03 0.05 0.00 0.02 0.03 0.00 0.00 0.00 0.02 0.00 0.00 99.03 100.09 100.78 101.54 100.7 101.27

Type D

Type A

Tj 57 TD 67C TD 67C TD 128 TD 128 TD 63 10 86 27 6 7 57 r c r c r c

35.87 36.85 39.95 37.36 0.12 0.07 0.00 0.06 22.17 21.37 23.02 21.69 0.04 0.06 0.02 0.00 26.71 35.17 24.02 35.49 24.41 33.84 23.52 34.7 2.55 1.48 0.56 0.88 0.29 0.82 0.29 0.98 11.02 4.88 13.18 4.93 1.11 1.52 0.85 0.88 0.06 0.02 0.01 0.01 0.03 0.01 0.02 0.02 97.68 100.92 101.42 101.51

36.74 0.03 21.68 0.09 30.68 29.51 1.30 1.15 7.36 1.20 0.04 0.03 99.13

TD 63 66 r

38.46 38.05 36.94 0.00 0.03 0.00 21.76 22.03 21.94 0.00 0.02 0.09 34.14 34.57 36.64 34.14 34.52 36.64 0.00 0.06 0.00 3.00 0.92 1.51 3.32 5.14 3.00 1.02 1.08 0.89 0.00 0.00 0.02 0.00 0.00 0.01 101.7 101.85 101.04

Si AlIV AlVI Ti Cr Fe3þ Fe2þ Mg Mn Ca Na K S

3.033 0.000 1.949 0.004 0.025 0.022 1.487 1.155 0.112 0.208 0.003 0.002 8.00

2.968 0.032 1.910 0.000 0.041 0.049 1.539 1.131 0.128 0.198 0.004 0.000 8.00

3.033 0.000 1.899 0.001 0.032 0.069 1.468 1.152 0.118 0.225 0.004 0.000 8.00

2.982 0.018 1.904 0.002 0.017 0.077 1.762 0.841 0.192 0.205 0.000 0.000 8.00

2.897 0.103 1.939 0.005 0.006 0.05 2.054 0.828 0.044 0.07 0.005 0.000 8.00

2.986 0.014 2.035 0.000 0.000 0.000 2.396 0.354 0.122 0.086 0.008 0.000 8.00

2.907 0.093 1.947 0.002 0.001 0.049 2.141 0.679 0.134 0.044 0.000 0.002 8.00

2.975 0.025 1.946 0.005 0.000 0.05 2.104 0.737 0.111 0.045 0.003 0.000 8.00

3.000 0.000 1.991 0.000 0.008 0.002 2.275 0.494 0.176 0.05 0.005 0.000 8.00

2.800 0.200 1.840 0.007 0.002 0.150 1.594 1.282 0.019 0.093 0.009 0.003 8.00

2.913 0.087 1.904 0.004 0.004 0.088 2.237 0.575 0.055 0.129 0.003 0.001 8.00

2.958 0.042 1.968 0.000 0.001 0.031 1.457 1.454 0.018 0.067 0.001 0.002 8.00

2.935 0.065 1.944 0.004 0.000 0.052 2.28 0.577 0.065 0.074 0.002 0.002 8.00

2.899 0.101 1.915 0.002 0.006 0.077 1.947 0.865 0.077 0.101 0.006 0.003 8.00

3.039 0.000 2.027 0.000 0.000 0.000 2.256 0.391 0.201 0.086 0.000 0.000 8.00

2.968 0.032 1.993 0.002 0.001 0.004 2.252 0.597 0.061 0.09 0.000 0.000 8.00

2.947 0.053 2.011 0.000 0.006 0.000 2.445 0.357 0.102 0.076 0.003 0.001 8.00

XMg Fe3þ/Fe3þ þ Fe2þ

0.44 0.01

0.42 0.03

0.44 0.04

0.32 0.04

0.29 0.02

0.13 0.00

0.24 0.02

0.26 0.02

0.18 0.00

0.45 0.09

0.20 0.04

0.50 0.02

0.20 0.02

0.31 0.04

0.15 0.00

0.21 0.00

0.13 0.00

Alm Sps Gr Py

0.5 0.04 0.07 0.39

0.51 0.04 0.07 0.38

0.5 0.04 0.08 0.39

0.59 0.06 0.07 0.28

0.69 0.01 0.02 0.28

0.81 0.04 0.03 0.12

0.71 0.04 0.01 0.23

0.7 0.04 0.01 0.25

0.76 0.06 0.02 0.16

0.53 0.01 0.03 0.43

0.75 0.02 0.04 0.19

0.49 0.01 0.02 0.49

0.76 0.02 0.02 0.19

0.65 0.03 0.03 0.29

0.77 0.07 0.03 0.13

0.75 0.02 0.03 0.2

0.82 0.03 0.03 0.12

(Continued)

A. BENDAOUD ET AL.

SiO2 TiO2 Al2O3 Cr2O3 FeOt FeO Fe2O3 MnO MgO CaO Na2O K2O Sum

Type E

Table 3. Continued Orthopyroxene Rock type:

Type D

Type C

Type E

TD 67A 45 /crdsp

TD 67C 101 /crd

Tj 57b 38 /qz

Tj 57b 16 /sp

TD 38 14 Sympl splcrd

TD 38 40 r/crd

TD 38 16 c

TD 38 Opx I reconstituted

SiO2 TiO2 Al2O3 Cr2O3 FeOt FeO Fe2O3 MnO MgO CaO Na2O K2O Sum

49.6 0.16 4.64 0.16 25.73 25.73 0.00 0.3 18.41 0.09 0.01 0.00 99.1

49.48 0.06 2.96 0.05 31.51 31.51 0.00 0.33 15 0.14 0.05 0.01 99.59

51.01 0.06 1.46 0.00 30.54 30.54 0.00 0.22 16.66 0.17 0.00 0.00 100.12

50.8 0.1 3.4 0.01 26.05 26.05 0.00 0.16 19.28 0.19 0.00 0.00 99.99

51.62 0.22 2.49 0.26 25.75 25.75 0.00 1.13 17.39 0.13 0.01 0.02 99.02

51.26 0.08 3.4 0.34 22.58 22.58 0.00 1.26 20.53 0.12 0.03 0.00 99.6

50.75 0.1 4.85 0.54 21.36 21.36 0.00 0.6 21.55 0.08 0.03 0.00 99.86

48.27 0.37 6.24 0.55 21.53 19.57 2.18 0.95 20.26 0.4 0.05 0.22 99.06

Si AlIV AlVI Alt Ti Cr Fe3þ Fe2þ Mg Mn Ca Na K Total

1.896 0.104 0.105 0.209 0.005 0.005 0.000 0.823 1.049 0.01 0.004 0.001 0.000 4.00

1.935 0.065 0.071 0.136 0.002 0.002 0.000 1.03 0.874 0.011 0.006 0.004 0.000 4.00

1.971 0.029 0.038 0.067 0.002 0.000 0.000 0.987 0.959 0.007 0.007 0.000 0.000 4.00

1.922 0.078 0.073 0.152 0.003 0.000 0.000 0.824 1.087 0.005 0.008 0.000 0.000 4.00

1.995 0.005 0.108 0.113 0.006 0.008 0.000 0.832 1.001 0.037 0.005 0.001 0.001 4.00

1.928 0.072 0.079 0.151 0.002 0.01 0.000 0.71 1.151 0.04 0.005 0.002 0.000 4.00

1.887 0.113 0.099 0.213 0.003 0.016 0.000 0.664 1.194 0.019 0.003 0.002 0.000 4.00

1.819 0.181 0.096 0.277 0.01 0.016 0.062 0.617 1.138 0.03 0.016 0.004 0.011 4.00

XMg

0.56

0.46

0.49

0.57

0.55

0.62

0.64

0.65 125

(Continued)



126

Table 3. Continued Orthoamphibole Rock type:

Type D TD 67 117 c

TD 67 47 c

TD 67 27 c

TD 67 100 c

TD 67 63 r

TD 67 12 r

TD 67 74 r

TD 67 21 s

TD 67 32 s

TD 67 21 s

SiO2 TiO2 Al2O3 Cr2O3 FeOt MnO MgO NiO ZnO CaO Na2O K2O F Cl Sum

42.72 0.62 15.91 0.00 23.19 0.16 13.37 0.00 0.22 0.69 1.74 0.04 0.22 0.12 99.00

40.44 1.27 16.1 0.17 23.9 0.25 11.74 0.00 0.00 0.69 2.05 0.05 0.14 0.12 96.92

40.17 0.51 19.34 0.2 22.85 0.28 12.41 0.00 0.00 0.6 2.14 0.03 0.37 0.09 98.99

38.47 0.16 20.45 0.00 20.22 0.12 14.6 0.00 0.04 0.31 2.34 0.02 0.43 0.03 97.19

49.72 0.06 14.81 0.08 19.21 0.28 12.48 0.00 0.01 0.17 0.12 0.14 0.13 0.00 97.21

44.2 0.9 13.5 0.16 22.64 0.2 14.25 0.00 0.02 0.53 1.7 0.06 0.36 0.00 98.52

43.43 0.78 14.39 0.16 21.77 0.33 14.08 0.00 0.19 0.53 1.92 0.00 0.2 0.09 97.87

50.09 0.07 6.58 0.14 25.68 0.11 15.72 0.00 0.00 0.09 0.07 0.01 0.00 0.02 98.58

45.89 0.56 8.43 0.08 26.94 0.19 14.73 0.00 0.00 0.49 0.75 0.03 0.02 0.03 98.14

45.97 0.69 12.15 0.11 22.6 0.24 14.91 0.00 0.00 0.6 1.45 0.00 0.46 0.07 99.25

Si AlIV AlVI Ti Cr Mg Fe2þ Mn Ni Zn Ca Na K F Cl P

6.275 1.725 1.03 0.069 0.000 2.928 2.849 0.02 0.000 0.024 0.109 0.496 0.007 0.102 0.03 15.53

6.123 1.877 0.996 0.145 0.02 2.65 3.026 0.032 0.000 0.000 0.112 0.602 0.01 0.067 0.031 15.592

5.918 2.082 1.276 0.057 0.023 2.726 2.815 0.035 0.000 0.000 0.095 0.611 0.006 0.172 0.022 15.643

5.718 2.282 1.3 0.018 0.000 3.235 2.513 0.015 0.000 0.004 0.049 0.674 0.004 0.202 0.008 15.813

7.098 0.902 1.59 0.006 0.009 2.656 2.293 0.034 0.000 0.001 0.026 0.033 0.025 0.059 0.000 14.674

6.499 1.501 0.838 0.1 0.019 3.124 2.784 0.025 0.000 0.002 0.083 0.485 0.011 0.167 0.000 15.47

6.415 1.585 0.92 0.087 0.019 3.1 2.689 0.041 0.000 0.021 0.084 0.55 0.000 0.093 0.023 15.511

7.326 0.674 0.46 0.008 0.016 3.427 3.141 0.014 0.000 0.000 0.014 0.02 0.002 0.000 0.005 15.102

6.874 1.126 0.362 0.063 0.009 3.289 3.375 0.024 0.000 0.000 0.079 0.218 0.006 0.009 0.008 15.426

6.693 1.307 0.778 0.076 0.013 3.236 2.752 0.03 0.000 0.000 0.094 0.409 0.000 0.212 0.017 15.387

XMg

0.507

0.467

0.492

0.563

0.537

0.529

0.536

0.522

0.494

0.54

A. BENDAOUD ET AL.


Table 3. Continued Spinel Rock type:

Type D

Type E

Type C

Type B

Type A

TD 67N 47

TD 67C 35

TD 38 27

TD 38 28

TD 38 55

Tj 57b 6

Tj 57b 15

Tj 56 48

Tj 56 75

Tj 130 19

TD 128 110

TD 159 10

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 MgO FeO MnO ZnO CaO Na2O K2O Sum

0.02 0.26 58.76 0.93 2.41 7.24 30.47 0.00 n.a. 0.02 0.01 0.00 100.13

0.01 0.1 59.58 0.16 2.17 5.53 33.14 0.12 0.08 0.00 0.00 0.00 100.91

2.79 0.74 29.57 24.86 0.00 1.96 34.94 0.76 n.a. 0.35 0.12 0.00 96.65

0.11 0.3 46.78 13.42 2.24 6.51 27.8 0.55 n.a. 0.21 0.05 0.16 98.28

0.00 0.21 43.98 16.13 0.16 5.99 27.66 0.34 n.a. 0.00 0.07 0.00 95.39

0.00 0.08 60.31 0.08 2.64 8.08 29.00 0.00 0.45 0.03 0.00 0.01 100.67

0.03 0.09 57.9 0.15 3.07 4.78 33.41 0.16 0.38 0.00 0.01 0.00 100.03

0.01 0.14 58.53 0.00 2.42 3.64 35.55 0.1 0.2 0.00 0.05 0.00 100.63

0.04 0.2 58.91 0.06 1.09 3.12 36.64 0.21 0.08 0.01 0.00 0.00 100.38

0.00 0.01 59.28 0.17 1.40 3.80 35.37 0.38 0.01 0.00 0.00 0.00 100.55

0.00 0.13 56.46 0.00 3.61 4.15 33.83 0.33 n.a. 0.00 0.00 0.02 98.53

0.07 0.02 61.67 0.00 0.00 0.04 36.78 0.02 n.a. 0.00 0.00 0.00 98.6

Si Ti Al Cr Fe3þ Mg Fe2þ Mn Zn Ca Na K Sum

0.001 0.005 1.918 0.02 0.05 0.299 0.706 0.000

0.092 0.018 1.145 0.646 0.000 0.096 0.96 0.021

0.003 0.007 1.626 0.313 0.05 0.286 0.686 0.014

0.000 0.005 1.598 0.393 0.004 0.275 0.713 0.009

0.001 0.001 0.000 3.000

0.000 0.002 1.947 0.003 0.045 0.229 0.768 0.003 0.002 0.000 0.000 0.000 3.001

XFe2þ

0.001 0.002 1.926 0.003 0.065 0.201 0.789 0.004 0.008 0.000 0.000 0.000 3.003

0.000 0.003 1.945 0.000 0.051 0.153 0.838 0.002 0.004 0.000 0.003 0.000 3.000

0.001 0.004 1.964 0.001 0.023 0.132 0.867 0.005 0.002 0.000 0.000 0.000 3.001

0.000 0.001 1.964 0.004 0.030 0.159 0.831 0.009 0.002 0.000 0.000 0.000 3.000

0.000 0.003 1.917 0.000 0.078 0.178 0.815 0.008

0.002 0.000 2.076 0.000 0.000 0.002 0.879 0.000

0.000 0.004 0.000 3.081

0.000 0.002 1.941 0.002 0.054 0.329 0.662 0.000 0.009 0.001 0.000 0.000 3.000

0.012 0.008 0.000 3.057

0.007 0.003 0.006 3.016

0.000 0.000 0.001 3.000

0.000 0.000 0.000 2.959

0.7

0.77

0.91

0.71

0.72

0.67

0.8

0.85

0.87

0.839

0.82

1


Sample: Analysis:

(Continued) 127

128

Table 3. Continued Plagioclase Rock type: Sample: Analysis: Position:

Type B

Type C

Type D

Type E

TD 63 45

TD 63 61

Tj 56 74 s

Tj 57b 17 s

Tj 59 56

TD 67C 75

TD 67C 43 /Qtz

TD 67C 26 /Qtz

TD 67C 44 opxcrdspl

TD 38 17 s

TD 38 25 s

TD 38 17 exsol

TD 38 21 matrix

60.37 0.09 26 0.00 0.00 0.02 0.00 0.00 6.55 7.68 0.21 101

61.46 0.00 25.2 0.00 0.00 0.00 0.11 0.03 6.01 7.96 0.37 101.18

56.2 0.02 26.49 0.00 0.00 0.01 0.22 0.07 8.76 7.02 0.04 98.89

45.01 0.01 36.13 0.05 0.00 0.00 0.34 0.03 19 0.48 0.00 101.06

62.34 0.01 25.31 0.03 0.00 0.02 0.12 0.01 5.89 7.81 0.23 101.77

62.16 0.02 24.45 0.00 0.00 0.03 0.05 0.00 5.28 8.79 0.34 101.17

59.84 0.03 25.69 0.09 0.00 0.00 0.39 0.02 6.9 7.52 0.11 100.65

56.76 0.00 28.31 0.06 0.00 0.00 0.41 0.00 9.73 6.05 0.12 101.52

43.78 0.07 35.86 0.00 0.00 0.02 0.37 0.00 19.25 0.76 0.02 100.14

45.86 0.08 34.03 0.00 0.00 0.03 0.44 0.00 17.89 1.12 0.05 99.6

48.72 0.05 32.56 0.16 0.00 0.02 0.48 0.07 15.85 2.21 0.02 100.15

54.97 0.00 28.19 0.04 0.00 0.18 0.82 0.04 10.89 5.1 0.06 100.28

55.06 0.01 28.86 0.00 0.00 0.00 0.37 0.00 11.54 5.11 0.04 101.09

Si Ti Al Cr Fe3þ Mg Fe2þ Mn Ca Na K Sum

2.661 0.003 1.351 0.000 0.000 0.001 0.000 0.000 0.309 0.656 0.012 4.99

2.702 0.000 1.306 0.000 0.000 0.000 0.004 0.001 0.283 0.678 0.021 5.00

2.558 0.001 1.421 0.000 0.000 0.001 0.009 0.003 0.427 0.62 0.002 5.04

2.056 0.000 1.946 0.002 0.000 0.000 0.013 0.001 0.93 0.043 0.000 4.99

2.716 0.000 1.299 0.001 0.000 0.001 0.004 0.001 0.275 0.66 0.013 4.97

2.731 0.001 1.266 0.000 0.000 0.002 0.002 0.000 0.249 0.749 0.019 5.02

2.654 0.001 1.343 0.003 0.000 0.000 0.014 0.001 0.328 0.647 0.006 5.00

2.518 0.000 1.48 0.002 0.000 0.000 0.015 0.000 0.463 0.521 0.007 5.01

2.027 0.002 1.957 0.000 0.000 0.001 0.014 0.000 0.955 0.068 0.001 5.03

2.124 0.003 1.858 0.000 0.000 0.002 0.017 0.000 0.888 0.101 0.003 5.00

2.228 0.002 1.755 0.006 0.000 0.002 0.018 0.003 0.777 0.196 0.001 4.99

2.478 0.000 1.498 0.001 0.000 0.012 0.031 0.001 0.526 0.446 0.003 5.00

2.463 0.000 1.522 0.000 0.000 0.000 0.014 0.000 0.553 0.443 0.002 5.00

Xan Xab Xor

0.32 0.67 0.01

0.29 0.69 0.02

0.41 0.59 0.000

0.956 0.044 0.000

0.29 0.7 0.01

0.245 0.736 0.019

0.334 0.66 0.006

0.467 0.526 0.007

0.933 0.067 0.001

0.895 0.102 0.003

0.798 0.201 0.001

0.539 0.457 0.003

0.554 0.444 0.002

c, core; r, rim; s, symplectite.

A. BENDAOUD ET AL.

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 MgO FeO MnO CaO Na2O K2O Sum

Type A


129

0.3 Pl

d

Sillimanite Free Metapelites Reconstituted primary Opx

x+

Cr

Grt

Op

% Mole 80

Core of primary Opx Rim of primary Opx Secondary Opx

0.2 Alt

70

Gedrite bearing granulites Secondary orthopyroxene bearing Metapelites

0.1

Td 57b

60 50

Opx + Crd

Td 57

Pl

40 XAlm XPy XGrs XSps

30 20

Td 59

0 0

0.2

0.4

0.6

0.8

1

XMg

Fig. 6. Plot of XMg v. Alt (cations p.f.u.) in orthopyroxene of the orthopyroxenebearing þ metapelites.

10 00 0

0.25

0.50

0.75

1.00

1.25 mm

Fig. 5. Compositional profile across garnet in gedrite-bearing granulites.

Orthopyroxene has the same composition in the gedrite-bearing granulites and metapelites with secondary orthopyroxene (types C and D, Table 3): XFe ranges between 0.43 and 0.58 (average of 0.50) and Al2O3 from 1.2 to 4.7 wt%; the most aluminous orthopyroxene is found in the symplectites (both the spinel þ cordierite þ orthopyroxene + plagioclase and the cordierite þ orthopyroxene symplectites) and in the orthopyroxene around quartz. In sillimanite-free orthopyroxene-bearing metapelites (type E, Table 3), the orthopyroxene in the symplectites show a XFe around 0.43 and Al2O3 contents between 2 and 3 wt%. Where exsolution occurs, the primary orthopyroxene is poorer in Fe (XFe 0.34–0.39) and richer in Al2O3 (3.65–5.8 wt%; Fig. 6). Image analysis and microprobe scanning (see analytical techniques above) have allowed us to estimate the composition of the primary orthopyroxene before exsolution: the two methods give similar results with XFe around 0.35 and Al2O3 contents close to 6.5 wt%. Orthoamphibole has a formula based on 23 equivalent oxygen-when calculated according to Spear (1980); this method gives the lowest Fe3þ compatible with stoichiometry, which corresponds to a maximum of Na assigned to the A-site. The composition of the Tidjenouine orthoamphibole is highly variable (XFe 0.43 –0.54, Al2O3 6.58 –21.50 wt%, Na2O 0.07 –2.47 wt%, TiO2 0.02 –1.27 wt%; type D, Table 3, Fig. 7a and b) but always have enough Al to be considered on the gedrite

side of the gedrite– anthophyllite solid solution. The variability in Al2O3 indicates, however, the absence of a miscibility gap, suggesting a temperature of crystallization above 600 8C (Spear 1980). The highest Al2O3 and Na2O values are found in the core of millimetre-sized elongated zoned grains. Orthoamphibole in symplectites have similar compositions to the rims of coarse-grained gedrite. Several substitutions have taken place (Fig. 8): a substitution of edenitic type (Si21 IV (Na,K)A þ2Alþ1; the slope of 0.56 in Figure 8 implies also a compensatory Tschermakitic substiVI tution in reverse (AlIV 21Al21Mgþ1Siþ1). These two substitutions correspond to the pargasitic type substitution (Robinson et al. 1971). An additional titano-Tschermakitic substitution (Siþ2Mgþ1Ti21 AlIV 22) also occurred. These three substitutions imply that AlIV ¼ A-site occupancy þ (AlVI þ Fe3þ þ 2Ti) (Robinson et al. 1971; Czarmanske & Wones 1973); indeed, the substitution of Na in the A-Site and Ti in the octohedral site must be compensated by the substitution of Al for Si in the tetrahedral sites. Spinel composition has a large variability related to the bulk-rock composition: Fe-rich hercynite–spinel solid solution in sillimanite-bearing metapelites (Table 3); Fe-rich (0.82 , XFe , 1) hercynite with nearly no chromite (,0.03%), no Zn and Fe3þ in orthopyroxene-free quartz-bearing metapelites (type A, Table 3) and corundum-bearing metapelites (type B; Table 3); spinel-rich hercynite (no Cr, 0.67 , XFe , 0.79) in metapelites with secondary orthopyroxene and gedrite-bearing rocks (types C and D, Table 3); and ternary solid solution between hercynite, chromite and spinel in sillimanitefree orthopyroxene-bearing metapelites (type E, Table 3). The spinel in the symplectites with orthopyroxene–cordierite–plagioclase–ilmenite–magnetite

130

(a)

A. BENDAOUD ET AL.

1 0.8

(Na + K)A

Na-Gedrite

Oam Ist core Oam Ist rim Oam II

0.6 Ideal Gedrite

0.4 0.2 Gedrite

Anthophyllite

0 0

0.2

0.4

0.6

0.8

1

1.2

1.4

1.6

1.8

2

2.2

2.4

AlIV (b)

1 Magnesio-Gedrite

Magnesio-Anthophyllite 0.8

Gedrite

Anthophyllite

XMg

0.6

0.4

0.2 Ferro-Gedrite

Ferro-Anthophyllite 0

8

7

Si

6

Fig. 7. Plot of orthoamphibole chemical compositions. (a) (Na þ K)A v. AlIV; (b) Si v. XMg, after Leake et al. (1997).

(Herc55Chr39Sp6, XFe ¼ 0.90) is consistently richer in in Mg and Cr (Cr2O3 ¼ 25.73 wt%) than the spinel in contact with quartz (Herc60Chr26Sp14, XFe ¼ 0.70); in both cases, Fe3þ is negligible. Plagioclase is highly variable in composition, but each given rock type and/or microdomain has its own characteristics. Plagioclase has a rather constant composition in the orthopyroxene-free metapelites (type A, Table 3: An25 – 34), with the richest. An composition found in the inclusions in garnet, whereas it has a highly variable composition in the gedritebearing granulites (type D, Table 3): An75 – 92 in the symplectites with orthopyroxene and cordierite; An27 – 47 in contact with quartz at the margin of symplectites, and An17 – 33 when included in quartz and sillimanite. Plagioclase around gedrite is zoned,

showing increasing XAn from the contact with gedrite (An23) towards the periphery (An47). In the secondary-orthopyroxene-bearing metapelites (type C, Table 3), the plagioclase in the leucosome is an unzoned oligoclase (An30) whereas the plagioclase in the spinel–orthopyroxene–cordierite symplectites in fissures in garnet has an almost pure anorthite composition (An95 – 97). Plagioclase from sillimanite-free orthopyroxene-bearing metapelites (type E, Table 3) shows large XAn variation according to the microdomain: between 0.75 and 0.92 in symplectites; 0.50 and 0.58 in plagioclase exsolved by orthopyroxene, and 0.45 , XAn , 0.57 in matrix plagioclase. Alkali-feldspar displays 60–99 mol% of orthoclase component.

TIDJENOUINE METAPELITES EVOLUTION (a)

131

8 Oam Ist core Oam Ist rim Oam II

Si

7

6

5

4 1

0

2

2(Na + K)A +

3

4

AlIV

NaA + NaM4

(b) 0.8 0.6

0.4

0.2

0 0

0.25

0.5

0.75

1

1.25

1.5

Ca + A0 (c) 11

Si + Mg

10

9

8

7

0

1

2

3

4

AlIV + AlIV (d) 19 18

2Si + Mg

17 16 15 14 13 12 0

0.5

1

1.5

2

2.5

Ti + 2AlIV Fig. 8. Orthoamphibole substitutions in gedrite-bearing granulites.

3

3.5

4

4.5

5

132

A. BENDAOUD ET AL.


The other minerals are: ilmenite (Ilm96 – 100 with Mg and Mn , 2–3 mol%); magnetite, present only in sample TD38 as very rare coarse intergrowths with spinel –ilmenite –orthopyroxene– cordierite –plagioclase and is pure Fe3O4; graphite, ubiquitous in metapelites; and pyrite, abundant in secondary-orthopyroxene-bearing metapelites.

Petrological and P– T evolution Several petrogenetical grids are presented. (1) A KFMASH petrogenetic grid involving garnet – orthopyroxene – sillimanite – biotite – melt –K-feldspar –quartz –cordierite –spinel, calculated using Thermocalc 3.1 software (Powell et al. 1998; Fig. 9). Compatibility diagrams were drawn interpret the textures and to work out the theoretical reactions in the KFMASH system. Representative analyses of coexisting phases have been projected from quartz and K-feldspar onto the AFM triangle (Fig. 10). These diagrams show the different stable assemblages derived from textural observations and mineral chemistry as well as the reaction sequences in the quartzbearing metapelites. These diagrams together with the textural relationships in the Tidjenouine metapelites indicate the prograde crossing of the univariant reaction (1), Sill þ Bt þ Qtz ! Grt þ Crd þ Ksp þ Melt, suggested by remnants of biotite, sillimanite and quartz in garnet and cordierite (Fig. 9a). The near metamorphic peak is represented by the crossing of the univariant reaction (8), Grt þ Bt þ Qtz ! Opx þ Crd þ Ksp þ Melt, which is observed in all orthopyroxene-bearing metapelites. During the decompressional stage, the degenerated reaction (11), Grt þ Sill ! Crd þ Spl þ Qtz, occurs. The XFe isopleths of garnet with Qtz and Melt in excess (divariant assemblages: Grt Sil Bt, Grt Crd Bt, Grt Sil Crd and Grt Opx Crd) are also represented in Figure 9b. These isopleths are very P-dependent and constitute a good geobarometer. The core composition of the most magnesian garnet (typical XFe of 0.5), which is observed in orthopyroxene-bearing assemblages, gives a good estimate of the maximum possible pressure, which can be fixed between 7 and 8 kbar. (2) A KFMASH petrogenetic grid involving garnet–corundum–sillimanite–biotite–cordierite– spinel –melt –K-feldpar and water (Fig. 11a). It

133

consists of two invariant points, [H2O] and [Cor], and the univariant reactions that emanate from them (Fig. 11b). The sequence of mineral reactions is well illustrated in Figure 11b. The corundum-consuming reaction (6), Grt þ Cor þ Melt þ Ksp ! Sill þ Spl þ Bt (H2O, Crd)), should occur before the breakdown of biotite and sillimanite with primary garnet to produce a cordierite assemblage (reaction (7), Grt þ Bt þ Sill ! Spl þ Crd þ Melt þ Ksp) (Fig. 11a and b). (3) An FMASH petrogenetic grid involving garnet–orthopyroxene–sillimanite–biotite–gedrite– quartz–cordierite –spinel and water is the same as that constructed by Ouzegane et al. (1996) for aH2O ¼ 1 (Fig. 11a). All reactions at the invariant points are dehydration reactions and therefore lowering aH2O to 0.6 or 0.2, which is in agreement with granulite-facies conditions, should lower the temperature of the invariant points. In this grid, only reactions producing garnet are observed, and the univariant FMASH reaction (13), Oam þ Sill þ Qtz ! Grt þ Crd, is crossed during the prograde stage. A P–T pseudosection has also been constructed for quartz-bearing microdomains (with representative composition: FeO 11.5 mol%, MgO 7 mol%, Al2O3 16 mol%, SiO2 65.5 mol% and aH2O ¼ 1; Fig. 11b). This pseudosection accounts qualitatively for the paragenetic evolution; thus, it shows a very complete history of the P –T evolution of the gedrite-bearing granulites by successive divariant and trivariant assemblages. The occurrence of sillimanite þ gedrite at an early stage of evolution, giving garnet þ gedrite þ sillimanite and garnet þ sillimanite (M1 peak assemblage), implies an increase of temperature before the decompression marked by the growth of cordierite þ orthopyroxene symplectites (M2). Afterwards, the assemblage orthopyroxene þ cordierite þ orthoamphibole (M20 ) indicates a decrease of temperature in the latest stage. This demonstrates that the Tidjenouine rocks have recorded a clockwise P– T evolution. All these stages (M1, M2 and M20 ) most probably occurred during the same metamorphic event. The evolution of pressure and temperature of the Tidjenouine granulite-facies metamorphism has been also determined using internally consistent datasets (average P –T option of Thermocalc, Powell & Holland 1988) and independently calibrated geothermometers and geobarometers. The results are summarized in Table 4. The

Fig. 9. Petrogenetic grid in KFMASH system representing quartz-bearing metapelites calculated using Thermocalc (Powell & Holland 1998). (a) Reactions and preferred P –T path; (b) plot of isopleths of XFe in garnet in different assemblages. Compatibility diagrams are derived from the KFMASH system after projection from quartz, water and K-feldspar (KSH) onto the AFM triangle. Reaction numbers are as in text.

134 A. BENDAOUD ET AL. Fig. 10. Petrogenetic grid in KFMASH system representing quartz-free metapelites, calculated using Thermocalc software (Powell & Holland 1998). (a) Reactions and preferred P– T path; (b) compatibility diagrams derived from the KFMASH system after projection from sillimanite, water, K-feldspar and melt onto the quartz–spinel– hercynite plane. Reaction numbers are as in text.


135

Fig. 11. Petrogenetic grid and P –T pseudosection in FMASH system representing gedrite-bearing metapelites calculated using Thermocalc software (Powell & Holland 1998). (a) Petrogenetic grid showing the displacement of invariant points and univariant reactions depending on aH2O. Reaction numbers are as in text. (b) P –T pseudosection for a fixed bulk composition (mol%: FeO 11.5, MgO 7, Al2O3 16, SiO2 65.5) and aH2O ¼ 1. The P –T path (bold continuous line) takes into account the textural observations in the gedrite-bearing granulites. The temperature is overestimated because of the activity of water which is, in reality, lower than unity.

metamorphic assemblages that we selected for P–T path reconstruction use the M1 peak metamorphism phases, the intermediate paragenesis corresponding to garnet exsolution in Opx, and the M2 decompressional metamorphic reactions between primary minerals. The prograde history is not accessible because of the chemical homogenization of garnet at high temperature. We combined the composition of the cores of the largest garnet grains with those of the cores of the matrix biotite and other primary minerals such as orthopyroxene or plagioclase, to obtain P–T conditions of the M1 peak paragenesis. The M2 retrograde conditions were estimated by using rims of garnet and plagioclase in contact with adjacent secondary biotite, cordierite and orthopyroxene. The M20 cooling stage was estimated by a later biotite product developed at the expense of orthopyroxene. Average P–T calculations were obtained using Thermocalc 3.1 (Powell & Holland 1988; Powell et al. 1998) with the expanded internally consistent dataset of September 1997. Components activities were estimated using the AX program (T.J.B.

Holland, unpublished). Quartz, sillimanite, ilmenite and rutile were assumed to be pure. For each rock, aH2O was chosen, after iterating on aH2O values, on the basis of the best fit (x2 test results; all quoted error estimates are at the 95% confidence level or 2s). The aH2O is additionally constrained by the presence of graphite in several samples. The results of the average P–T calculations are summarized in Table 4.

Peak metamorphism (M1) Sillimanite-free orthopyroxene-bearing and orthopyroxene-free quartz-bearing assemblages allow us to calculate average peak temperature and pressure: sample TD38 (garnet–primary reconstituted orthopyroxene–biotite–plagioclase–K-feldspar–quartz–ilmenite–rutile) gives 7.9 + 1.1 kbar and 863 + 43 8C with an optimum aH2O of 0.3; sample 80–128 (garnet – biotite – plagioclase – K-feldspar – quartz – sillimanite –ilmenite –rutile) gives 7.5 + 1.3 kbar and 855 + 77 8C with an optimum aH2O (in graphite-bearing sample) of 0.7 (the results with an aH2O , 0.4 overlap at 2s uncertainty, as fit values

136

Table 4. Summary of P–T estimates Geothermometers (8C) Grt –Bt PL 83

Exsolution conditions Sil-free metapelites

Grt –Opx Opx –Bt Grt –Crd H 84 S et al. 90 P et al. 85

Grt –Sil –Pl–Qtz NH 81 KN 88

Grt –Opx –Pl– Qtz Grt –Bt –Pl –Qtz NP 82 H 90 (Mg) H 90 (Fe)

7.35 + 0.7 7.8 + 0.6

855*

8.1 + 1.2

850*

6.5 + 1

GRIPS BL 86

Grt –Crd –Sil –Qtz aH2O P & al. 85

T

P

Best fit

7.7 + 0.9

7.2 + 1.2

8 + 0.6

8.7 + 0.5

8.5 + 1.1

8.2 + 0.8 7.2 + 0.4

0.3

863 + 43 7.9 + 1.1

0.95

6.8 + 0.5

0.2

814 + 38 5.5 + 1.1

0.71

0.1

731 + 98 4.3 + 1.2

1.10

0.1

697 + 39 4.3 + 0.5

1.32

7.9 + 1.2 8.5 + 0.6

812 + 25 800 + 33

Thermocalc software Average P –T

6.2 + 1.1

Retrograde conditions Opx-free metapelites 705 + 35 Secondary 610 + 86 740 + 25 660 + 45 Opx-bearing metapelites Silfree metapelites 715 + 32 745 + 30 705 + 40 Ged-bearing 690 + 43 granulites

695 + 23 705 + 60

4 + 1.2 4.8 + 1.1 2.9 + 1.1 3.5 + 0.9

675 + 25 670 + 55

3 + 1.5 3.1 + 1.3

3.1 + 1

4.1 + 1.1 3.4 + 1.2

4.6 + 1.4 4.2 + 1.3 6.25 + 1.3 3+1 2.75 + 0.9 4.1 + 1

5.2 + 1.4

4.8 + 1

6+1 5.5 + 0.8

4.8 + 1.1 5.5 + 1

5.3 + 0.9

PL 83, Perchuck & Lavrent’eva (1983); H 84, Harley (1984); S et al. 90, Sengupta et al. (1990); P et al. 85, Perchuck et al. (1985); NH 81, Newton & Haselton (1981); KN 88, Koziol & Newton (1988); NP 82, Newton & Perkins (1982) H 90, Hoisch (1990); BL 86, Bohlen & Liotta (1986). *With reconstituted orthopyroxene.

A. BENDAOUD ET AL.

Peak conditions Opx-free metapelites 857 + 45 Opx-bearing 827 + 34 metapelites Sil-free metapelites 798 + 25 Ged-bearing granulites

Geobarometers (kbar)


are outside statistical limits; see other calculations in Table 4). Temperatures were calculated for an assumed pressure of 8 kbar, using the garnet – biotite (Perchuck & Lavrent’eva 1983), garnet –orthopyroxene (Harley 1984) and orthopyroxene– biotite (Sengupta et al. 1990) geothermometers. The calculated temperatures are around 857 + 45 8C, 827 + 34 8C and 798 + 25 8C for orthopyroxene-free quartz-bearing metapelites, secondary-orthopyroxene-bearing metapelites and sillimanite-free orthopyroxene-bearing metapelites, respectively, using the calibration of Perchuk & Lavrent’eva (1983). The temperatures calculated using the estimation, by image analysis and microprobe scanning, of the primary orthopyroxene compositions before exsolution are around 865 8C (Grt–Opx: Harley 1984) and 848 8C (Bt –Opx: Sengupta et al. 1990). Pressure estimates were based on the garnet – sillimanite –plagioclase –quartz, garnet –orthopyroxene–plagioclase–quartz, garnet–biotite–plagioclase–quartz and garnet–rutile–ilmenite–plagioclase–quartz assemblages. All these geobarometers give a pressure between 7 and 8.5 kbar. The M1

137

granulite-facies event can thus be set at 800–875 8C and 7–8.5 kbar (Fig. 12a).

Decompressional (M2) and cooling evolution (M20 ) The P–T conditions of the exsolutions in orthopyroxene of sillimanite-free metapelites can be also calculated. Orthopyroxene–garnet–plagioclase– biotite–K-feldspar–quartz–ilmenite–rutile assemblage gives 5.5 + 1.1 kbar and 814 + 38 8C with aH2O ¼ 0.2 (with best results of average P–T of Thermocalc). The later stage is calculated with sample TD38 (sillimanite-free orthopyroxenebearing metapelites) and sample TD57b (secondary orthopyroxene-bearing metapelites). Sample TD38 contains garnet–orthopyroxene–biotite–plagioclase– spinel – cordierite – quartz – K-feldspar – ilmenite – rutile assemblage and gives 4.3 + 0.5 kbar and 697 + 39 8C for an optimum aH2O of 0.1. Sample TD57b is a metapelite in which garnet displays cracks filled with orthopyroxene – spinel – cordierite – plagioclase; this latter assemblage suggests 4.3 + 1.2 kbar and 731 + 98 8C for an

P kbar 14

(b)

Tin Begane

Garnet Pyroxenite

Peak of Metamorphism

12

Metapelite

P kbar

+M rd +C ion Grt Bt t u l > l+ evo z == Sp e + t d r l gra ua Sil Pro te + Q and ==> p ni s a K m illi lt + +S Me r+ tite o o i C B t+ Gr


8 V=

V=

Symplectitic Stage

Ky 4

Sill

Panafrican reheating ?

And

Panafrican reheating ?

And

45

Symplectitic Stage 2 of Garnet Pyroxenite

Amphibolitization

Ky Amphibolitization

42

Symplectitic Stage

6

Sill 4


elt

Tidjenouine

6

Tamanrasset

10

10

8

Symplectitic Stage 1 of Garnet Pyroxenite

Amphibolite (Retrogressed Garnet Pyroxenite)

(a)

2

2 500

600

700

T (°C)

800

900

500

600

700

800

900

T (°C)

Fig. 12. P –T evolution of Tidjenouine metapelites. (a) P –T path; (b) comparison of Tidjenouine metapelites evolution with the metamorphic evolution of Tamanrasset (Ouzegane et al. 2001) and Tin Begane (Derridj et al. 2003).

138

A. BENDAOUD ET AL.

optimum aH2O of 0.1. Results with an aH2O . 0.3 for these two samples show fit values outside statistical limits. Combination of classical geobarometers and geothermometers (Grt–Bt, Grt–Opx, Bt–Opx and Grt–Crd: Table 4) indicates that the later cooling stage occurred at a pressure of 3– 4 kbar and temperatures from 745 to 610 8C (from M2 to M20 ). However, the lowest temperature may correspond to a lower diffusion or to the amphibolitization stage. These results show a good agreement between the P –T conditions obtained from Thermocalc and those obtained from the calibrated geothermometers and geobarometers. The M2 granulitefacies event can thus be estimated at 700 + 50 8C and 3–4 kbar (Fig. 12a).

The late M3 heating metamorphism The M20 amphibolite-facies retrogression is evidenced by the appearance of anthophyllite in the gedrite granulites and of cummingtonite and brown–green hornblende in the metabasic rocks. This stage is followed, along the mega-shear zone, by the crystallization of sillimanite in the metapelites and by the breakdown of amphibole, if quartz is present, to orthopyroxene and plagioclase. The recrystallizations are considered as distinct from M2 and M20 because: (1) sillimanite crosscuts sharply the former mineral orientation; (2) the assemblages indicate a reheating compared with the M20 stage at the amphibolite– granulite transition; (3) in contrast to M2 and M20 , its development is associated spatially with the Pan-African mega-shear zones. This late M3 phase should have occurred at c. 650 –700 8C. This temperature and the association with the mega-shear zones suggests that this phase could be linked with the Pan-African batholiths, whose emplacement is also associated with the mega-shear zones, particularly the Anfeg and Tin Amzi batholiths present in the vicinity of the Tidjenouine granulites (Acef et al. 2003; Fig. 1c).

Zircon U– Pb ages of the Tidjenouine granulites Zircons were hand-picked in alcohol from the least magnetic concentrates (18 tilt at full amperage). Selected crystals were then embedded in epoxy resin, ground and polished to expose the internal structure. They were subsequently observed by back-scattered electron (BSE) imaging using a scanning electron microscope (SEM) at the University of Montpellier II. The sample mounts were later

used for U –Th– Pb microanalyses using a Lambda Physik COMPex 102 excimer laser generating 15 ns pulses of radiation at a wavelength of 193 nm. For analyses, the laser was coupled to a VG Plasmaquad II ICP-MS and analytical procedures followed those outlined by Bruguier et al. (2001) and described in earlier reports (e.g. Neves et al. 2006). Analyses where acquired during two analytical sessions where the spot size of the laser beam was 26 and 51 mm. Unknowns were bracketed by measurements of the G91500 zircon standard (Wiedenbeck et al. 1995), which were used for mass bias and inter-element fractionation corrections. The calculated bias factors and their associated errors were then added in quadrature to the errors measured on each unknown. Accurate common Pb correction during laser ablation analyses is difficult to achieve, mainly because of the isobaric interference of 204Hg with 204Pb. The contribution of 204 Hg to 204Pb was estimated by measuring the 202 Hg and assuming a 202Hg/204Hg natural isotopic composition of 0.2298. This allows monitoring of the common Pb content of the analysed zircon domain, but corrections often resulted in spurious ages. Analyses yielding 204Pb close to or above the limit of detection were thus rejected, and in Table 5 we report only analyses that were found to contain no common Pb. Zircons were separated from the Tidjenouine TJ5 granulitic-facies orthogneiss, a sample with a simple mineralogy comprising quartz, Kfeldspar, plagioclase, biotite, opaque minerals, zircon and apatite. These zircons typically present an internal structure characterized by three concentric zones (Fig. 13): (1) a central zone that is most often grey and homogeneous in BSE but sometimes has a faint oscillatory zoning (e.g. Zr4, Fig. 13); (2) a first rim, brighter in BSE, with a spongy appearance, containing numerous tiny inclusions of calcite; (3) a second rim, not always developed, which is homogeneous and grey in BSE and has no inclusions. Most grains have rounded terminations but still preserve a prismatic shape, suggesting a metamorphic corrosion of originally magmatic grains. In addition, a few grains are not prismatic and display more simple internal structure (Fig. 13, Zr10). The spongy BSE-bright areas are still zircon and the BSE-dark tiny inclusions are calcite. Thus there has not been a destabilization of a pre-existing zircon, but a syncrystallization of zircon and calcite from a melt. This abundance of calcite in these intermediate zones can be correlated to the granulitic-facies metamorphism: (1) fluid inclusions linked to the granulitic decompression stage in the Tamanrasset area are rich in CO2 (Ouzegane et al. 2001); (2) calcite has been

Table 5. U –Th– Pb LA-ICP-MS results for zircon grains from Tidjenouine granulite TJ5 Sample Pb* U Th Th/U (ppm) (ppm) (ppm)

206

Pb/204Pb

208

Pb/206Pb

207

Pb/206Pb +(1s)

207

Pb/235U +(1s)

206

Pb/238U +(1s)

r

Apparent

ages (Ma)

206

Pb/238U +(1s)

159797 163551 38535 174901 138113 129133 62533 104594 95732 60632 326906 220634 506884 333740 400226 319632 512480 599168 280694 474308 418714 249846 388996 470528 413632 513796 493368 235016 30015 435708 365896 631400 460598

0.215 0.218 0.159 0.125 0.217 0.213 0.161 0.134 0.169 0.158 0.148 0.094 0.187 0.194 0.117 0.170 0.222 0.216 0.187 0.190 0.168 0.146 0.183 0.162 0.209 0.157 0.186 0.176 0.076 0.200 0.163 0.226 0.186

0.13076 0.13418 0.13042 0.12763 0.13445 0.13257 0.13323 0.11517 0.12872 0.13232 0.11281 0.11711 0.12378 0.13127 0.12274 0.12940 0.13427 0.13117 0.13436 0.13120 0.13313 0.12741 0.12971 0.13112 0.13150 0.13138 0.12958 0.12503 0.09180 0.13130 0.12718 0.13247 0.12991

0.00045 0.00054 0.00048 0.00044 0.00055 0.00053 0.00054 0.00351 0.00088 0.00089 0.00275 0.00067 0.00085 0.00118 0.00052 0.00111 0.00034 0.00079 0.00026 0.00091 0.00067 0.00034 0.00061 0.00104 0.00083 0.00207 0.00092 0.00061 0.00190 0.00144 0.00096 0.00059 0.00209

6.651 7.000 6.606 5.818 7.260 6.735 7.179 3.727 6.181 6.743 3.086 4.008 4.664 6.735 4.739 6.227 7.258 6.711 7.346 6.731 6.904 5.891 5.931 7.103 6.797 5.176 6.228 4.729 1.595 6.202 5.318 6.238 5.585

0.153 0.160 0.170 0.144 0.163 0.079 0.135 0.271 0.085 0.095 0.155 0.084 0.142 0.156 0.132 0.102 0.134 0.056 0.204 0.109 0.077 0.068 0.195 0.161 0.058 0.171 0.145 0.052 0.101 0.255 0.226 0.088 0.217

0.36888 0.37830 0.36734 0.33058 0.39165 0.36848 0.39077 0.23472 0.34826 0.36960 0.19842 0.24820 0.27324 0.37210 0.28004 0.34900 0.39205 0.37109 0.39653 0.37208 0.37612 0.33532 0.33162 0.39294 0.37486 0.28575 0.34859 0.27429 0.12602 0.34255 0.30325 0.34155 0.31182

0.00838 0.00851 0.00936 0.00811 0.00863 0.00405 0.00716 0.01547 0.00414 0.00457 0.00872 0.00501 0.00812 0.00792 0.00772 0.00490 0.00716 0.00216 0.01101 0.00547 0.00378 0.00377 0.01078 0.00836 0.00214 0.00917 0.00775 0.00267 0.00755 0.01356 0.01270 0.00456 0.01102

0.99 0.98 0.99 0.99 0.98 0.94 0.98 0.91 0.87 0.88 0.87 0.96 0.97 0.92 0.99 0.85 0.99 0.69 1.00 0.90 0.89 0.97 0.99 0.94 0.67 0.97 0.95 0.89 0.94 0.96 0.98 0.95 0.91

2024 2068 2017 1841 2130 2022 2126 1359 1926 2028 1167 1429 1557 2039 1592 1930 2132 2035 2153 2039 2058 1864 1846 2136 2052 1620 1928 1563 765 1899 1707 1894 1750

39 40 44 39 40 19 33 80 20 21 47 26 41 37 39 23 33 10 51 26 18 18 52 39 10 46 37 13 43 65 63 22 54

Pb/206Pb +(1s)

2108 2153 2104 2066 2157 2132 2141 1883 2081 2129 1845 1913 2011 2115 1996 2090 2155 2114 2156 2114 2140 2063 2094 2113 2118 2116 2092 2029 1463 2115 2059 2131 2097

6 7 6 6 7 7 7 55 12 12 44 10 12 16 8 15 4 11 3 12 9 5 8 14 11 28 12 9 39 19 13 8 28

4.0 4.0 4.1 10.9 1.2 5.2 0.7 27.8 7.4 4.8 36.8 25.3 22.6 3.6 20.3 7.7 1.0 3.7 0.1 3.5 3.8 9.6 11.8 21.1 3.1 23.4 7.9 23.0 47.7 10.2 17.1 11.1 16.6

139

(Continued)


Spots on the 2151 Ma discordia li02 140 318 244 0.77 li03 135 306 235 0.77 li04 30 70 40 0.57 li07 143 397 170 0.43 li08 120 251 195 0.78 li10 112 256 190 0.74 li16 55 125 73 0.58 li17 83 326 110 0.34 li18 81 205 125 0.61 li24 52 126 70 0.56 qs02 48 219 77 0.35 qs03 24 93 24 0.26 qs04 85 274 148 0.54 qs05 53 123 76 0.62 qs07 61 200 74 0.37 qs08 45 114 56 0.49 qs09 90 204 158 0.78 qs10 95 220 158 0.72 qs11 55 128 80 0.62 qs12 83 193 132 0.68 qs15 62 148 90 0.61 qs16 47 128 60 0.47 qs17 71 191 117 0.61 qs19 80 181 104 0.58 qs20 76 178 127 0.71 qs22 81 221 115 0.52 qs24 87 223 138 0.62 qs25 39 130 75 0.58 qs27 3 24 1 0.04 qs28 72 181 112 0.62 qs29 59 174 83 0.48 qs31 124 309 238 0.77 qs32 87 243 151 0.62

207

Disc.

140

Table 5. Continued Sample Pb* U Th Th/U (ppm) (ppm) (ppm)

206

Pb/204Pb

208

Pb/206Pb

207

Pb/206Pb +(1s)

207

Pb/235U +(1s)

206

Pb/238U +(1s)

r

Apparent 206

Pb/

qs33 qs34 qs35

41 64 122

98 274 295

42 97 199

ages (Ma)

238

U +(1s)

207

206

Pb/

Disc.

Pb +(1s)

217364 336960 560518

0.118 0.164 0.192

0.13169 0.12129 0.13123

0.00121 0.00131 0.00052

6.881 3.652 6.910

0.136 0.074 0.072

0.37902 0.00666 0.89 0.21839 0.00377 0.85 0.38190 0.00373 0.93

2072 1273 2085

31 20 18

2121 1975 2115

16 19 7

2.3 35.5 1.4

Spots on the 2062 Ma discordia li01 45 314 42 0.13 li05 51 230 55 0.24 li06 88 743 30 0.04 li09 36 135 38 0.28 li11 79 643 17 0.03 li12 85 666 19 0.03 li19 31 130 51 0.39 qs06 91 221 145 0.66 qs14 8 50 4 0.08 qs18 73 204 72 0.35 qs21 61 199 83 0.42 qs23 68 181 32 0.18 qs30 72 193 86 0.44

65431 78852 126845 47865 111001 115340 37315 521178 83814 441820 369426 450526 468496

0.047 0.058 0.016 0.083 0.015 0.024 0.149 0.203 0.069 0.113 0.111 0.076 0.131

0.07952 0.10386 0.07834 0.11461 0.08021 0.08773 0.10696 0.12592 0.08833 0.12675 0.11877 0.12713 0.12402

0.00069 0.00222 0.00057 0.00179 0.00038 0.00042 0.00230 0.00097 0.00145 0.00059 0.00068 0.00093 0.00136

1.683 3.179 1.326 4.124 1.425 1.686 3.289 6.222 2.027 5.953 4.679 6.437 5.843

0.185 0.186 0.022 0.202 0.028 0.058 0.179 0.062 0.038 0.091 0.117 0.106 0.227

0.15346 0.22198 0.12274 0.26097 0.12883 0.13941 0.22299 0.35837 0.16645 0.34065 0.28573 0.36724 0.34169

0.01682 0.01206 0.00178 0.01209 0.00250 0.00472 0.01112 0.00229 0.00146 0.00496 0.00693 0.00542 0.01274

1.00 0.93 0.89 0.95 0.97 0.99 0.92 0.64 0.47 0.95 0.97 0.90 0.96

920 1292 746 1495 781 841 1298 1974 993 1890 1620 2016 1895

93 63 10 62 14 27 58 11 8 24 35 25 61

1185 1694 1155 1874 1202 1377 1748 2042 1390 2053 1938 2059 2015

17 39 14 28 9 9 39 14 32 8 10 13 19

22.3 23.7 35.4 20.2 35.0 38.9 25.8 3.3 28.6 8.0 16.4 2.1 6.0

Spots on the concordia at 614 Ma li15 10 105 1 0.01 li13 9 102 1 0.01 li21 8 89 1 0.02 li14 11 120 2 0.01 qs13 6 65 1 0.01 qs1 5 55 1 0.01

13132 12645 11241 15700 41346 34388

0.006 0.006 0.006 0.005 0.011 0.018

0.06081 0.06032 0.06066 0.06004 0.06146 0.06179

0.00086 0.00058 0.00048 0.00093 0.00182 0.00048

0.848 0.836 0.835 0.826 0.831 0.840

0.052 0.018 0.017 0.044 0.030 0.018

0.10108 0.10052 0.09982 0.09982 0.09801 0.09865

0.00598 0.00190 0.00190 0.00506 0.00194 0.00197

0.97 0.89 0.92 0.96 0.56 0.93

621 617 613 613 603 606

35 11 11 30 11 12

633 615 627 605 655 667

31 21 17 34 64 17

1.9 20.4 2.2 21.4 8.0 9.0

A. BENDAOUD ET AL.

0.43 0.35 0.68


141

Fig. 13. Texture of the dated Tidjenouine zircon using SEM (back-scattered electrons). White circles indicate the location of spot analyses. Ages indicated are the discordia or concordia ages shown Figure 14. ‘% disc.’ gives the degree of discordance of the considered spot. ‘Spongy’ areas are made of zircon with tiny inclusions of apatite. In crystal Zr6, there is one spot on a central grey zone with an age of 2062 Ma: this is attributed to the presence of the spongy zone present very close to the spot just below the analysed surface. This is just visible on close inspection of the picture.

142

A. BENDAOUD ET AL.

Fig. 14. Zircon U–Pb concordia diagrams showing concordia and discordia ages: the grey ellipses correspond to the zircon grey central zones, the hatched ellipses to the zircon ‘spongy’ zones; within the inset, the ellipses correspond to single zircons not displaying the corona texture of most of the Tidjenouine zircons (grey: used in the calculation; white: not used (for calculation including that spot, see text); black: result of the concordia age calculation).

described as a granulitic metamorphic phase in the same area (Ouzegane 1981); (3) the presence of Ca-rich minerals (Ca-plagioclase, apatite) in melanosome in the Tidjenouine granulite-facies migmatite suggests that Ca was in excess during the granulitic migmatitization. This means that these inclusion-rich zones should be related to the granulite-facies migmatitic event. Sixty spots have been analysed on these zircons. They show a broad alignment from c. 2100 Ma to c. 600 Ma. When considering these results and the relation between ages and the different zircon domains (Fig. 14), the following patterns arise. (1) Thirty-four spots in central grey zones define a discordia line with an upper intercept of 2144 + 9 Ma and a lower intercept of 597 + 27 Ma (2s, MSWD ¼ 1.5). Among these analyses, five concordant spots provide a slightly older but consistent age of 2151 + 8 Ma (2s, five zircons, MSWD ¼ 1.5). We consider this last age as the best estimate for the crystallization of these

central zones. Th/U ratios of this group vary between 0.78 and 0.43 for spots with 206Pb/238U ages above 1700 Ma, those with younger 206 Pb/238U ages having ratios between 0.58 and 0.26. (2) Thirteen spots in spongy intermediate zones define a discordia line with an upper intercept of 2062 + 39 Ma and a lower intercept of 681 + 63 Ma (2s, MSWD ¼ 4.1); there are no true concordant spots in this group but four spots have only a slight discordance below 8%: their mean 207Pb/206Pb age is 2049 + 22 Ma; Th/U ratios of this group vary between 0.66 and 0.24 for spots with 206Pb/238U ages above 1200 Ma, those with 206Pb/238U ages below 1000 Ma having ratios between 0.03 and 0.13. We note that the U and Pb concentrations in these analyses are not significantly different from those of the first group (Table 5), indicating that the calcite inclusions present in these zones do not interfere in these analyses, as we would expect.


(3) Five spots obtained in the non-prismatic core-free zircons are concordant close to the previous discordia lower intercepts and a sixth one is nearly concordant. Our best estimate for this batch of analyses is 614 + 11 Ma (five zircons, MSWD ¼ 0.71). Their Th/U ratios are very low, between 0.01 and 0.02. The outer rims displayed by some zircons were too thin to be analysed by the laser ablation technique but we propose the hypothesis that a similar Pan-African age would have been acquired on these zones. Finally two spots are slightly below the two discordias and have not been included in the age calculations. The oldest age of 2151 + 8 Ma has been determined on central parts of the grains, some of which are zoned and characteristic of a magmatic crystallization. This age is thus attributed to the magmatic protolith of the granulite. The slightly younger age of 2062 + 39 Ma is questionable, as it has been calculated from discordant analyses sampling the intermediate coronas linked to the granulitic migmatitic event (M1 and M2 phase). The location of these data points on the left of the c. 2.15 – 0.60 Ga discordia line indicates that these zones have undergone U – Pb disturbances, at some times in the past, between these two ages. The limited degree of discordance of some of these analyses (,10%) is taken as evidence for a Palaeoproterozoic age for this event. This would imply that both the prograde M1 and retrograde M2 metamorphic phases are Eburnean in age and most probably correspond to one metamorphic path. A younger age (i.e. Neoproterozoic) for the granulitic event cannot be strictly ruled out in the absence of concordant analyses but is unlikely: in this case, the spots acquired on the intermediate zones should lie on a discordia line pointing to c. 2.15 Ga and not as much to the left. The rare independent crystals unzoned and unaffected by the reaction coronas are dated at 614 + 11 Ma, an age that can probably be applied to the thin external rims of most zircons. This age corresponds to that of the intrusion of the neighbouring granitic batholiths such as the Anfeg batholith (608 + 7 Ma; U – Pb zircon, Bertrand et al. 1986, recalculated by Lie´geois et al. 2003) and thus to the M3 thermal metamorphic phase, which is thus Pan-African in age. The fact that this phase was the most effective in lowering the Th/U ratio indicates that during the Pan-African M3 metamorphism, only solid-state reactions occurred, whereas melts were produced during the Eburnean M1 – M2 granulite-facies migmatitic event, the lowering of the Th/U ratio being favoured by metamorphic fluids (Williams et al. 1996), which probably eased the exchange of Th between zircon and minerals such as monazite.

143

Discussion and conclusion In several areas of the Laouni terrane, observed granulitic formations are commonly associated with an important migmatitic event. The textural relationships and the P– T estimates suggest that the beginning and maintenance of melt production occurred during the prograde metamorphic evolution (M1) culminating at 850 8C and 7.5 kbar. A large part of the retrograde evolution (M2) down to 700 8C and 4 kbar, also occurred under granulitefacies conditions: the presence of early, strongly restitic granulites (corundum metapelites earlier than the garnet– sillimanite– biotite metamorphic peak) indicates that migmatitization was already important before the M1 climax and some melt was also produced during the late breakdown of biotite (M2). The M2 stage evolved eventually to an M20 phase in the amphibolite facies at 600 8C, which is evidenced by some late minerals such as anthophyllite, secondary biotite and cummingtonite, depending on the rock type. This granulitic metamorphism is Eburnean (2062 + 39 Ma). This clockwise retrograde P–T segment is similar to that constructed using a variety of different rock types (metapelitic and metabasic rocks) from the basement of the Laouni terrane (Ouzegane et al. 2001; Bendaoud et al. 2003; Derridj et al. 2003). During this evolution aH2O generally decreased, probably because of absorption of H2O in anatectic melts, preserving most of the granulite-facies parageneses (M20 is local). Our petrological and thermobarometric study indicates a clockwise P–T path marked by a decompression stage generating spectacular coronitic and symplectitic textures in both the para- and ortho-derived metamorphic units. The succession of parageneses during this decompression depends on the chemical composition of the rocks. In Tidjenouine, the metapelites and the microdomains rich in Si and Mg are characterized by the appearance of an orthopyroxene–cordierite association at the expense of garnet, quartz and biotite, in the absence of sillimanite. On the other hand, the metapelites and the microdomains rich in Al and Fe display the spinel –cordierite assemblage, without orthopyroxene, following the destabilization of garnet, sillimanite and biotite. The occurrence of sillimanite inclusions in the core of primary garnet in quartz-bearing metapelites confirms that this mineral was present during the prograde stage. The peak pressures obtained at Tamanrasset (10 kbar: Ouzegane et al. 2001) and at Tin Begane (12 kbar: Derridj et al. 2003) are higher than those obtained in the study area (7–8 kbar). This can be related to different exposed crustal levels (Bendaoud et al. 2004). Coupled with the observation of the abundance of often subhorizontal

144

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shear zones, this suggests that the LATEA microcontinent is composed of a series of Eburnean nappes, probably resulting from a collisional orogeny. It is thus possible that the shear zones interpreted as Pan-African in age (Bertrand et al. 1986) were initiated during the Eburnean orogeny and reactivated during the Pan-African orogeny. More work is needed to assess this hypothesis. The age of the protolith of the dated sample (2151 + 8 Ma) is thus probably related to a precollisional event such as a subduction regime. No Archaean age is recorded here as in the other regions of the southern LATEA (Bertrand et al. 1986; Barbey et al. 1989); Archaean ages are currently only known in the Gour Oumelalen region (NE LATEA; Peucat et al. 2003; Fig. 1). This could suggest the existence of an Archaean continent to the NE involved to the SW in a collisional orogeny with a Palaeoproterozoic terrane, but more geochronological, metamorphic and geochemical data are needed to proceed in this interpretation. We can point that the Eburnean granulitic metamorphism in the Archaean Gour Oumelalen area is younger (c. 1900 Ma; Peucat et al. 2003) than in SW LATEA (c. 2100 Ma; Barbey et al. 1989; this study). The geodynamic understanding of the Eburnean evolution of Hoggar is still in its infancy. The age of 614 + 11 Ma obtained on single unzoned zircons and the large discordance of many Eburnean zircons indicate that the effect of the Pan-African orogeny was important in LATEA although the Eburnean granulite-facies parageneses are well preserved. Similar ages have been obtained on the Telohat migmatites (609 + 17 Ma; U– Pb zircon lower intercept; Barbey et al. 1989). The Pan-African event is marked by the M3 thermal metamorphism (650 8C; 3–4 kbar) that led to the destabilization of the amphibole in the metabasic rocks and probably of the biotite in the metapelites, and allowed the crystallization of a new generation of sillimanite not linked to the M1 –M2 metamorphic phase, as postulated by Caby (2003). The M3 metamorphism is synchronous with the large Pan-African batholiths such as the Anfeg batholith (608 + 7 Ma: U –Pb zircon, Bertrand et al. 1986, recalculated by Lie´geois et al. 2003); these in turn are synchronous with the development of the large shear zones characteristic of the Tuareg shield (Fig. 1). These batholiths are rooted in the subvertical major shear zones and were emplaced as sheets along reactivated pre-existing subhorizontal shear zones (Acef et al. 2003; Lie´geois et al. 2003). We can here confirm the pre-existence of these subhorizontal shear zones, to which we attribute an initial Eburnean age on the basis of the above petrological results linked to

the dated c. 2060 Ma granulitic-facies metamorphism. True dating of these shear zones remains to be done. These findings shed light on the LATEA Pan-African metacratonic evolution (Lie´geois et al. 2003): the LATEA microcontinent was mainly built during the Eburnean orogeny, which generated a regional granulite-facies metamorphism, and became a craton by lithospheric thickening (Black & Lie´geois 1993) during the Mesoproterozoic, a quiet period for LATEA (no Mesoproterozoic events are recorded in central Hoggar) as for most of West Africa. This rigid cratonic behaviour allowed LATEA to become amalgamated with several Neoproterozoic island arcs (Lie´geois et al. 2003): the Iskel terrane at 870–850 Ma (Caby et al. 1982), and the Tin Begane unit at c. 685 Ma (Lie´geois et al. 2003) among others, which are not yet dated. These accretion events are not recorded in the Tidjenouine granulites. The main Pan-African orogenic phase is characterized by large horizontal movements along mega-shear zones and the intrusion of granitoid batholiths in the 620–580 Ma age range (Bertrand et al. 1986; Caby & Andreopoulos-Renaud 1989; Black et al. 1994; Lie´geois et al. 1994, 2003). This phase dismembered the LATEA craton and heat transfer was caused by the magmas rising along the shear zones, although many of the cratonic features were preserved, including the Eburnean granulitic paragenesis and probably many Eburnean structures, although they were slightly to strongly reworked. This corresponds to the notion of metacraton (Abdelsalam et al. 2002) that can be applied to LATEA (Lie´geois et al. 2003). Taking into account the relatively small area of LATEA, we can suggest that it belonged, before the Pan-African orogeny, to a larger craton probably constituting its margin. Whether LATEA represents the former eastern boundary of the West African craton or the western boundary of the Saharan craton is still a matter of debate. The Tidjenouine area demonstrates the complexity of metacratonic areas that result from the interplay of two orogenies on a rigid block. This is the reason why metacratonic areas are most often not well understood and are probably now among the most fascinating regions to study with modern techniques.

We warmly thank G. Rebay and P. Goncalves for their reviews, which significantly improved the final version of the manuscript. Lively discussions with R. Caby on the Eburnean v. Pan-African effects in Hoggar were appreciated. We thank N. Ennih for his editorial comments. This work was supported by the TASSILI 05

TIDJENOUINE METAPELITES EVOLUTION MDU 653 project ‘Imagerie tridimentionnelle et e´volution spatio-temporelle du Hoggar’ and by the NATO grant EST/CLE 979766 and CNRS PICS project ‘Architecture lithosphe´rique et dynamique du manteau sous le Hoggar’. We are also extremely grateful to ORGM and OPNA for logistic support during fieldwork.

References A BDELSALAM , M., L IE´ GEOIS , J. P. & S TERN , R. J. 2002. The Saharan metacraton. Journal of African Earth Science, 34, 119 –136. A CEF , K., L IE´ GEOIS , J. P., O UABADI , A. & L ATOUCHE , L. 2003. The Anfeg post-collisional Pan-African high-K calc-alkaline batholith (Central Hoggar, Algeria), result of the LATEA microcontinent metacratonisation. Journal of African Earth Sciences, 37, 95– 311. A DJERID , Z., O UZEGANE , K., G ODARD , G. & K IENAST , J. R. 2008. First report of ultrahigh- temperature sapphirine spinel quartz and orthopyroxene þ spinel þ quartz parageneses discovered in Al–Mg granulites from the Khanfous area (In Ouzzal metacraton, Hoggar, Algeria). In: E NNIH , N. & L IE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton Geological Society, London, Special Publications, 297, 147–167. B ARBEY , P., B ERTRAND , J. M., A NGOUA , S. & D AUTEL , D. 1989. Petrology and U/Pb geochronology of the Telohat migmatites, Aleksod, Central Hoggar, Algeria. Contributions to Mineralogy and Petrology, 101, 207–219. B ENDAOUD , A., O UZEGANE , K. & K IENAST , J. R. 2003. Textures and phase relationships in ferrous granulites from Tidjenouine (Hoggar, Algeria): fayalite– ferrossilite– quartz secondary assemblage. Journal of African Earth Sciences, 37, 241–255. B ENDAOUD , A., D ERRIDJ , A., O UZEGANE , K. & K IENAST , J. R. 2004. Granulites of the Laouni terrane, (Tamanrasset, Tidjenouine, Tin Begane). Journal of African Earth Sciences, 39, 187– 192. B ENYAHIA , O., H ADDOUM , H., O UZEGANE , K., B ENDAOUD , A., D JEMAI , S. & K IENAST , J.-R. 2005. Fonctionnement et roˆle des me´ga zones de cisaillement dans la structuration du me´tacraton e´burneén du LATEA au panafricain puis au phane´rozoı¨que (re´gion de Tamanrasset, Hoggar, Alge´rie). African Geosciences Review, 12, 261– 274. B ERTRAND , J.-M. & J ARDIM DE S A´ , E. F. 1990. Where are the Eburnean–Transamazonian collisional belts? Canadian Journal of Earth Sciences, 27, 1382–1393. B ERTRAND , J. M., M ICHARD , A., B OULLIER , A. M. & D AUTEL , D. 1986. Structure and U/Pb geochronology of Central Hoggar (Algeria): a reappraisal of its Pan-African evolution. Tectonics, 5, 955 –972. B LACK , R. & L IE´ GEOIS , J. P. 1993. Cratons, mobile belts, alkaline rocks and continental lithospheric mantle: the Pan-African testimony. Journal of the Geological Society, London, 150, 89–98. B LACK , R., L ATOUCHE , L., L IE´ GEOIS , J. P., C ABY , R. & B ERTRAND , J. M. 1994. Pan-African displaced terranes in the Tuareg shield (central Sahara). Geology, 22, 641–644.

145

B OHLEN , S. R. & L IOTTA , J. J. 1986. A barometer for garnet amphibolites and granulites. Journal of Petrology, 27, 1025–1034. B RUGUIER , O., T ELOUK , P., C OCHERIE , A., F OUILLAC , A. M. & A LBARE` DE , F. 2001. Evaluation of Pb–Pb and U– Pb laser ablation ICP-MS zircon dating using matrix-matched calibration samples with a frequency quadrupled (266 nm) Nd:YAG laser. Geostandards Newsletter, 25, 361– 373. C ABY , R. 2003. Terrane assembly and geodynamic evolution of central– western Hoggar: a synthesis. Journal of African Earth Sciences, 37, 133 –159. C ABY , R. & A NDREOPOULOS -R ENAUD , U. 1989. Age U– Pb a` 620 Ma d’un pluton synorogeńique de l’Adrar des Iforas (Mali). conse´quences pour l’aˆge de la phase majeure de l’oroge`ne pan-africaine. Comptes Rendus de l’Acade´mie des Sciences, 308, 307– 314. R., A NDREOPOULOS -R ENAUD , U. & C ABY , G RAVELLE , M. 1982. Cadre ge´ ologique et ge´ ochronologique U/Pb sur zircon des batholites pre´ coces dans le segment pan-africain du Hoggar central (Alge´ rie). Bulletin de la Socie´ te´ Ge´ ologique de France, 24, 677 – 684. C ZAMANSKE , G. D. & W ONES , D. R. 1973. Oxidation during magmatic differentiation, Finnmarka Complex, Oslo area, Norway. II. The mafic silicates. Journal of Petrology, 14, 349–380. D ERRIDJ , A., O UZEGANE , K., K IENAST , J. R. & B ELHAI¨ , D. 2003. P–T –X evolution in garnet pyroxenites from Tin Begane (Central Hoggar, Algeria). Journal of African Earth Sciences, 37, 257–268. H ANSON , G. N. 1989. An approach to trace element modeling using a simple igneous system as an example. In: LIPIN , R. B. & MC KAY , G. A. (eds) The Geology and Geochemistry of Rare Earth Elements. Mineralogical Society of America, Reviews in Mineralogy, 21, 79–97. H ARLEY , S. L. 1984. An experimental study of the partitioning of Fe and Mg between garnet and orthopyroxene. Contributions to Mineralogy and Petrology, 86, 359– 373. H ARLEY , S. L. 1989. The origins of granulites: a metamorphic perspective. Geological Magazine, 126, 215– 247. H OISCH , T. D. 1990. Empirical calibration of six geobarometers for the mineral assemblage quartz þ muscovite þ biotite þ plagioclase þ garnet. Contributions to Mineralogy and Petrology, 104, 225–234. H OLLAND , T. J. B. & P OWELL , R. 1990. An enlarged and updated internally consistent thermodynamic dataset with uncertainties and correlations: the system K2O– Na2O–CaO–MgO–FeO–Fe2O3 –Al2O3 –TiO2 –SiO2 – C–H–O. Journal of Metamorphic Geology, 8, 89–124. K OZIOL , A. M. & N EWTON , R. C. 1988. Redetermination of the anorthite breakdown reaction and improvement of the plagioclase –garnet– Al2SiO5 –quartz barometer. American Mineralogist, 73, 216– 223. L EAKE , B. E., W OOLLEY , A. R. & B IRCH , W. D. et al. 1997. Nomenclature of amphiboles. Canadian Mineralogist, 9, 623– 651. L IE´ GEOIS , J. P., B LACK , R., N AVEZ , J. & L ATOUCHE , L. 1994. Early and late Pan-African orogenies in the Aı¨r assembly of terranes (Tuareg shield, Niger). Precambrian Research, 67, 59– 88.

146

A. BENDAOUD ET AL.

L IE´ GEOIS , J. P., L ATOUCHE , L., B OUGHRARA , M., N AVEZ , J. & G UIRAUD , M. 2003. The LATEA metacraton (Central Hoggar, Tuareg shield, Algeria): behaviour of an old passive margin during the Pan-African orogeny. Journal of African Earth Sciences, 37, 161– 190. N EVES , S., B RUGUIER , O., V AUCHEZ , A., B OSCH , D., R ANGEL DA S ILVA , J. M. & M ARIANO , G. 2006. Timing of crust formation, deposition of supracrustal sequences, and Transamazonian and Brasiliano metamorphism in the East Pernambuco belt (central domain, Borborema Province, NE Brazil): implications for western Gondwana assembly. Precambrian Research, 149, 197–216. N EWTON , R. C. & H ASELTON , H. T. 1981. Thermodynamics of the garnet–plagioclase– Al2SiO5 –quartz geobarometer. In: N EWTON , R. C., N AVROTSKY , A. & W OOD , B. J. (eds) Thermodynamics of Minerals and Melts. Springer-Verlag, New York, 131–147. N EWTON , R. C. & P ERKINS , D. I. 1982. Thermodynamic calibration of geobarometers based on assemblages garnet–plagioclase–orthopyroxene (clinopyroxene)– quartz. American Mineralogist, 67, 203– 222. O UZEGANE , K. 1981. Le me´tamorphisme polyphase´ granulitique de la region de Tamanrasset (Hoggar central). These´ de 3e´me cycle, Universite´, Paris VII, France. O UZEGANE , K. & B OUMAZA , S. 1996. An example of ultrahigh-temperature metamorphism: orthopyroxene –sillimanite– garnet, sapphirine–quartz and spinel-quartz parageneses in Al–Mg granulites from In Hihaou, In Ouzzal, Hoggar. Journal of Metamorphic Geology, 14, 693–708. O UZEGANE , K., D JEMAI , S. & G UIRAUD , M. 1996. Gedrite garnet sillimanite bearing granulites from Amesmessa area, south In Ouzzal, Hoggar, Algeria. Journal of Metamorphic Geology, 14, 739–753. O UZEGANE , K., B ENDAOUD , A., K IENAST , J. R. & T OURET , J. L. R. 2001. Pressure–temperature–fluid evolution in the Eburnean metabasites and metapelites from Tamanrasset (Hoggar, Algeria). Journal of Geology, 109, 247– 263. O UZEGANE , K., K IENAST , J. R., B ENDAOUD , A. & D RARENI , A. 2003. A review of Archaean and Paleoproterozoic evolution of the In Ouzzal granulitic terrane (Western Hoggar, Algeria). Journal of African Earth Sciences, 37, 207 –227. P ERCHUK , L. L. & L AVRENT ’ EVA , I. V. 1983. Experimental investigation of exchange equilibria in the system cordierite–garnet– biotite. In: S AXENA , S. K. (ed.) Kinetics and Equilibrium in Mineral Reactions. Advances in Physical Geochemistry, 3, 199– 239.

P ERCHUK , L. L., A RANOVICH , L. Y. & P ODLESSKII , K. K. et al. 1985. Precambrian granulites of the Aldan shield, eastern Siberia, USSR. Journal of Metamorphic Geology, 3, 265– 310. P EUCAT , J. J., D RARENI , A., L ATOUCHE , L., D ELOULE , E. & V IDAL , P. 2003. U–Pb zircon (TIMS and SIMS) and Sm– Nd whole-rock geochronology of the Gour Oumelalen granulitic basement, Hoggar massif, Tuareg shield, Algeria. Journal of African Earth Sciences, 37, 229–239. P OWELL , R. & H OLLAND , T. J. B. 1988. An internally consistent data set with uncertainties and correlations: 3. Applications to geobarometry worked examples and a computer program. Journal of Metamorphic Geology, 6, 173–204. P OWELL , R., H OLLAND , T. J. B. & W ORLEY , B. 1998. Calculating phase diagrams involving solid solutions via nonlinear equations, with examples using THERMOCALC. Journal of Metamorphic Geology, 16, 577–588. R OBINSON , P., R OSS , M. & J AFFE , H. W. 1971. Composition of the anthophyllite –gedrite series: comparisons of gedrite–hornblende and the anthophyllite– gedrite solvus. American Mineralogist, 56, 1005– 1041. S ENGUPTA , P., D ASGUPTA , S., B HATTACHARYA , P. K. & M UKHERJEE , M. 1990. An orthopyroxene –biotite geothermometer and its application in crustal granulites and mantle-derived rocks. Journal of Metamorphic Geology, 8, 191– 197. S PEAR , F. S. 1980. The gedrite– anthoplyllite solvus and the composition limits of orthoamphibole from the Post Pond volcanics, Vermont. American Mineralogist, 65, 1103–1118. S UN , S. S. & M C D ONOUGH , W. F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: S AUNDERS , A. D. & N ORRY , M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications, 42, 313 –345. T AYLOR , S. R. & M C L ENNAN , S. M. 1985. The Continental Crust: its Composition and Evolution. Blackwell, Oxford. W IEDENBECK , M., A LLE´ , P., C ORFU , F., G RIFFIN , W. L. & M EIER , M. 1995. Three natural zircon standards for U–Th– Pb, Lu– Hf, trace element and REE analyses. Geostandards Newsletter, 19, 1– 23. W ILLIAM , I. S., B UICK , I. S. & C ARTWRIGHT , I. 1996. An extended episode of early Mesoproterozoic metamorphic fluid flow in the Reynold Range, central Australia. Journal of Metamorphic Geology, 14, 29–47.

First report of ultrahigh-temperature sapphirine 1 spinel 1 quartz and orthopyroxene 1 spinel 1 quartz parageneses discovered in Al –Mg granulites from the Khanfous area (In Ouzzal metacraton, Hoggar, Algeria) Z. ADJERID1, K. OUZEGANE2, G. GODARD3 & J. R. KIENAST4 1

Ecole Normale Supe´rieure, B.P. 92, Vieux Kouba, 16500 Alger, Alge´rie (e-mail: [email protected])

2

Laboratoire de Geódynamique, Geólogie de l’Ingeńieur et de Plane´tologie, FSTGAT –USTHB, B.P. 32 El Alia, Dar el Beida, 16111 Alger, Alge´rie

3

Equipe Geóbiosphe`re actuelle et primitive, CNRS IPGP, Universite´ Denis-Diderot (Paris 7), case 89, 4 place Jussieu, 75252 Paris, France 4

Laboratoire de Geósciences marines, IPGP, universitie´ Denis-Diderot (Paris 7), case 89, 4 place Jussieu, 75252 Paris, France Abstract: The Archaean to Palaeoproterozoic Khanfous area from the Archaean In Ouzzal granulite terrane (Western Hoggar, Algeria) preserves exceptional thermal-peak (1150 , T , 1300 8C) mineral parageneses, consisting of orthopyroxene þ spinel þ quartz, sapphirine þ spinel þ quartz and sapphirine þ orthopyroxene þ quartz, in quartz-rich Al–Mg granulites. Reaction textures coupled with P –T FMASH pseudosections indicate that rocks experienced complex multi-stage evolution. Our results suggest that the Khanfous area, as well as the entire northern In Ouzzal metacraton, experienced ultrahigh-temperature crustal metamorphism attributed to a 2 Ga Palaeoproterozoic event, followed by exhumation along a clockwise P –T path. The extreme temperatures attained suggest delamination of the lithosphere and ascent of the asthenosphere after crustal thickening.

Extreme crustal metamorphism at temperatures of 950 –1100 8C is mainly restricted to Proterozoic and Archaean granulite-facies terranes. Al –Mgrich granulites formed under these ultrahightemperature (UHT) conditions are characterized by the presence of orthopyroxene þ sillimanite + quartz. Other indicators of UHT metamorphism are the thermal-peak parageneses spinel þ quartz (Waters 1991) and sapphirine þ quartz (e.g. Dallwitz 1968; Hensen 1971; Hensen & Green 1973; Harley 1985, 1998), assemblages with osumilite (e.g. Ellis 1980; Grew 1982; Audibert et al. 1995; Carrington & Harley 1995; Sajeev & Osanai 2004), high-Al orthopyroxene (8– 12 wt% Al2O3: Harley & Motoyoshi 2000; Harley 2004) and/or ternary feldspar preserved as mesoperthite or antiperthite in metapelites (e.g. Harley 1985; Sandiford 1985; Sheraton et al. 1980, 1987). Thermodynamics can predict the stability at even higher temperatures (.1100 8C) of assemblages never previously observed in nature, such as sapphirine þ spinel þ quartz and orthopyroxene þ spinel þ quartz. We describe here the first occurrence of these parageneses, which were

found in some Al– Mg granulites from the In Ouzzal metacraton (northwestern Hoggar, Algeria) together with sapphirine þ orthopyroxene þ quartz and spinel þ quartz parageneses previously reported in the same area by Bertrand et al. (1992) and Ouzegane & Boumaza (1996). After a section devoted to the geological setting, we present the main assemblages (petrography) and their minerals (mineralogy). Metamorphic conditions and P –T paths based on P –T and T– X pseudosections are then reported. Implications for the evolution of the In Ouzzal metacraton are discussed in the conclusions.

Geological setting The In Ouzzal metacraton (Hoggar, Fig. 1a and 1b) is a well-known example of a deep Archaean crust that experienced ultrahigh-temperature (UHT) metamorphism (peak T . 1050 8C at 10 kbar) during a Palaeoproterozoic tectonometamorphic event (Kienast & Ouzegane 1987; Bertrand et al. 1992; Mouri et al. 1993, 1994;


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Fig. 1. (a) Geological map of the Hoggar shield, after Black et al. (1994). (b) Geological and sketch map of the northern part of the In Ouzzal terrane. (c) Geological map of the Khanfous area, with location of investigated samples (see (b) for location).

Haddoum et al. 1994; Guiraud et al. 1996; Ouzegane & Boumaza 1996; Ouzegane & Kienast 1996; Ouzegane et al. 2003a, b, and references therein). It consists of high-grade orthogneisses,

tholeiitic– komatiitic basic or ultrabasic lenses and metasediments, including marbles, magnetitebearing quartzites and Al –Mg granulites (e.g. Ouzegane et al. 2003a).

ULTRAHIGH-I MINERALS, IN OUZZAL

Geochronological data suggest that the oldest rocks of the In Ouzzal are 3.3 –3.2 Ga enderbites (U/Pb on zircon: Peucat et al. 1996; Ouzegane et al. 2003a). The sedimentary series is thought to have deposited between 2.65 and 2.70 Ga. The youngest Archaean igneous event, at 2.5 Ga, involved calc-alkaline granites generated by partial melting of precursors (tonalites and interbedded metasediments) from the lower to middle continental crust (Peucat et al. 1996). All these rocks underwent UHT metamorphism at 2.0 Ga (e.g. Peucat et al. 1996). In contrast, the PanAfrican orogeny had a negligible impact on the inner In Ouzzal metacraton. The studied rocks occur on the northern foothill of Jebel Khanfous (228530 4100 N, 028480 1000 E), in the central Tekhamalt region (In Ouzzal; Fig. 1c). The Khanfous complex is a 700 m wide structure extending 2 km in an east–west direction. It consists of a variety of rocks arranged in two main series. (1) The first series consists of alternating bands of high-grade metasediments dominated by quartzite, Al– Mg granulites, lenses of marble and magnetite-bearing quartzites (Fig. 1c). The series locally comprises mafic and ultramafic rocks. (2) The second series consists of a clinopyroxenebearing orthogneiss, which forms the top of the hill more than 60 m above the ‘reg’. The alkaline protolith, dated at 2.65 Ga, ranges in composition from granodiorite–monzogranite to granite (69– 75 wt% SiO2: Peucat et al. 1996). It is now metamorphosed as a granulite-facies orthogneiss characterized by the paragenesis perthitic K-feldspar þ quartz þ ferro-augite + oligoclase + green hornblende. The accessory minerals are apatite, zircon, magnetite, ilmenite, pyrochlore and REE-rich chevkinite (Drareni et al. 2007).

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The studied rocks are Al –Mg granulites, which represent the dominant rock type of the metasedimentary series. They are dense, massive rocks characterized by a subvertical foliation parallel to the lithological layering. The stretching lineation, marked by the preferred orientation of orthopyroxene and sillimanite, shows a highly variable but generally steep plunge (70 –908). These granulites form horizons 15 cm to several metres in thickness with a russet-red patina. Centimetre-thick lenses of a quartz-free, sapphirine-rich granulite with a particular dark blue colour are also present in places.

Petrology According to the classification of Bernard-Griffiths et al. (1996) and Ouzegane et al. (2003a), the studied samples are quartzitic Al –Mg granulites (Tek96: 85 . SiO2 . 65 wt%) and quartz-bearing Al –Mg granulites (Tek58, Tek100, Tek102: 65 . SiO2 . 45 wt %). These two types have the same mineral assemblage but different silica contents. They contain abundant quartz (45–60 vol% Qtz), garnet (10 –19% Grt), and orthopyroxene (8– 14% Opx). Sillimanite (Sil), cordierite (Crd), sapphirine (Spr), and spinel (Spl) are present in varying amounts. Accessory phases include biotite (Bt , 2%), K-feldspar (Kfs , 1%), plagioclase (Pl , 1%), ilmenite (Ilm), rutile (Rt), zircon, monazite, apatite and pyrite. The main characteristic of our samples is their peculiar composition (Table 1): SiO2, Al2O3, MgO and FeO account for 98 wt%, with low amounts of CaO and alkalis (K2O . Na2O). Bernard-Griffiths et al. (1996) and Fourcade et al.

Table 1. Element data for some whole-rock samples Sample: Al –Mg type:

Tek96 Quartzitic granulite

Tek101 Quartzitic granulite

Tek102 Quartz-bearing granulite

SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 Cr2O3 (ppm) NiO (ppm) Sum XMg

77.20 7.22 7.24 0.09 7.01 0.13 0.00 0.25 0.28 0.02 429 130 99.43 0.69

64.80 12.00 12.50 0.10 8.13 0.41 0.00 1.22 0.61 0.04 1090 234 99.81 0.60

56.80 20.10 13.00 0.22 6.09 0.91 0.32 1.32 1.17 0.03 994 230 99.86 0.52

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(1996) interpreted their composition, particularly the d18O ratio and high Ni and Cr contents, in terms of a sedimentary mixture of a granitic source and a hydrothermally altered basic to ultrabasic material. These rocks show complex textures as a result of the extremely heterogeneous mineral associations that developed in closely spaced domains during the various metamorphic stages. Textural relationships indicate the succession of four main metamorphic stages, which are described in detail below. (1) Stage I: the mineral assemblages defining this stage are preserved unevenly and consist of quartz þ garnet þ biotite. (2) Stage II: orthopyroxene–sillimanite-bearing symplectites around garnet developed during this stage. (3) Stage III: this stage is characterized by the development of peculiar UHT parageneses such as sapphirine þ spinel þ quartz and orthopyroxene þ spinel þ quartz. (4) Stage IV: cordierite appeared widely after peak metamorphism. All the cracks in garnet were filled by very fine, late cordierite-bearing symplectites.

Stage I The medium- to fine-grained assemblages that define this stage include quartz, garnet and biotite, which are systematically separated by orthopyroxene–sillimanite symplectites (i.e. stage II). Garnet occurs as both small inclusion-free anhedral grains (,0.5 mm) and large subidioblastic crystals (up to 3 cm in diameter). The latter often contain inclusions such as quartz, Fe–Ti oxides, biotite and sillimanite, which may be ascribed to an earlier stage. The presence, among these inclusions, of some scarce sapphirine, sillimanite and spinel suggests that the early paragenesis fossilized in garnet formed in high-temperature, low-pressure conditions.

Stage II The most spectacular symplectites that developed during this stage are observed around garnet crystals and along their cracks. The resorption of garnet occurred through reactions (1) and (2). At the contact with quartz, garnet grains are surrounded by fine- to coarse-grained symplectites of orthopyroxene þ sillimanite + plagioclase (Fig. 2a). Relics of garnet and quartz are preserved in some of the orthopyroxene and sillimanite vermicules (Fig. 2a). Symplectites in garnet are unevenly distributed, suggesting the former

Fig. 2. Backscattered scanning electron (SEM) images. (a) Stage II large Opx þ Sil symplectites between Grt þ Qtz (reaction (1)). The development of Crd þ Opx symplectites at the interface between quartz and garnet occurred later (reaction (8)). (b) The development of Opx þ Sil symplectites in garnet crystals suggests the former presence of quartz (reaction (1)). (c) Reactions (1) and (2) occur in close microdomains.


presence of quartz (Fig. 2b). These features indicate the reaction Grt þ Qtz , Opx þ Sil þ R (1) 1 (Fe, Mg, Mn)3:000 Si3:002 Al1:990 Ca0:037 O12 þ 0:700 SiO2 , 1:600 (Fe, Mg, Mn)1:864 Al0:297 Si1:851 Ca0:001 O6 þ 0:760 Al2 SiO5 þ 1 (Fe, Mg, Mn)0:014 Si0:032 Al0:005 Ca0:035 : The reaction was balanced in the CaMnFMAS (CaO– Na2O –FeO –MgO –Al2O3 –SiO2) system, neglecting minor components (K2O, TiO2, Cr2O3 . . .). The vector R represents the residue of the least-squares method. The compositions used for the phases are means of several punctual analyses, obtained by electron microprobe. In silica-deficient microdomains, this stage is marked by the development of sapphirine, orthopyroxene and sillimanite around garnet. This symplectite developed through the reaction Grt , Spr þ Opx þ Sil þ R 1 (Fe, Mg, Mn)3:002 Al1:980 Si3:020 Ca0:032 O12

(2)

, 0:211 (Fe, Mg, Mn)1:770 Al4:490 Si0:740 O10 þ 1:495 (Fe, Mg)1:751 Al0:399 Si1:749 O6 þ 0:233 Al2 SiO5 þ 1 (Fe, Mg, Mn)0:001 Si0:003 Al0:009 Ca0:032 : Figure 2c shows both reactions (1) and (2) in close microdomains.

Stage III Characteristic UHT assemblages comprising sapphirine and/or spinel with quartz and Al-rich orthopyroxene frequently developed during this stage (Fig. 3). They formed at the expense of the previous stage I and stage II associations consisting of garnet þ orthopyroxene þ sillimanite þ quartz. The occurrence of the peculiar sapphirine– spinel –quartz paragenesis indicates that the UHT peak was attained during this stage. At these high temperatures, the stage II orthopyroxene–sillimanite assemblage is replaced by a sapphirine –quartz symplectite, with sapphirine developing mainly around sillimanite crystals (Fig. 3a). This feature suggests the reaction

151

This exceptional UHT association is also observed between garnet and sillimanite or in garnet fractures (Fig. 3b), where sillimanite is completely consumed and replaced by a coarse sapphirine –quartz symplectite through the multivariant reaction Grt þ Sil , Spr þ Qtz þ R

(4)

1 (Fe, Mg, Mn)2:880 Al1:998 Si3:001 Ca0:045 O12 þ 2:230 Al2 SiO5 , 1:520 (Fe, Mg, Mn)1:876 Al4:250 Si0:830 O10 þ 3:960 SiO2 þ 1(Fe, Mg, Mn)0:001 Si0:009 Al0:002 Ca0:045 : Another reaction, reaction (5), between sillimanite and garnet produced a second exceptional assemblage, namely spinel þ quartz. This assemblage is only observed in the garnet fractures of sample Tek96 (Fig. 3c), together with strongly zoned plagioclase coronas that mantle garnet crystals, recycling Ca from garnet (Fig. 3c): Grt þ Sil , Spl þ Qtz þ Pl þ R 1 (Fe, Mg, Mn)2:952 Al2:019 Si2:966 Ca0:057 O12

(5)

þ 1:493 Al2 SiO5 , 2:728 (Fe, Mg, Mn)1:020 Al1:798 Si0:000 Cr0:150 Zn0:033 O10 þ 4:298 SiO2 þ 0:062Al1:638 Si2:351 Na0:359 Ca0:659 O8 þ 1 (Fe, Mg, Mn)0:061 Si0:000 Al0:000 Ca0:014 Na0:026 Cr0:396 Zn0:089 :

Reactions (4) and (5) are similar, as both have garnet and sillimanite as reactants. Their combination led to the formation of the unique assemblage sapphirine þ quartz þ spinel. As sapphirine, quartz and spinel are almost collinear (i.e. spinel þ quartz sapphirine), one of the two reactions (reaction (4) or (5)) can easily operate in place of the other depending on the microdomain. They can even interchange through time, explaining how sapphirine grew at the spinel – quartz interface in some symplectites (Fig. 3c) and how spinel grew from sapphirine in other microdomains (Fig. 3d). The occurrence of sapOpx þ Sil , Spr þ Qtz þ R (3) phirine, spinel and quartz in mutual contact (Fig. 3c and d), described here for the first time, 1(Fe, Mg, Mn)1:820 Al0:430 Si1:742 Ca0:002 Cr0:003 O6 provides further evidence of UHT metamorphism þ 1:853Al2 SiO5 (T . 1170 8C), as discussed below. In sillimanite-free microdomains, sapphirine– , 0:970(Fe, Mg)1:860 Al4:287 Si0:810 Cr0:052 O10 quartz– orthopyroxene UHT parageneses deveþ 2:809SiO2 loped mostly around garnet porphyroblasts and þ 1(Fe,Mg,Mn)0:001 Si0:000 Al0:022 Ca0:002 Cr0:047 : along their fractures (Fig. 3e). In some cases

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Fig. 3. Images of UHT symplectites. (a) A complex texture representing two progressive reactions: the Opx–Sil association, crystallized between Grt þ Qtz during stage II (reaction (1)), is replaced by a sapphirine þ quartz symplectite (reaction (3)) during stage III. (b) A UHT Spr þ Qtz symplectite developed at the expense of garnet and sillimanite, now completely consumed (reaction (4)). (c) The development of Spl þ Qtz þ Spr at the interface between


orthopyroxene is concentrated around garnet, whereas the sapphirine and quartz grew preferentially in the cracks (Fig. 3e). This peculiar texture indicates that the symplectite developed through the reaction Grt , Opx þ Spr þ Qtz þ R 1 (Fe, Mg, Mn)2:962 Si2:970 Al2:000 Ca0:038 O12 , 1:210 (Fe, Mg, Mn)1:843 Al0:297 Si1:840

(6)

Na0:002 Ca0:027 O6 þ 0:385(Fe, Mg, Mn)1:860 Si0:780 Al4:286 O10 þ 0:443 SiO2 þ 1(Fe, Mg, Mn)0:007 Si0:000 Al0:009 Ca0:045 Mn0:022 : Another unusual UHT paragenesis that formed during this stage consists mostly of spinel, quartz and orthopyroxene (Fig. 3f). This exceptional association, which to our knowledge has never been reported before, is observed around garnet crystals of the most ferrous rocks (XMg 0.50). This reaction texture was partially obliterated by the development of late Crd-bearing coronas during retrogression (Fig. 3f): Grt , Opx þ Spl þ Qtz þ R

(7)

1 (Fe, Mg, Mn)2:943 Si3:028 Al1:980 Ca0:029 O12 , 0:106 (Fe, Mg, Mn)6:970 Al14:670 Si0:000 Cr0:978 O32 þ 1:203 (Fe, Mg, Mn)1:800 Si1:810 Al0:350 O6 þ 0:852 SiO2 þ 1 (Fe, Mg, Mn)0:006 Si0:001 Al0:004 Ca0:029 Cr0:104 :

Stage IV The main characteristic of this stage is the widespread development of cordierite, which grew as thin films at the interface between a pre-existing Al-rich phase (sillimanite, sapphirine or spinel) and orthopyroxene, garnet or quartz (Figs 3d,f,

153

and 4a,b). Cordierite, together with orthopyroxene, sapphirine or spinel, also formed symplectites, mainly at the expense of garnet, quartz and sillimanite (Fig. 4a and c). As the size of the vermicular symplectite crystals is temperaturedependent, the coarse symplectites are thought to have developed at relatively high temperatures, whereas fine intergrowths appeared after substantial cooling. Among these symplectites, we observe cordierite – sapphirine intergrowths and cordierite coronas between relict sillimanite and orthopyroxene ([Sil] j Spr þ Crd j Crd j [Opx]): Opx þ Sil , Crd þ Spr þ R 1 (Mg, Fe)1:820 Al0:380 Si1:790 O6 þ 1:720 Al2 SiO5 , 0:660 (Mg, Fe)1:979 Al4:027 Si4:989 O18 þ 0:270 (Fe, Mg)1:890 Al4:280 Si0:803 O10 þ 1 (Fe, Mg)0:004 Si0:000 Al0:006 :

(8)

The most common cordierite-bearing symplectite occurs between garnet and quartz, forming with orthopyroxene and plagioclase a complex corona texture ([Qtz] j Opx j Pl j Opx þ Crd j [Grt]; Fig. 4a). These textural relationships suggest the general equilibrium reaction Grt þ Qtz , Crd þ Opx þ Pl þ R

(9)

1 (Fe, Mg, Mn)2:970 Si2:990 Al1:99 Ca0:054 Na0:0186 O12 þ 1:350 SiO2 , 0:420 (Fe, Mg, Mn)1:96 Al3:95 Si5:01 Na0:042 Ca0:046 O18 þ 1:13 (Fe, Mg, Mn)1:870 Al0:230 Si1:870 Na0:001 Ca0:003 O6 þ 0:050 Si2:560 Al1:440 Ca0:420 Na0:580 O8 þ 1 (Fe, Mg, Mn)0:004 Si0:005 Al0:01 Ca0:010 Na0:016 : Where garnet is isolated from quartz, a very fine intergrowth cordierite þ spinel þ orthopyroxene + plagioclase developed principally

Fig. 3. (Continued) garnet and sillimanite, which has completely disappeared, suggests the combination of reactions (4) and (5) during stage III. In this case, sapphirine is thought to have developed between spinel and quartz. It should be noted that the An content of plagioclase increases adjacent to garnet (red, Al; green, Mg; blue, Ca). (d) The UHT sapphirine þ spinel þ quartz paragenesis (stage III); spinel is thought to have grown at the expense of sapphirine. The triple junction, indicating equilibrium between these phases, should be noted. (e) The UHT Spr þ Opx þ Qtz assemblage crystallized around garnet crystals and along fractures in garnet (reaction (6)) during stage III. (f) The stability of the peculiar UHT Spl þ Qtz þ Opx paragenesis that probably developed through reaction (7). It should be noted that quartz is sometimes in direct contact with spinel and orthopyroxene. The appearance of cordierite as thin films around all phases is related to stage IV.

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Fig. 4. Backscattered scanning electron image showing the main reaction textures that developed during stage IV. (a) Coronal cordierite (Crd) separating sillimanite (Sil) from garnet (Grt) and quartz (Qtz) (upper part), and the reaction between quartz (Qtz) and garnet (Grt) yielding cordierite (Crd), orthopyroxene (Opx) and plagioclase (Pl) according to reaction (9) (lower part). It should be noted that orthopyroxene preferentially mantles quartz, whereas Crd þ Opx þ Pl symplectites developed around garnet. (b) Details of a Crd þ Spr þ Opx symplectite that developed at the margin of a partially resorbed garnet. (c) The development of Crd þ Opx þ Spl along garnet cracks (reaction (10)).


along fractures in garnet (Fig. 4c). This suggests the reaction Grt , Crd þ Opx þ Spl þ R (10) 1 (Fe, Mg, Mn)3:966 Si2:990 Al1:986 Ca0:029 O12 , 0:187 (Fe, Mg, Mn)1:965 Si4:980 Al4:039 O18 þ 1:065 (Fe, Mg, Mn)1:939 Al0:149 Si1:939 O6 þ 0:575 (Fe, Mg, Mn)0:949 Si0:002 Al1:857 Cr0:086 Zn0:047 O4 þ 1 (Fe, Mg, Mn)0:008 Si0:002 Al0:002 Ca0:029 Cr0:550 Zn0:380: : All the reaction textures involving cordierite formed during retrogression, as discussed below.

Mineralogy Representative analyses of the main mineral phases are reported in Tables 2– 7. Chemical analyses were performed using a Cameca SX50 electron microprobe at the University of Paris. The operating conditions were 15 kV accelerating voltage and 10 nA sample current. Natural silicates and synthetic oxides were used as standards for all elements except fluorine and zinc; the latter were calibrated on fluorite and sphalerite, respectively.

Garnet The garnet solid solution is dominated by pyrope– almandine (Prp36 – 61 –Alm59 – 37) (Table 2). Andradite and uvarovite contents are low, showing no significant variations within single crystals, whereas spessartine and grossular contents are more variable. The main chemical variations are in Fe and Mg contents, with the XMg (¼Mg/ (Fe2þ þ Mg)) value in all samples varying widely from 0.36 to 0.61. Garnet is chemically heterogeneous at the sample scale, as a result of a compositional layering of the rock. At the microdomain scale, this variation controlled the type of reaction in which garnet was involved (Fig. 5a). Reaction (1) occurred only in the presence of garnet with an XMg value of 0.49– 0.56. Reaction (2) was likewise linked to Mg-rich garnet (0.53 , XMg , 0.58). The scarcity of the assemblage sapphirine þ orthopyroxene þ quartz (reaction (6)) is related to the rarity of garnet with the highest pyrope-content (XMg . 0.56). The growth of spinel in association with cordierite and orthopyroxene (reaction (10)) implies almandine-rich compositions (0.43 , The most Fe-rich garnet XMg , 0.50).

155

(0.38 , XMg , 0.43) broke down in contact with sillimanite and quartz to form cordierite coronas. Garnet crystals generally show a plateau-like zoning profile, displaying strong chemical variations only at rims or along fractures (Fig. 5b). The plateau composition is probably related to equilibrium at peak metamorphism, whereas the rims re-equilibrated during retrogression, with a decrease in the XMg value especially at the contact with biotite (Fig. 5b).

Orthopyroxene Orthopyroxene presents three modes of occurrence: (1) as coarse-grained porphyroblasts (5–20 mm), generally strewn with small rutile needles parallel to the {110} planes and thought to result from exsolution during retrogression (Perchuk et al. 1985); (2) as coarse UHT symplectites together with sillimanite (reaction (1)), sillimanite and sapphirine (reaction (2)) sapphirine and quartz (reaction (6)), or spinel and quartz (reaction (7)); (3) as late cordierite-bearing coronas and symplectites between garnet and quartz (reaction (9)) or after garnet (reaction (10)). In all cases, orthopyroxene is bronzite (0.73 , XMg , 0.78). In contrast to late orthopyroxene, early coarse-grained porphyroblasts contain as much as 11.4 wt% Al2O3 (i.e. 23 mol% of the Tschermak end-member; Table 3); that is, values equal to the highest ones reported for the Al–Mg granulites of the In Ouzzal (Bertrand et al. 1992). Such high contents are found in similar rocks of Antarctica, where the stability of the sapphirine– quartz association has also been reported (c. 12 wt%; Harley & Motoyoshi 2000). At the rim of porphyroblasts, the decrease in the Al content to 16 mol% marks the drop in temperature during retrogression. In the UHT symplectites, the Tschermak component ranges from 23 to 17 mol%, whereas it is generally ,8 mol% in the late cordierite-bearing symplectites (Table 3).

Cordierite The composition of cordierite approaches that of the Mg end-member (Table 4). Although its composition in all Crd-bearing assemblages is homogeneous at the thin-section scale, XMg varies slightly from sample to sample in the 0.86–0.92 range. In all cases, the K, Ca, Mn and Ti contents do not usually exceed 0.5 p.f.u. Cordierite grains may contain minor amounts of CO2 and H2O, because their analytical total is 1– 2% short of 100 wt%; this suggests the presence of channelfilling volatiles, although the H2O content was not estimated.

156

Table 2. Representative chemical analyses of garnet Stage: Sample no.: Reaction:

Stage III

Stage IV

Tek96 (3) 1

Tek96 (1) 2

Tek96 (1) 4 and 5

Tek96 (3) 6

Tek96 (3) 6

Tek96 (3) 9

Tek96 (2) 10

Tek96 (1) 10

40.67 0.01 23.54 0.00 20.58 0.17 14.62 0.60 0.00 100.19

39.88 0.03 22.92 0.19 21.56 0.33 15.08 0.67 0.00 100.68

41.20 0.03 24.77 0.29 21.73 0.42 12.46 1.35 0.00 102.25

39.97 0.01 23.01 0.33 20.22 0.53 15.55 0.39 0.00 100.01

40.13 0.07 22.89 0.35 20.32 0.32 15.77 0.23 0.00 100.08

40.67 0.01 23.54 0.05 20.58 0.17 14.62 0.60 0.00 100.24

39.73 0.01 21.97 0.05 26.02 0.51 11.23 0.35 0.00 99.87

39.91 0.04 21.82 0.27 23.71 0.43 12.77 0.69 0.01 99.67

3.028 0.000 1.972 0.000 0.000 0.003 1.658 1.277 0.033 0.029 0.000 54.869 0.000 0.816 43.046 1.122 0.147 0.43

3.017 0.000 1.942 0.000 0.002 0.016 1.498 1.439 0.027 0.056 0.002 48.245 0.009 1.046 48.931 0.933 0.837 0.49

3.007 0.000 2.050 0.000 0.001 0.000 1.273 1.612 0.011 0.048 0.000 43.251 0.000 1.616 54.771 0.362 0.000 0.56

2.939 0.061 1.928 0.113 0.002 0.011 1.216 1.657 0.021 0.053 0.000 33.518 1.414 0.000 63.646 0.798 0.624 0.58

3.019 0.000 2.137 0.000 0.002 0.017 1.331 1.361 0.026 0.106 0.000 47.129 0.000 2.847 48.18 0.928 0.917 0.51

2.953 0.047 1.954 0.066 0.001 0.019 1.183 1.712 0.033 0.031 0.000 32.105 0.075 0.000 65.448 1.267 1.105 0.59

2.961 0.039 1.949 0.056 0.004 0.020 1.198 1.734 0.020 0.018 0.000 31.658 0.701 0.000 66.870 0.771 0.000 0.59

3.007 0.000 2.050 0.000 0.001 0.000 1.273 1.612 0.011 0.048 0.000 43.251 0.000 1.616 54.771 0.362 0.000 0.56

Z. ADJERID ET AL.

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O Sum Cations per 12 oxygens TSi AlIV AlVI Fe3þ Ti Cr Fe2þ Mg Mn Ca Na Alm Adr Grs Prp Sps Uv XMg

Stage II


157

Fig. 5. (a) Plot of garnet compositions in the almandine (Alm) þ spessartine (Sps)– grossular (Grs)–pyrope (Prp) diagram. This composition is controlled at the microdomain scale by the type of reaction in which garnet is involved. (b) Profile of electron microprobe data across the rim of a compositionally zoned garnet grain in contact with biotite. Garnet is more Fe-rich when in direct contact with biotite, whereas the core composition is Mg-rich. The latter is thought to be acquired during prograde metamorphism. The Fe content also increases at the contact with fine spinel-bearing symplectites in cracks. The equidistance between points is 1 mm.

158

Table 3. Representative chemical analyses of orthopyroxene Coarse-grained porphyroblasts

Late fine symplectites

UHT coarse symplectites

Sample no.: Tek58 (1) Tek58 (1) Tek58 (1) Tek58 (1) Tek58 (1) Tek58 (1) Tek96 (1) Tek96 (1) Tek96 (1) Tek96 (2) Tek96 (3) Tek96 (2) Tek96 (2) Reaction Reaction Reaction Reaction Reation Reaction Reaction Type: Core Core Core Rim Rim Rim (4) (4) (9) (9) (9) (10) (10) 48.31 11.48 0.06 0.81 15.00 0.05 24.01 0.02 0.03 0.00 99.77 1.741 0.258 0.229 0.488 0.002 0.023 0.005 0.446 1.289 0.001 0.001 0.002 0.000 4 0.74 0.26 0.74 0.23

48.91 11.34 0.01 0.07 15.50 0.11 24.33 0.10 0.00 0.00 100.37 1.750 0.249 0.229 0.478 0.000 0.001 0.018 0.445 1.297 0.003 0.003 0.000 0.000 4 0.74 0.25 0.74 0.23

49.35 9.08 0.22 0.04 17.32 0.12 23.43 0.03 0.01 0.01 99.61 1.798 0.202 0.187 0.389 0.006 0.001 0.002 0.525 1.272 0.004 0.001 0.000 0.000 4 70.48 29.44 0.71 0.19

49.4 8.5 0.20 0.06 18.23 0.13 23.48 0.02 0.00 0.00 100.05 1.796 0.204 0.160 0.364 0.006 0.002 0.030 0.523 1.273 0.004 0.001 0.000 0.000 4 69.48 30.47 0.71 0.16

50.55 8.45 0.12 0.13 16.62 0.06 24.86 0.01 0.02 0.00 100.83 1.809 0.191 0.166 0.357 0.003 0.004 0.017 0.480 1.326 0.002 0.000 0.002 0.000 4 72.63 27.34 0.73 0.17

48.11 10.62 0.11 0.33 15.50 0.01 25.24 0.07 0.00 0.00 99.99 1.720 0.276 0.172 0.448 0.002 0.009 0.088 0.375 1.347 0.000 0.002 0.000 0.000 4 0.78 0.22 0.78 0.17

49.74 10.53 0.28 0.00 14.41 0.00 24.81 0.03 0.03 0.00 99.83 1.784 0.215 0.230 0.445 0.007 0.000 0.000 0.432 1.326 0.000 0.001 0.002 0.000 4 0.75 0.24 0.75 0.23

52.61 2.73 0.02 0.19 17.33 0.05 27.18 0.05 0.02 0.00 100.18 1.89 0.106 0.010 0.116 0.000 0.005 0.091 0.430 1.458 0.001 0.002 0.001 0.000 4 73.55 26.34 0.77 0.01

52.69 4.16 0.05 0.31 15.74 0.11 26.65 0.05 0.04 0.00 99.8 1.899 0.101 0.076 0.177 0.001 0.009 0.016 0.458 1.432 0.003 0.002 0.003 0.000 4 74.91 24.99 0.76 0.08

51.66 5.17 0.08 0.24 15.37 0.15 27.3 0.09 0.01 0.00 100.07 1.849 0.151 0.067 0.218 0.002 0.008 0.074 0.385 1.456 0.004 0.003 0.001 0.000 4 75.71 24.11 0.79 0.07

54.81 1.61 0.09 0.33 16.60 0.18 26.23 0.09 0.01 0.00 99.95 1.988 0.011 0.056 0.068 0.002 0.009 0.000 0.503 1.417 0.005 0.003 0.001 0.000 4 0.72 0.26 0.74 0.06

54.50 1.87 0.06 0.28 16.79 0.04 26.41 0.07 0.01 0.00 100.03 1.972 0.027 0.052 0.079 0.001 0.008 0.000 0.508 1.424 0.001 0.002 0.001 0.000 4 0.73 0.26 0.74 0.052

Z. ADJERID ET AL.

SiO2 47.98 Al2O3 11.05 0.15 TiO2 Cr2O3 0.31 FeOt 15.77 MnO 0.09 MgO 24.81 CaO 0.06 Na2O 0.03 K2O 0.00 Sum 100.25 Cations per 6 oxygens Si 1.717 0.282 AlIV VI 0.184 Al Altot 0.466 Ti 0.004 Cr 0.008 3þ Fe 0.083 2þ Fe 0.388 Mg 1.323 Mn 0.003 Ca 0.002 Na 0.002 K 0.000 Sum 4 En 0.77 Fs 0.22 Xmg 0.77 Xmgts 0.18


159

Table 4. Representative chemical analyses of cordierite Stage: Sample no.: Reaction:

Stage IV Tek96 (1) Reaction (9)

50.25 SiO2 TiO2 0.00 Al2O3 34.59 FeO 1.96 MnO 0.00 MgO 12.37 CaO 0.00 0.00 Na2O 0.00 K2O Sum 99.17 Cations per 18 oxygens Si 4.977 Al 4.035 Ti 0.000 2þ Fe 0.162 Mn 0.000 Mg 1.827 Ca 0.000 Na 0.000 K 0.000 Sum 11.00 0.92 XMg

Tek96 (1) Reaction (9)

Tek58 (2) Corona

Tek58 (2) Corona



49.54 0.00 33.88 2.37 0.00 12.34 0.06 0.07 0.01 98.27

49.53 0.02 34.30 3.18 0.12 11.48 0.00 0.02 0.00 98.65

49.87 0.00 34.11 3.41 0.08 11.47 0.03 0.01 0.00 98.99

49.21 0.01 33.90 3.53 0.00 11.47 0.00 0.02 0.05 98.19

48.80 0.07 33.56 3.06 0.00 11.14 0.00 0.01 0.02 96.66

4.962 3.996 0.000 0.199 0.000 1.843 0.006 0.014 0.001 11.02 0.90

4.961 4.046 0.002 0.267 0.011 1.714 0.000 0.004 0.000 11.00 0.87

4.984 4.015 0.000 0.285 0.007 1.709 0.004 0.003 0.000 11.01 0.86

4.957 4.022 0.001 0.298 0.000 1.722 0.000 0.003 0.007 11.01 0.85

4.992 4.042 0.006 0.262 0.000 1.699 0.000 0.002 0.003 11.01 0.87

Sapphirine Sapphirine occurs in all samples and is compositionally homogeneous within individual grains. It is Mg-rich (XMg ¼ 0.79 2 0.87) and close to the 7:9:3 stoichiometry (i.e. 7[Mg, Fe]O:9[Al, Cr, Fe3þ]:3[SiO2]; Higgins et al. 1979) (Table 5). The Al content ranges from 60 to 64 wt%. This variable composition is related to the different textural settings, as UHT sapphirine intergrown with quartz or spinel is less aluminous and less magnesian (Al2O3 60 wt%; XMg 0.82) than sapphirine grains developed within cordierite –orthopyroxene symplectites (Al2O3 63 wt%; XMg 0.86). The Fe3þ/Fetot stoichiometric ratio (see Higgins et al. 1979) reaches 0.27 and depends on the mineral assemblage. UHT sapphirine crystallized with quartz is less ferric (Fe3þ/Fetot 0.14) than the fine-grained sapphirine intergrown with cordierite and/or orthopyroxene (Fe3þ/Fetot ¼ 0.17–0.27).

Spinel Spinel is present in various symplectites (reactions (5), (7) and (10); see stages III and IV of the petrological section), and as tiny inclusions within garnet porphyroblasts (stage I). In all occurrences, it is a hercynite–spinel solid solution (0.38 , XMg , 0.47) with variable amounts of magnetite (2.0–4.2 mol%)

(Table 6). The Zn content is appreciable, reaching 5.89 wt% ZnO. The most Zn-rich spinel belongs to the quartz–orthopyroxene–spinel symplectite (reaction (7); Fig. 3f; Table 6). This feature is important, as a high content in this component extends the stability field of the quartz–spinel association towards low temperatures and high pressures (Schulters & Bohlen 1989; Nichols et al. 1992). UHT spinel in direct contact with sapphirine and quartz (i.e. reactions (4) and (5); Fig. 3c and d) shows the highest Cr2O3 content (6.0– 6.8 wt%: Table 6), whereas the late spinel of stage IV has the lowest content (c. 4.00 wt% Cr2O3).

Biotite All samples contain low modal amounts of biotite, which shows small variations in Fe and Mg contents (Table 7; 0.82 , XMg , 0.88). Zoning is only sometimes discernible at the contact with garnet or other ferromagnesian minerals, where it is linked to retrograde exchange of Fe and Mg with these minerals (e.g. Spear 1993). Generally, an increase in Ti content (0.24 –0.28 p.f.u., for 22 equivalent O) is generally correlated with a decrease in Mg and F (0.46 –0.94 p.f.u.) and an increase in Fe. The high XF value (F/(F þ OH) ¼ 0.25 –0.49) is a typical feature of UHT metamorphism (e.g. Hensen & Osanai 1994).

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Table 5. Representative chemical analyses of sapphirine Stage: Sample no.: Reaction:

Stage I Tek96 (3) 4

13.44 SiO2 TiO2 0.17 Al2O3 62.02 FeOt 6.55 MnO 0.20 MgO 16.51 1.36 Cr2O3 ZnO 0.06 Sum 100.32 Cations per 20 oxygens Si 1.590 Ti 0.020 Al 8.660 0.000 Fe3þ 2þ Fe 0.650 Mn 0.020 Mg 2.920 Cr 0.130 Zn 0.000 3þ 2þ 0.00 Fe /Fe 0.82 XMg

Tek96 (1) 4

Tek96 (1) 4

Tek96 (3) 3

Tek96 (1) 6

13.61 0.02 60.12 7.27 0.00 16.19 1.91 0.05 99.26

13.46 0.11 60.30 7.89 0.02 16.31 1.93 0.00 100.05

12.97 0.00 60.62 8.37 0.03 16.41 1.33 0.00 99.72

13.46 0.06 60.58 7.94 0.00 16.31 1.18 0.27 99.94

1.640 0.000 8.520 0.030 0.700 0.000 2.900 0.180 0.000 0.04 0.80

Feldspars The alkali feldspar of sample Tek96 is mainly K-rich (Or53 – 90Ab42 – 10An5 – 0). Plagioclase exists

1.600 0.010 8.490 0.090 0.700 0.000 2.910 0.180 0.000 0.13 0.81

1.550 0.000 8.550 0.220 0.620 0.000 2.930 0.130 0.000 0.35 0.82

1.610 0.010 8.530 0.140 0.650 0.000 2.900 0.110 0.020 0.21 0.82

in samples Tek58 and Tek102, where it exhibits two main habits. An early andesine –labradorite (An30 – 65) occurs as unzoned porphyroblasts in the matrix; it was slightly ternary, as it exsolved

Table 6. Representative chemical analyses of spinel Stage: Sample no.: Reaction:

Stage III Tek96 (3) 5

0.00 SiO2 TiO2 0.02 Al2O3 55.63 FeOt 24.43 MnO 0.20 MgO 11.12 6.83 Cr2O3 ZnO 1.63 Sum 98.00 Cations per 32 oxygens Ti 0.000 Al 14.370 3þ 0.450 Fe Fe2þ 4.030 Mn 0.040 Mg 3.630 Cr 1.184 Zn 0.270 Fe3þ/Fe2þ 0.11 XMg 0.47

Stage IV

Tek96 (3) 5

Tek58 (2) 6

Tek58 (2) 6

Tek58 (1) 10

Tek58 (1) 10

Tek58 (1) 10

0.02 0.00 55.58 24.55 0.00 11.76 6.08 1.68 98.00

0.02 0.07 57.54 21.61 0.00 11.02 4.96 5.89 101.13

0.06 0.00 56.96 20.10 0.09 10.84 4.95 5.58 98.59

0.06 0.00 58.40 22.45 0.00 12.46 3.88 2.25 99.5

0.06 0.00 57.91 22.44 0.12 12.39 4.03 2.38 99.33

0.05 0.00 58.09 22.18 0.03 12.57 4.00 2.35 99.27

0.000 14.830 0.350 3.360 0.020 3.570 0.860 0.930 0.10 0.51

0.000 14.880 0.430 3.630 0.000 4.010 0.660 0.360 0.12 0.52

0.000 14.800 0.480 3.590 0.020 4.000 0.690 0.390 0.13 0.53

0.000 14.830 0.460 3.550 0.000 4.060 0.680 0.380 0.13 0.53

0.001 14.310 0.670 3.810 0.000 3.830 1.050 0.280 0.17 0.50

0.010 14.670 0.490 3.420 0.000 3.550 0.850 0.960 0.14 0.51


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Table 7. Representative chemical analyses of biotite Type: Sample no.:

Core Tek96 (1)

SiO2 39.01 4.62 TiO2 Al2O3 14.63 Cr2O3 0.51 FeO 6.52 MnO 0.00 MgO 19.59 0.16 Na2O K2O 10.01 CaO 0.00 F 2.02 Cl 0.51 3.07 H2O Sum 100.65 Cations per 22 oxygens Si 2.811 Ti 0.251 Al 1.243 Cr 0.029 Mn 0.000 0.393 Fe2þ Mg 2.105 Na 0.023 K 0.920 Ca 0.000 OH 1.477 F 0.460 Cl 0.063 Sum 9.774 0.84 XMg XF(F/F þ OH) 0.24

Intermediate

Rim in contact with garnet

Tek96 (1)

Tek96 (1)

Tek96 (1)

Tek96 (1)

Tek96 (1)

Tek96 (1)

38.12 4.70 13.93 0.09 5.27 0.02 20.29 0.14 10.03 0.09 2.09 0.49 2.95 98.20

39.42 4.30 14.08 0.37 7.57 0.00 20.03 0.13 9.98 0.00 2.22 0.57 2.98 101.64

40.10 3.85 13.73 0.16 5.36 0.00 21.51 0.10 10.30 0.09 2.61 0.61 2.80 101.19

38.09 5.11 14.33 0.64 7.36 0.10 19.82 0.07 9.99 0.00 2.92 0.56 2.62 101.62

37.15 4.36 13.82 0.49 6.64 0.00 20.15 0.12 10.12 0.00 3.45 0.44 2.29 99.04

37.93 5.19 14.25 0.27 6.99 0.08 20.13 0.11 9.80 0.00 4.11 0.62 2.02 101.50

2.808 0.260 1.210 0.005 0.001 0.325 2.227 0.020 0.943 0.007 1.452 0.487 0.061 9.806 0.87 0.25

2.828 0.232 1.191 0.021 0.000 0.455 2.142 0.018 0.913 0.000 1.428 0.503 0.069 9.799 0.82 0.26

an antiperthitic K-feldspar. A secondary Ca-rich plagioclase (An . 80 mol%) developed as vermicular intergrowths and/or coronas around garnet (Figs 3c and 4a).

P – T evolution The mineral associations and textures reported in this study provide evidence of UHT metamorphism in the Khanfous area. This UHT event is also documented by the high Al content in orthopyroxene (i.e. 23 mol% of the Tschermak end-member). The described assemblages include cordierite, sapphirine, orthopyroxene, garnet, sillimanite, quartz and spinel. Plagioclase, rutile and ilmenite are also present in variable but small proportions. Biotite is an accessory mineral rarely involved in metamorphic reactions. By ignoring these minor phases, all reactions can be adequately modelled in the simplified FMASH system

2.869 0.207 1.158 0.009 0.000 0.321 2.294 0.013 0.940 0.007 1.336 0.590 0.074 9.818 0.88 0.31

2.753 0.278 1.221 0.037 0.006 0.445 2.136 0.010 0.921 0.000 1.264 0.667 0.069 9.806 0.83 0.34

2.760 0.244 1.211 0.029 0.000 0.413 2.232 0.017 0.960 0.000 1.134 0.810 0.056 9.865 0.84 0.42

2.755 0.284 1.220 0.015 0.005 0.425 2.179 0.015 0.908 0.000 0.979 0.945 0.076 9.806 0.84 0.49

(FeO –MgO –Al2O3 –SiO2 – H2O). Pioneer models in this system based on theoretical and experimental data at high-T and low-f O2 conditions (e.g. Hensen 1971, 1986, 1987; Hensen & Green 1973; Bertrand et al. 1991) produced a petrogenetic grid with stable [Spl], [Opx] and [Sil] invariant points, as shown in Figure 6 (Hensen & Green 1973; Hensen 1987). Using the modern thermodynamic database of Holland & Powell (1998) and the solid-solution model of Ouzegane et al. (2003a) for sapphirine, a similar P–T petrogenetic grid was modelled in the same system (Fig. 7) with the software Thermocalc v.2.75 (Powell & Holland 1988). Although the [Spl], [Opx] and [Sil] invariant points are stable, their arrangement in the P –T diagram is different, with the [Sil] invariant point lying at a higher temperature (compare Fig. 7 with Fig. 6). This difference could be due to (1) the amount of water, whose activity has an important effect on the position of univariant curves involving cordierite (Fig. 7), and (2) differences in the chemical composition of

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Fig. 6. A partial petrogenetic grid for the FMASH system at low-f O2 from Hensen & Harley (1990), modified after the experimental constraints of Bertrand et al. (1991). The arrow indicates the preferred P –T path based upon the sequence of reaction textures. Phases are projected from quartz onto the Al2O3 – FeO– MgO triangle.

mineral phases used in prior studies (Hensen & Green 1973; Bertrand et al. 1991) and those obtained by modelling. Our sequence of metamorphic stages, with the crystallization of the unusual assemblages sapphirine þ spinel þ quartz and orthopyroxene þ spinel þ quartz, entails a P –T path that oversteps the [Sil] invariant point (see arrow in Figs 6 and 7). This suggests a peak temperature greater than 1170 8C or 1300 8C, depending on the data considered (early experimental data (Fig. 6) or modern thermodynamic data (Fig. 7), respectively). To better constrain the P –T path, we constructed P –T pseudosections, valid for particular compositions. We considered the FMASH system and used the dataset of Holland & Powell (1998) with the solid-solution model of Ouzegane et al. (2003a) for sapphirine, which yields similar results to those of Kelsey et al. (2004). The first pseudosection (Fig. 8) was modelled for the bulk composition of a quartzitic Al– Mg granulite (Table 1) and low water activity (aH2O ¼ 0.1). The metamorphic evolution, from

stage II to stage IV, is in agreement with a single metamorphic cycle. Stage II would correspond to the divariant Opx –Sill– Qtz–Grt field, which implies conditions (,1000 8C; P 9 kbar) consistent with the experimental data of Annersten & Seifert (1981), Audibert et al. (1995) and Carrington & Harley (1995), and with the P– T conditions estimated by Bertrand et al. (1992) for this assemblage in granulites from the In Roccan region (Fig. 1b). The assemblage sapphirine þ quartz + orthopyroxene + sillimanite (i.e. stage III) is stable at temperatures greater than 950 8C, which correlates well with the high content in Tschermak end-member (i.e. 23 mol%) of the orthopyroxene. The transition between stages II and III entails an important increase of T. However, the pseudosection does not show fields corresponding to the crucial parageneses orthopyroxene þ spinel þ quartz and sapphirine þ spinel þ quartz (i.e. reactions (4), (5) and (7) of stage III), probably because these are stable for microdomain compositions different from that considered for the pseudosection. The growth of cordierite–orthopyroxene-


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Fig. 7. Petrogenetic grid calculated in the FMASH system using Thermocalc. Circles represent the evolution of invariant points at different aH2O.

Fig. 8. P– T pseudosection for the bulk composition of a quartz-rich Al–Mg granulite sample. The bold continuous curve represents the P– T path during the metamorphic stages on the basis of textural analysis. The numbered univariant curves correspond to: (1) Opx þ Sil , Grt þ Spr þ Qtz (Crd, Spl); (2) Opx þ Spr þ Qtz , Crd þ Grt (Sil, Spl); (3) Spr þ Qtz , Grt þ Crd þ Sil (Opx, Spl); (4) Opx þ Sil , Crd þ Spr þ Grt (Spl, Qtz); (5) Opx þ Sil þ Qtz , Crd þ Grt (Spl, Spr).

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Fig. 9. P– T pseudosection for a specific bulk composition corresponding to the sapphirine þ orthopyroxene þ quartz þ garnet assemblage. The stippled region denotes the UHT metamorphic peak estimated by the intersection Opx of the XGrt pyrope and XTschermak isopleths. The univariant reaction highlighted in bold is Opx þ Sil , Grt þ Spr þ Qtz (Crd, Spl). Phases in parentheses have a negligible abundance.

Motoyoshi (2000), who applied the same method to a similar assemblage (Al-rich orthopyroxene þ sapphirine þ quartz) from the Napier Complex (Antarctica). We also tried to construct P –T pseudosections for the composition of the UHT spinel-bearing symplectites, using the database of Holland & Powell (1998). Unfortunately, we did not obtain

10 kbar

8 kbar 0.25

1000 °C

0.20

XAI

bearing symplectites and coronas (i.e. stage IV) is consistent with a decrease of P (,7.5 kbar) and T, towards fields where cordierite and orthopyroxene are co-stable. However, the precise arrangement of these cordierite-bearing fields depends on the water activity (see Fig. 7), in contrast to the anhydrous cordierite-free assemblages. To unravel the P– T conditions at the thermal peak, we calculated a second P –T pseudosection for a specific composition corresponding to the UHT symplectite orthopyroxene þ sapphirine þ quartz (Fig. 9), obtained by balancing the reaction Grt , Opx þ Spr þ Qtz (reaction (6)). The trivariant and divariant fields Grt ( –Qtz –Sil) and Opx– Spr–Qtz ( –Grt) show the stability fields for the reactant and products of reaction (6), respectively. Therefore, a P– T evolution from the first field towards the second is necessary to account for the reaction. The median stability temperature for this symplectite, estimated by the intersection of Opr isopleths (XGrt pyrope –Xenstatite), is close to 1200 8C (Fig. 9). These extreme conditions are confirmed by orthopyroxene thermometry based on the solubility of Al in orthopyroxene (Fig. 10), using Opx Opx and Xenstatite isopleths given by the XTschermak Hensen & Harley (1990). According to this diagram, metamorphic conditions attained during the UHT stage are about 1100 8C and 10 kbar. These results agree with those of Harley &

6 kbar 0.15

0.10

0.05 0.5

900 °C

800 °C

0.55

0.60

0.65

0.70

0.75

XMg Fig. 10. Orthopyroxene compositions plotted in the Opx Opx (XAl) v. Xenstatite (XMg) diagram. Isotherms and XTschermak isobars are from Hensen & Harley (1990).


the expected fields (Opx– Spl–Qtz –Grt and Spl–Qtz –Sil –Grt). This could be due to the simplified FMAS system that is considered while modelling the pseudosections, whereas our spinel is rich in Fe3þ, Cr and Zn (Table 6). These features shift towards lower temperatures and extend the P–T stability domains of the spinel-bearing assemblages (e.g. Guiraud et al. 1996; White et al. 2002). The P–T estimates of the Khanfous Al –Mg granulites indicate a clockwise P–T evolution that occurred in static and anhydrous conditions, with a thermal peak exceptionally high in temperature (1150–1300 8C), which was responsible for the appearance of unique assemblages.

Conclusions Al– Mg granulites from the Khanfous area preserve remarkable UHT mineral assemblages that were only partially consumed during the late retrograde evolution. The estimated P –T path indicates a peak temperature as high as 1150– 1300 8C, at low water activity. Peculiar parageneses, namely spinel þ orthopyroxene þ quartz, sapphirine þ spinel þ quartz, spinel þ quartz and sapphirine þ quartz, developed before the cordierite-bearing assemblages and are a record of the UHT event. Whereas the rocks underwent an important ductile deformation prior to stage II, all the subsequent UHT clockwise P–T evolution took place without any deformation, as evidenced by the microtextures of the described symplectites and coronas, which grew statically. Our data suggest that the Khanfous area, as well as the whole of the northern In Ouzzal terrane, experienced UHT crustal metamorphism followed by exhumation along a clockwise P–T path. The P–T history of the In Ouzzal metacraton is considered to be continuous and related to the Eburnean event (2.0 Ga, Ouzegane et al. 2003b). Several features of this metacraton distinguish it from the others. In most UHT terranes, prograde P –T histories are unknown or obliterated by extreme peak temperatures. On the other hand, in the In Ouzzal, the prograde P–T path is well documented by the orthopyroxene-bearing assemblages prior to sapphirine-bearing parageneses (Ouzegane et al. 2003a; this study). The peak temperature deduced from this study is consistent with decompression and heating to temperatures higher than those of the [Spl], [Opx] and [Sil] FMASH invariant points. This is an exceptional feature, as all samples that have experienced isothermal decompression (e.g. Harley 1998; Raith et al. 1997) do not exceed temperatures higher than the [Spl] invariant point.

165

UHT metamorphism is characterized by a variety of P–T paths, some typical of isothermal decompression (ITD) and others of isobaric cooling (IBC) or hybrid ITD–IBC (Harley 2004). The UHT clockwise P–T path and the rapid postpeak exhumation of the deep crust characteristic of ITD can only be explained by an ad hoc thermal model. The cause of the thermal anomaly lies below the crust and may be ascribed to, for example, heat from mantle plumes or advection of heat as a result of magma emplacement. The UHT –IBC P–T path is often considered to be the result of advection of heat into the continental crust as a result of underplating by mantle-derived basaltic magmas and subsequent cooling (Bohlen 1987). However, the convective removal of the lithospheric mantle after crustal thickening may explain UHT metamorphism followed by ITD (Platt et al. 1998). In the case of the In Ouzzal terrane, to bring the lower crust to ultra-high temperatures Ouzegane et al. (2003b) suggested a frontal hypercollision during the Eburnean orogeny at 2.0 Ga that removed the lithospheric mantle, bringing the asthenosphere close to the Moho. The transfer of heat from the asthenosphere to the lower crust would lead to UHT metamorphic conditions adequate to determine the static growth of our peculiar parageneses. Thermal anomalies are unstable, and the return to more normal conditions is indicated by the last cooling coupled with decompression (stage IV). The recovery of the lithospheric mantle during the long post-Eburnean period led to the formation of a thick cratonic lithosphere that protected the In Ouzzal during the Pan-African collision (0.6 Ga) and the subduction-related stages documented in other parts of the Hoggar shield. The Pan-African orogeny did not affect the studied rocks and only reactivated discrete brittle faults: the rigidity of the In Ouzzal ‘metacraton’, although limited in size, helped preserve these Eburnean UHT parageneses. We thank Y. Osanai and K. Sato for their detailed and constructive reviews, which helped us to improve the manuscript. We also express our gratitude to A. F. Palladino, who corrected the English style. This work is a contribution to the projects TASSILI 05 MDU 653 ‘Imagerie tridimentionnelle et e´volution spaciotemporelle du Hoggar’ and NATO EST/CLE 979766. We are also extremely grateful to ORGM and OPNA for logistic support during fieldwork.

References A NNERSTEN , H. & S EIFERT , F. 1981. Stability of the assemblage orthopyroxene–sillimanite– quartz in the system MgO–FeOFe2O3 –Al2O3 –SiO2 – H2O. Contributions to Mineralogy and Petrology, 77, 158– 165.

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A UDIBERT , N., H ENSEN , B. J. & B ERTRAND , P. 1995. Experimental study of phase relations involving osumilite in the system K2O–FeO– MgO–Al2O3 – SiO2 –H2O at high pressure and temperature. Journal of Metamorphic Geology, 13, 331– 344. B ERNARD -G RIFFITHS , J., F OURCADE , S., K IENAST , J. R., P EUCAT , J. J., M ARTINEAU , F. & R AHMANI , A. 1996. Geochemistry and isotope Sr, Nd, O study of Al– Mg granulites from the In Ouzzal Archaean Block Hoggar, Algeria. Journal of Metamorphic Geology, 14, 709–724. B ERTRAND , P., E LLIS , D. J. & G REEN , D. H. 1991. The stability of sapphirine –quartz and hypersthene– sillimanite– quartz assemblages: an experimental investigation in the system FeO–MgO –Al2O3 –SiO2 under H2O and CO2. Contributions to Mineralogy and Petrology, 108, 55– 71. B ERTRAND , P., O UZEGANE , K. & K IENAST , J. R. 1992. P –T –X relationships in the Precambrian Al– Mg rich granulites from In Ouzzal Hoggar, Algeria. Journal of Metamorphic Geology, 10, 17– 31. B LACK , R., L ATOUCHE , L., L IE´ GEOIS , J. P., C ABY , R. & B ERTRAND , J. M. 1994. Pan-African displaced terranes in the Tuareg shield (central Sahara). Geology, 22, 641 –644. B OHLEN , S. R. 1987. Pressure– temperature– time paths and a tectonic model for the evolution of granulites. Journal of Geology, 95, 617– 632. C ARRINGTON , D. P. & H ARLEY , S. L. 1995. The stability of osumilite in metapelitic granulites. Journal of Metamorphic Geology, 13, 613–625. D ALLWITZ , W. B. 1968. Coexisting sapphirine and quartz in granulite from Enderby Land, Antartica. Nature, 219, 476–477. D RARENI , A., O UZEGANE , K. & B ENDAOUD , A. 2007. L’Archeén du Hoggar: geóchimie et geódynamique. Bulletin du Service Geólogique d’ Alge´rie, 18, 1– 21. E LLIS , D. J. 1980. Osumilite–sapphirine–quartz granulites from Enderby Land, Antarctica: P– T conditions of metamorphism, implications for garnet–cordierite equilibria and the evolution of the deep crust. Contributions to Mineralogy and Petrology, 74, 201–210. F OURCADE , S., K IENAST , J. R. & O UZEGANE , K. 1996. Metasomatic effects related to channeled fluid streaming through deep crust: fenites and associated carbonatites In Ouzzal Proterozoic granulites, Hoggar, Algeria. Journal of Metamorphic Geology, 14, 763 –781. G REW , E. S. 1982. Osumilite in the sapphirine–quartz terrane of Enderby Land, Antartica: Implications for osumilite petrogenesis in the granulitic facies. American Mineralogist, 67, 762–787. G UIRAUD , M., K IENAST , J. R. & R AHMANI , A. 1996. Petrological study of high temperature granulites from In Ouzzal, Algeria: some implications of the phase relationships in the FMASTOCr system. European Journal of Mineralogy, 8, 1375– 1390. H ADDOUM , H., C HOUKROUNE , P. & P EUCAT , J. J. 1994. Evolution of the Precambrian In Ouzzal block (Central Sahara, Algeria). Precambrian Research, 65, 155–166. H ARLEY , S. L. 1985. Garnet– orthopyroxene bearing granulites from Enderby Land, Antartica: implication for FMAS petrogenetic grids in the granulite facies.

Contributions to Mineralogy and Petrology, 94, 452–460. H ARLEY , S. L. 1998. Ultrahigh temperature granulite metamorphism (1050 8C, 12 kbar) and decompression in garnet (Mg70)–orthopyroxene– sillimanite gneisses from the Rauer Group, East Antarctica. Journal of Metamorphic Geology, 16, 541 –562. H ARLEY , S. L. 2004. Extending our understanding of ultrahigh temperature crustal metamorphism. Journal of Mineralogical and Petrological Sciences, 99, 140–158. H ARLEY , S. L. & M OTOYOSHI , Y. 2000. Al zoning in orthopyroxene in a sapphirine quartzite: evidence for .1120 8C UHT metamorphism in the Napier Complex, Antarctica, and implications for the entropy of sapphirine. Contributions to Mineralogy and Petrology, 138, 293–307. H ENSEN , B. J. 1971. Theoretical phase relations involving cordierite and garnet in the system MgO–FeO– Al2O3 – SiO2. Contributions to Mineralogy and Petrology, 33, 191–194. H ENSEN , B. J. 1986. Theoretical phase relations involving garnet and cordierite revisited: the influence of oxygen fugacity on the stability of sapphirine and spinel in the system Mg–Fe– Al– Si– O. Contributions to Mineralogy and Petrology, 92, 362– 367. H ENSEN , B. J. 1987. P– T grids for silica undersaturated granulites in the system MAS (n þ 4) and FMAS (n þ 3): tools for the understanding of P– T paths of metamorphism. Journal of Metamorphic Geology, 5, 255–271. H ENSEN , B. J. & G REEN , D. H. 1973. Experimental study of the stability of cordierite and garnet in pelitic compositions at high pressures and temperatures III. Synthesis of experimental data and geological applications. Contributions to Mineralogy and Petrology, 38, 151–166. H ENSEN , B. J. & H ARLEY , S. L. 1990. Graphical analysis of P –T –X relations in granulite facies metapelites. In: A SHWORTH , J. R. & B ROWN , M. (eds), Hightemperature Metamorphism and Crustal Anatexis. Unwin Hyman, London, 19– 56. H ENSEN , B. J. & O SANAI , Y. 1994. Experimental study of dehydration melting of F-bearing biotite in model pelitic compositions. Mineralogical Magazine, 58A, 410–411. H IGGINS , J. B., R IBBE , P. H. & H ERD , R. K. 1979. Sapphirine I: crystal chemical contributions. Contributions to Mineralogy and Petrology, 68, 349– 356. H OLLAND , T. J. B. & P OWELL , R. 1998. An internallyconsistent thermodynamic data set for phases of petrological interest. Journal of Metamorphic Geology, 16, 309– 343. K ELSEY , D. E., W HITE , R. W., H OLLAND , T. J. B. & P OWELL , R. 2004. Calculated phase equilibria in K2O–FeO– MgO–Al2O3 –SiO2 –H2O for sapphirine–quartz-bearing assemblages. Journal of Metamorphic Geology, 22, 559– 578. K IENAST , J. R. & O UZEGANE , K. 1987. Polymetamorphic Al–Mg rich granulites with orthopyroxene-sillimanite and sapphirine parageneses in Archean rocks from Hoggar, Algeria. In: K INNAIRD , J. & B OWDEN , P. (eds) African Geology Reviews Geological Journal (thematic issue), 22, 57–79.

ULTRAHIGH-I MINERALS, IN OUZZAL M OURI , H., G UIRAUD , M. & K IENAST , J. R. 1993. Al– Mg granulites of Ihouhaouene, Hoggar, Algeria: An example of phase relationships in the KFMASH system and melt absent equilibria. Comptes Rendus de l’Acade´mie des Sciences, 316, 1565–1572. M OURI , H., G UIRAUD , M. & K IENAST , J. R. 1994. L’origine des granulites Al–Mg d’Ihouhaouene (Hoggar–Alge´rie): alte´ration de roches basiques et ultrabasiques a` l’Archeén. Comptes Rendus de l’Acade´mie des Sciences, 318, 941– 948. N ICHOLS , G. T., B ERRY , R. F. & G REEN , D. H. 1992. Internally consistent gahnite spinel– cordierite– garnet equilibria in the FMASHZn system: geothermobarometry and applications. Contributions to Mineralogy and Petrology, 111, 362–377. O UZEGANE , K. & B OUMAZA , S. 1996. An example of very high temperature metamorphism: orthopyroxene–sillimanite– garnet, sapphirine– quartz, and spinel–quartz. Journal of metamorphic Geology, 14, 693–708. O UZEGANE , K. & K IENAST , J. R. 1996. Nature et e´volution des se´ries me´tamorphiques de tre`s haute tempe´rature de l’Unite´ Granulitique de l’In Ouzzal Ouest Hoggar. Bulletin du Service Geólogique de l’Alge´rie, 7, 133– 157. O UZEGANE , K., G UIRAUD , M. & K IENAST , J. R. 2003a. Prograde and retrograde evolution in high temperature corundum granulites. FMAS and KFMASH systems from In Ouzzal terrane, NW Hoggar, Algeria. Journal of Petrology, 44, 517–545. O UZEGANE , K., K IENAST , J. R., B ENDAOUD , A. & D RARENI , A. 2003b. A review of Archaean and Paleoproterozoic evolution of the In Ouzzal granulitic terrane (Western Hoggar, Algeria). Journal of African Earth Sciences, 37, 207–227. P ERCHUK , L. L, A RANOVICH , L. YA ., P ODELESSKII , K. K. ET AL . 1985. Precambrian granulites of the Aldan Shield, eastern Siberia, U.S.S.R. Journal of Metamorphic Geology, 3, 265–310. P EUCAT , J. J., C APDEVILA , R., D RARENI , A., C HOUKROUNE , P., F ANNING , M., B ERNARD -G RIFFITHS , J. & F OURCADE , S. 1996. Major and trace element geochemistry and isotope Sr, Nd, Pb, O systematics of an Archaean basement involved in a 2.0 Ga VHT 1000 8C metamorphic event: In Ouzzal massif, Hoggar, Algeria. Journal of Metamorphic Geology, 14, 667–692. P LATT , J. P., S OTO , J. I., W HITEHOUSE , M. J., H URFORD , A. J. & K ELLEY , S. P. 1998. Thermal

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evolution, rate of exhumation, and tectonic significance of metamorphic rocks from the floor of the Alboran extensional basin, western, Mediterranean. Tectonics, 17, 671– 689. P OWELL , R. & H OLLAND , T. J. B. 1988. An internally consistent thermodynamic dataset with uncertainties and correlations: 3. Application methods, worked examples and a computer program. Journal of Metamorphic Geology, 6, 173 –204. R AITH , M., K ARMAKAR , S. & B ROWN , M. 1997. Ultrahigh temperature metamorphism and multistage decompressional evolution of sapphirine granulites from the Palni Hill Ranges, southern India. Journal of Metamorphic Geology, 15, 379– 400. S AJEEV , K. & O SANAI , Y. 2004. Osumilite and spinel þ quartz from Sri Lanka: Implications for UHT conditions and retrograde P –T path. Journal of Mineralogical and Petrological Sciences, 99, 320– 327. S ANDIFORD , M. 1985. The metamorphic evolution of granulites at Fyfe Hills: implications for Archaean crustal thickness in Enderby Land, Antarctica. Journal of Metamorphic Geology, 3, 155– 178. S CHULTERS , J. & B OHLEN , S. R. 1989. The stability of hercynite and hercynite–gahnite spinels in corundumor quartz-bearing assemblages. Journal of Petrology, 30, 1017–1031. S HERATON , J. W., O FFE , L. A., T INGEY , R. J. & E LLIS , D. J. 1980. Enderby Land, Antarctica, an unusual Precambrian high-grade metamorphic terrain. Journal of the Geological Society of Australia, 27, 1 –18. S HERATON , J. W., T INGEY , R. J., B LACK , L. P., O FFE , L. A. & E LLIS , D. J. 1987. Geology of Enderby Land and Western Kemp Land, Antarctica. Bulletin of Australian Bureau of Mineralogical Resources, 223, 1–51. S PEAR , F. S. 1993. Metamorphism of ultramafic and cordierite anthophyllite rocks. In: Metamorphic Phase Equilibria and Pressure Temperature Time Paths. Mineralogical Society of America, Monograph Series, 393–489. W ATERS , D. J. 1991. Hercynite– quartz granulites: phase relations, and implications for crustal processes. European Journal of Mineralogy, 3, 367–386. W HITE , R. W., P OWELL , R. & C LARKE , G. L. 2002. The interpretation of reaction textures in Fe-rich metapelitic granulites of the Musgrave Block, central Australia: constraints from mineral equilibria calculations in the system K2O–FeO– MgO–Al2O3 –SiO2 –H2O– TiO2 – Fe2O3. Journal of Metamorphic Geology, 20, 41– 55.

Review of the orogenic belts on the western side of the West African craton: the Bassarides, Rokelides and Mauritanides MICHEL VILLENEUVE FRE CNRS 2761, Universite´ de Provence, case 67, 3, place Victor Hugo, 13331, Marseille, Cedex 03, France (e-mail: [email protected]) Abstract: The West African craton is fringed along its western side by a 3000 km long orogenic belt subdivided into three separate orogens: the Bassaride (Pan-African I orogeny), Rokelide (Pan-African II orogeny) and Mauritanide (Hercynian orogeny) thrust belts. The Bassarides are cut to the north by the Mauritanides and to the south by the Rokelides but parts of this Bassaride belt are incorporated in the other two younger belts. This review presents the main geological, geophysical and geochronological results from the western side of the West African craton, collected over the past 90 years, concentrating on those from the last 15 years. Former geological investigations underlined the thin-skinned structure model within these thrust belts, whereas the geophysical results gave prominence to the major importance of block faulting resulting from the Pan-African I orogeny and its strong influence on the subsequent orogenic belt features. The geochronological data allow us to distinguish major tectonothermal events related to the Pan-African I (660–650 Ma), Pan-African II (550– 530 Ma) and Hercynian (330–300 Ma and 280–270 Ma) orogenies. However, they also reveal five other tectonothermal events (at 1200– 1000, 750 –700, 600– 580, 510–480 and 450–380 Ma), which are still very poorly understood. The 1200– 1000 Ma tectonothermal event recently revealed in the northern Mauritanides may correspond to a remanent orogenic belt segment that witnessed the Grenvillian orogeny.

The West African craton (WAC) (Fig. 1) is surrounded by several Late Neoproterozoic to Palaeozoic orogenic belts. Those linked with the Trans-Saharan suture zone fringe the eastern margin of the WAC. This mobile zone can be traced from Benin to Morocco and includes the Dahomeyides, Hoggar –Iforas, Ougarta and Pharusian belts. These belts were folded and metamorphosed at around 600 Ma, has been proposed by various workers (Black et al. 1979; Caby et al. 1981; Affaton et al. 1991; Castaing 1993). The western margin is a polyphase mobile zone that can be traced from Liberia to Morocco. The Rokelide and Bassaride belts are the orogenic result of two distinct Pan-African orogenic phases and have been crosscut by the Hercynian orogeny, which mainly contributed to the formation of the Mauritanide belt. However, some signs of an older phase, which possibly formed the Souttoufide tectonothermal event (Villeneuve et al. 2006) around 1200–1000 Ma, are suspected in the northern Mauritanide belt (Villeneuve et al. 2006). The WAC itself is composed of a crystalline basement that stabilized around 2000 Ma, with a Neoproterozoic to Palaeozoic sedimentary cover. The Archaean (.2500 Ma) and Birimian (2000– 1700 Ma) basements of the WAC (Fig. 1) are mainly exposed within the Reguibat shield (north) and the Ivory Coast shield (ICS) or Leo uplift (south). Small inliers also exist in the

western part of the craton, along the Bassaride and Mauritanide belts (Kenieba and Kayes inliers) as well as within the Anti-Atlas belt. The Saõ Luı´s craton is also considered as a fragment of the WAC and finally appeared, after the breakup of Pangaea, on the South American platform (Klein et al. 2005). The Mauritanide belt is the name given by Sougy (1962a) to the fold belt extending from southern Morocco to northern Senegal. The Rokelide belt (from Guinea–Bissau to Liberia) was named by Allen (1967), and the Bassaride belt (southern Senegal to Guinea-Bissau) was named by Villeneuve (1984). The contact between the Hercynian Mauritanide belt and the sedimentary cover of the Taoudeni and Tindouf basins is called the MFT (Mauritanian frontal thrust or hercynian frontal thrust), implying a transpressive tectonic regime with eastward overthrusting of the Mauritanide belt on top of both foreland basins. However, it clearly appears that the origin and the depositional infilling of the Taoudeni and peripheral basins were partly controlled by the geodynamic evolution of the surrounding fold belts and had already started some 1000 Ma ago. For the time being, it can be considered that three main tectonic events were involved in the structuring of the belts of the WAC: the PanAfrican I tectonic event linked to the Bassaride belt (660 –650 Ma), the Pan-African II tectonic


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Fig. 1. The main structural features of the West African craton (after Villeneuve 1984, modified) and the location of the Mauritanide, Bassaride and Rokelide belts. Circled numbers: 1, Mauritanide belt; 2, Bassaride belt; 3, Rokelide belt; 4, Anti-Atlas –Ougarta belt; 5 –7, Trans-Saharan mobile zone (subdivided into: 5, Pharusian belt; 6, Hoggar– Iforas belt; 7, Dahomeyide belt); 8, possible Souttoufide belt. Tf.B, Tindouf basin; Td.B, Taoudeni basin; WAC, West African craton; ICS, Ivory Coast shield (Leo uplift); SLC, Saõ Luis craton. Legend: 1, crystalline basement; 2, main Hercynian belts; 3, Pan-African belts; 4, foreland basin; 5, thrust directions; 6, thrusts; 7, basin boundaries; 8, study areas; 9, shear zone.

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event linked to the Rokelide belt in the southwestern part (550 –530 Ma), and the Hercynian event linked to Mauritanide belt in the northern part (Sougy 1962a; Allen 1967; Villeneuve & Dallmeyer 1987). The main disconformities that are recognized in the foreland basins together with the main tectonic events identified within the thrust belts are summarized in Figure 2. The main stratigraphical markers, apart from Silurian, Devonian and Carboniferous fossils, are deposits from glaciations, which occurred at the end of the Neoproterozoic and at the end of the Ordovician.

Previous work The geological exploration of West Africa began in the 20th century. Four main exploration periods can be distinguished: (1) from 1917 up to the first geological review of West Africa by Roques (1948); (2) from 1948 up to the special issue of the Bulletin de la Socie´te´ Geólogique de France devoted to the Mauritanides and their foreland, edited by Sougy (1969); (3) from 1969 until 1991, with the review of Dallmeyer & Lecorche´ (1991); (4) from 1991 until now (this paper). At least 120 geologists and geophysicists from more than 20 countries collaborated in this research.

Structural geology General overview The orogenic thrust belt fringing the western margin of the WAC is less than 200 km wide, and extends from southern Morocco to Liberia over a distance of more than 3000 km. From north to south, we distinguish three orogenic belts: (1) the Mauritanide belt (Fig. 3), which represents the main segment and extends from the Anti-Atlas in Morocco to the north of the Bove´ basin in Guinea-Bissau, covering some 2200 km; (2) the Bassaride belt (Fig. 5), from southern Senegal to southern Guinea-Bissau; (3) the Rokelide belt (Fig. 5) from the Bove´ basin to southwestern Liberia. Classically two main structures, a ‘foreland’ and a ‘thrust belt’ are distinguished in each of these orogenic belts. The ‘foreland’, which is the external zone of the belt, and which is a stable area marginal to to the orogenic thrust belt, corresponds to the deformed margin of the WAC, whereas the ‘thrust belt’ structure itself represents the tectonized and metamorphosed part of the orogenic belt, corresponding to its internal zone. The contact between the two structures is mainly tectonic. The thrust plane separating the thrust belt from its foreland zone is named the ‘frontal thrust’. Therefore,

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considering the three orogenic thrust belts, we have to delineate three frontal thrusts: the Mauritanide frontal thrust (MFT), the Bassaride frontal thrust (BFT) and the Rokelide frontal thrust (RFT).

The Mauritanide orogenic belt Figure 3 provides a geological sketch map of the entire Mauritanide thrust belt and its foreland, from the western Anti-Atlas in Morocco, to the north of the Bove´ basin in Guinea-Bissau. Three segments can be distinguished from north to south: the northern, central and southern Mauritanides. Seven profiles (1–7) displayed in Figure 3 will be discussed in detail. They are located in the following areas, from north to south: (I) the AntiAtlas; (II) the Dhlou–Sekkem belt, fringing the western part of the Tindouf basin; (III) the Adrar Souttouf (or Oulad Dlim) area; (IV) the Akjoujt area; (V) the central part of the belt from Aouker to Kidira; (VI) within its southern part, the Koulountou –Bove´ area in southern Senegal and northern Guinea-Bissau. The Mauritanide frontal thrust (MFT). This extends from the western end of the Anti-Atlas to Adrar Souttouf (III in Fig. 3) and from the south of the Reguibat uplift up to the north of Guinea-Bissau. In the Anti-Atlas (Hoepffner et al. 2005), the MFT crops out west of Goulimine, between the ‘plage blanche’ and the ‘Bas Draa’ inlier (Hoepffner et al. 2005). In the Dhlou belt, Dacheux (1967) observed many north–south-trending west-dipping thrust faults. In the Adrar Souttouf the phyllitic series of the internal ‘nappes’ are thrust over the Ordovician to Devonian sediments (Sougy 1962b). In the Akjoujt much evidence for thrusting come from the observation of metamorphic ‘nappes’ overriding the Neoproterozoic and Palaeozoic Taoudeni sedimentary cover. The most famous thrust is the ‘Guelb el Hadej’ klippe (30 km NE of Akjoujt city), described by Teissier et al. (1961). In the central Mauritanides, the MFT is, according to Houdry (1990), very close to the Ordovician Assaba cliff located in the south of the ‘Diouk pass’. This interpretation has been confirmed by Lafrance et al. (1993). Lepage (1983) and Diop (1996) extended the MFT to Senegal. To the south, the MFT crops out in southeastern Senegal, in the Niokola–Koba area (Villeneuve 1984), where it separates the Koulountou Group (to the west) from the Niokolo– Koba Group (to the east). In Guinea-Bissau the MFT is supposedly located between Pirada and Canquelifa. The Mauritanide foreland. The foreland includes the crystalline basement and the Neoproterozoic and Palaeozoic sedimentary cover of the Tindouf,

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Fig. 2. The main geological events in West Africa, based on field observations made on the West African orogenic belts and their overrided eastern foreland basins (after Villeneuve 2005, modified). Circled numbers: 1, tillitic formation; 2, volcanic formation; 3, coarse clastic formation; 4, basal conglomerates; 5, unconformity; 6, basin or trough. Sout. Orog., Adrar Souttouf orogeny (Souttoufides).


Taoudeni and Bove´ basins (Fig. 4). The crystalline basement crops out within several tectonic windows, such as in the ‘Reguibat uplift’ in the Anti-Atlas belt, in the small inliers of Kayes and Kedougou (Figs 1 and 3), and in the western end of the ‘Leo uplift’. Sheets of the crystalline basement have been incorporated within the thrust belt area (unit 12 in Fig. 3). The Neoproterozoic and Palaeozoic sedimentary cover, rocks crop out in several basins; from north to south, these are in the Tindouf basin, the Taoudeni basin, the Fale´me´ trough, the Youkounkoun basin and the Bove´ basin. (1) The Tindouf basin. This is separated from the Taoudeni basin by the Reguibat uplift. It forms an elongated WSW–ENE-trending (800 km long) asymmetrical trough. Only the two flanks of the basin are exposed; the central part is concealed below the so-called Hamadian formations of Cretaceous, Tertiary and Quaternary age. Its gently dipping southern flank wedges out southward, onlapping the Reguibat shield, whereas along the northern flank in the Anti-Atlas area, a thicker and more complete Palaeozoic cover sequence is observed; this is up to 10 000 m, close to the city of Tindouf. On the southern flank, Gevin (1960) noted an onlapping of Late Ordovician sandstones on top of the crystalline basement of the Reguibat uplift. These sandstones are capped by a Late Ordovician ‘tillite’, Silurian graptolitic shales and Devonian limestones and sandstones. On top of this lithostratigraphic pile occurs a conglomeratic unit capped by Late Carbonifereous red argillites (red beds). Along the northern flank, the Palaeozoic cover starts with the Adoudounian limestones (Jeannette & Schumacher 1976; Jeannette et al. 1981), which consist of a lower and upper limestone formation with a reddish sandy to schistoid formation (Se´rie Lie de Vin) between them (Destombes et al. 1985). The Adoudounian limestones are capped by the Tabanites sandstones and by the Zini schists and limestones (Ordovician). Late Ordovician, Silurian, Devonian and Early Carboniferous sediments, starting with the second ‘Bani sandstone’ marker bed, are also well represented in this northern flank of the basin. In the Zemmour area (profile 1, Fig. 3) older rocks appear below the Late Ordovician erosional surface. They consist of dolomitic carbonates with interfingered shales (El Tlethyate Group), which thin out rapidly toward the basement. Because of the presence of stromatolites similar to those of the Atar Group (Supergroup 1 of the Taoudeni Basin), a Late Proterozoic age has been proposed by Sougy (1964). East of the Adrar Souttouf area (profile 2, Fig. 3), the allochthonous formations of the Adrar Souttouf are tectonically overlapping Palaeozoic

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sedimentary deposits. They start with a Late Ordovician tillitic level (Bronner & Sougy 1969) capped by red and black sandstones and red and blue limestones interbedded with shales and sandstones ascribed to the Silurian– Devonian. However, because of the folding and faulting of these formations, no realistic thickness was proposed by Sougy (1962b) or Rjimati & Zemmouri (2002). The onlapping geometry of the sedimentary sequences above the basement suggests that the Reguibat shield was an uplifted area during most of Late Proterozoic and Early Palaeozoic times. It probably acted as a shoal between the subsiding Tindouf basin and the epeirogenic Taoudeni basin. The Tindouf basin’s northern flank underwent a Hercynian tectonic compression as demonstrated by Burkhard et al. (2006), who considered a shortening accommodated by crustal faulting of at least 15 km to at most 25 km. The southern flank is not affected by this tectonic compressional regime. (2) The Taoudeni basin. The Taoudeni basin is a large basin cropping out in the central ‘depression’ of the WAC and extending over 2 106 km2. It has been divided into eight sub-basins but only two are in close relationship with the western fold belt: the Taganet sub-basin to the north and the Tambaoura sub-basin to the south. They are separated by the Affole´ high (Fig. 3) and by the Bissau – Kidira– Kayes fault zone (BKKF). The Taganet sub-basin is a gently dipping syncline with depocentre located in the Taganet ‘depression’. The most complete sedimentary sequence is exposed in the Adrar Cliffs. According to Trompette (1973) and Deynoux (1980) this succession is divided into four supergroups separated by three main disconformities. The lithostratigraphic succession is shown in Figure 4. However, this lithostratigraphic succession includes the Bakoye (or Wassangara) Group, which crops out only in the southern Bakoye sub-basin. Supergroup 1 is of Proterozoic age and includes four groups (the Bakoye Group does not occur in the Adrar area). The older deposits have an age of c. 1000 Ma and the younger ones are capped by a Neoproterozoic marker level, the ‘Jbeliat group’, also called the ‘Triad’ (TR in Fig. 4) (comprising a triple association of diamictites, dolomites and bedded cherts, further clarified below, in the paragraph on Supergroup 2). This ‘Triad’, which eroded the Hassabet-el-Hassiane Group in the Adrar area, is supposedly a witness of the last glacial deposits related to the main glaciogenic Neoproterozoic event, which is dominant in the southern Bakoye (or Wassangara) Group. Supergroup 2 comprises glacial deposits and starts distinctively with the ‘Triad’ unit. It rests with an erosional and angular unconformity upon Supergroup 1 or directly upon the basement.

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Fig. 3. Geological scheme of the Mauritanide belt. Circled numbers: I, Anti-Atlas area; II, Zemmour area; III, Adrar Souttouf area; IV, Akjoujt area; V, Moudjeria–Bakel area; VI, Koulountou area. 1– 7, cross-sections shown in Figure 7; HML, high metamorphism line; TZTF, Tizi n’Test fault zone; BKKF, Bissau–Kidira–Kayes fault zone; MKB, Madina–Kouta basin; Yb, Youkounkoun basin; ICS, Ivory Coast shield; KDG, Kedougou inlier; IF, Ifni inlier; Tf, Tarfaya; Boujd, Boujdour; Sm, Smara; Tn, Tindouf; Dh, Dahkla; Aou, Aoucert; AkT, Akjoujt; Ndb, Nouadhibou; NK, Nouakchott; DKR, Dakar; CNK, Conakry; Ky, Kayes; KDG, Kedougou; ML, Maghta Lahjar;


The glacial deposits are capped by a thin and discontinuous dolomitic horizon (cap carbonates), in turn overlain by bedded cherts and green shales. The lithostratigraphic association (diamictites –dolomites –bedded cherts) known as the ‘Triad’ has been used for a long time as a marker horizon (Zimmermann 1960). The age of the last Neoproterozoic glacial event is strongly debated because the shales associated with the glacial level have been dated between 630 and 595 Ma by Clauer & Deynoux (1987). The microfossils from the cap carbonates of the Senegalese ‘Triad’ unit provide, however, an early Cambrian age (Culver et al. 1988; Culver & Hunt 1991). Although the age of the Precambrian–Cambrian boundary does not coincide with the unconformity between Supergroups 1 and 2, Villeneuve (2005) considered this major unconformity as the onset of the Palaeozoic sedimentary cover sequence in the Taoudeni basin. These glacial deposits are capped with 200 –300 m thick marine green shales and siltstones, which pass upward into 250 –300 m thick shallow marine to lagoonal detrital deposits. The rest of the megasequence consists of 300 m thick fluvial cross-bedded sandstones unconformably overlain by 400 m thick transgressive shallow marine Scolithus sandstones, the upper part of which contains inarticulate brachiopods, suggesting an approximate age ranging from Late Cambrian to Early Ordovician. An age of 595 + 43 Ma is shown in Figure 4 within the Taniagouri Group. Supergroup 3 includes, above an erosional and locally angular unconformity (Dia et al. 1969), a Late Ordovician glacial level overlain by graptolitic Silurian shales. This second erosional disconformity capped by glacial deposits occurred at the end of the Ordovician (Hirnantian). Supergroup 4 corresponds to Carboniferous deposits, which rest unconformably on Silurian shales or on Ordovician glacial deposits. In the southern part of the Taganet sub-basin (Assaba area), the lithostratigraphic sequence shows a very thick and complex basal sequence with several volcanic and glaciogenic levels associated with banded iron formations (BIFs), indicating a marine volcanogenic environment. An erosional unconformity occurs inside the Cambrian–Ordovician

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sandstones. However, the Silurian black shales are missing (Lepage 1983; Lafrance 1996). East of the MFT, in the Akjoujt area, the sedimentary covers are slightly folded, but in the southwestern part, along the Assaba cliffs, this foreland seems to be unaffected or very slightly affected by the Hercynian tectonic event. The Tambaoura sub-basin’s stratigraphy is not quite the same as that of the Taganet basin. The Hassabet-el-Hassianne Group and probably a large part of the Atar Group are missing, whereas the Bakoye Group (missing in the Taganet sub-basin) is widely represented (Deynoux et al. 1989). Supergroup 2 is limited to green argillites overlying the ‘Triad’, and Supergroups 3 and 4 are missing in this area. This basin has not been affected by the Hercynian tectonic event. The Fale´me´ trough (Fig. 5), studied by Bassot (1966), Chiron (1973), Lepage (1983), Dia (1984) and Lafrance (1996), fringes the western rim. It is partly (to the west) tectonized and metamorphosed. The lithostratigraphic sequence changes from the north to the south (Fig. 6, logs 1 and 2). Although the Fale´me´ trough is the westernmost part of the Taoudeni basin, it has a different structure and lithostratigraphy from that of the rest of the basin: it has a north– south orientation parallel to the margin of the WAC, and it does not contain any sediment from the Taoudeni basin’s Supergroup 1. It is related to the widespread tectonic event that occurred in West Africa, after the deposition of the last glacial Neoproterozoic sediments. The Early Cambrian sequence (Tichilit-al-Beida Group) is more complicated in the southern part, with a large amount of volcano-sedimentary formations (Nagara and Bouly Groups). Diamictites, laminites with lonestones, and turbiditic sandstones are generally present in the lower part of the sequence and are interpreted as glacio-marine deposits (e.g. Bassot 1966; Culver & Williams 1979; Villeneuve 1984; Culver & Magee 1987). The Late Cambrian succession, which consists of sandstones, limestones and red shales to the north (Mejeria Group), exhibits only red sandstones to the south (Ndoumeli Group). The Cambrian–Ordovician deposits consist of cross-bedded white sandstones to the south (Ndeio Group) and Scolithus-bearing sandstones to the north (Agouaoujeft Group). The

Fig. 3. (Continued ) MED, Moujeria; MBT, M’Bout; KD, Kidira. 1, Mesozoic– Cenozoic sedimentary rocks of coastal basins; 2, Carbonifereous deposits; 3, Palaeozoic deposits of the foreland; 4, Youkounkoun Group; 5, shales of the Late Neoproterozoic to Devonian in the foreland; 6, sediments of the Late Neoproterozoic to Devonian in the Mauritanide belt; 7, Late Neoproterozoic metamorphic rocks up to Devonian sediments in the western unit of the Mauritanide belt; 8, Bakoye Group; 9, calc-alkaline Pan-African granites in the Mauritanide belt; 10, ultrabasic rocks in the belt; 11, Late Proterozoic cover (Supergroup 1); 12, West African cratonic basement included in the Mauritanide belt; 13, West African craton basement; 14, thrusts, MFT, Mauritanide frontal thrust; 15, faults; 16, main folds.

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Fig. 4. Synthetic stratigraphic sequence of the Taoudeni Basin (after Villeneuve 2005, modified). 1, basement; 2, conglomerates and sandstones; 3, limestones; 4, shales and cherts; 5, cross-bedded sandstones; 6, sandstones with Scolithus; 7, glacial deposits; 8, pelites; 9, shales; 10, sandstones; 11, glacial conglomeratic shales; 12, stromatolites; 13, graptolites; 14, brachiopods; 15, microfossils; 16, erosional unconformity; 17, angular unconformity; 18, stratigraphic unconformity; Sg, Supergroup.


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Fig. 5. Geological sketch map of the central and southern parts of the Mauritanide belt and the Bassaride and Rokelide belts. Mkb, Madina–Kouta basin; Tgt, Taganet; Ass, Assaba; Tbb, Tambaoura; SMb, Senegalo-Mauritanian basin; Dbl.b, Diourbel basin; Yb, Youkounkoun basin; Tbb, Taban trough; RT, Richat trough; Ft, Fale´me´ trough; Kbt, Komba trough; Kct, Kolente´ trough; Rrt, Rokel River trough; Mt, Mali or Komba trough; Akj, Akjoujt; Nd, Nouadhibou; NK, Nouakchott; Dkr, Dakar; Cnk, Conakry; Frt, Freetown; M, Monrovia. 1, West African craton basement; 2, Pan-African belt; 3, Hercynian belt; 4, volcanic unit; 5, Pan-African reworked basement; 6, Late Proterozoic cover (Supergroup 1); 7, shales of the Late Neoproterozoic to Devonian; 8, Archaean to Birimian reworked basement; 9, Youkounkoun Group and equivalents; 10, Palaeozoic deposits of the foreland; 11, Cambrian –Ordovician; 12, Silurian– Devonian; 13, Mesozoic and Cenozoic deposits; 14, tillitic levels.

178 M. VILLENEUVE Fig. 6. Stratigraphic successions of the western basins and troughs. DR, erosional unconformity; DA, angular unconformity; Teb.gr., Tichilit al Beida Group; Mej.gr., Mejeria Group; Agh.gr., Agouaoujeft Group; DKL. gr., Dikkel Group; GNG.gr., Gneigara Group; Nga.gr., Nagara Group; BIF, banded iron formations; Boul.gr., Bouly Group; Ndoum. gr., Ndoumeli Group; Ndei.gr., Ndeio Group; SAKH.gr., Sahka Group; GOD.gr., Godiovol Group; TR.seq, Triade; Tich.gr., Tichilit al Beida Group; MK basin, Madina– Kouta Basin; SKT.gr., Soukouta Group; KB. gr., Komba Group; Pnp.gr., Panampou Group; Yk.gr., Youkounkoun Group; Pita gr., Pita Group; Telim.gr, Telimele Group;


Late Ordovician glaciogenic sediments (Dikkel and Sakha Groups) and the Devonian shelly sandstones (with brachiopods) exist in both parts, but the Early Silurian black shales are lacking in the south (Drot et al. 1978). This foreland shows gentle folds and reverse faults, such as the Massar Fault, interpreted by Lafrance (1996) as a normal fault linked to the extensive event of the Fale´me´ trough and reworked during the Hercynian tectonic event. The Youkounkoun Group (Fig. 6, log 4) fills a triangular and very deep (almost 1500 m) basin: the Youkounkoun basin. The Youkounkoun Group is slightly folded to the north by the Hercynian Mauritanide tectonic event. It consists of conglomerates and red coarse feldspathic sandstones. The direction of transportation indicates a centripetal deposition in sub-basins. The deposits of the Youkounkoun Group most probably have a Late Cambrian to Early Ordovician age. (3) The Bove´ basin groups. All the terranes previously described are unconformably overlain by the sediments of the Bove´ basin. The Bove´ basin covers the main part of the Fouta-Djalon massif. The contact between the Youkounkoun Group and the white sandstones consists of an erosional disconformity marked by a 1–2 m thick conglomeratic level. The Bove´ basin is almost flat in the Fouta –Djalon area but is deformed to the north (in Senegal and northern Guinea-Bissau) over a 30– 50 km wide area south of the HFT. According to Villeneuve (1984) and Villeneuve & Da Rocha Araujo (1984), the stratigraphic succession includes three groups; from the base to the top (Fig. 6, logs 4 and 5), these are the Pita Group, the Telime´le´ Group and the Bafata Group. The Pita Group (250– 930 m) includes three formations: from the base to the top, these are the Guemeta formation, the Kindia formation and the Mont Gangan formation. The Guemeta formation includes red sandstones and conglomerates, and represents the Rokelide molasses. The Kindia formation is mainly represented by white fluviatile and deltaic sandstones. The Mont Gangan formation includes conglomeratic argillites or sandstones, which were correlated to the Late Ordovician tillite (Villeneuve 1984). Fossils have not been found in any of these formations, except the latest Kindia levels, where Roman’ko (1974) found some Late Ordovician graptolites.

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The Telime´le´ Group (150 –330 m) includes black shales with interbedded grey laminated sandstones. Seven formations have been distinguished and dated by graptolites, brachiopods and microfossils (Villeneuve 1984). The Bafata Group (150– 430 m) is an alternation of sandstones and shales with ripple-mark structures indicating a deltaic environment (Villeneuve 1984). The lower formation contains Gedinian brachiopods (Villeneuve 1984) whereas the upper one contains Famennian brachiopods (Bechennec 1980). Thus we conclude that the lithostratigraphy observed in the foreland of the orogenic thrust belt (MFT) fringing the western margin of the WAC is far from uniform. In particular, significant differences are noticed within the reported lithostratigraphic sequences when comparing the sedimentary infill of the cratonic basins (Taoudeni basin) with that of the adjacent pericratonic basin such as the Fale´me´ trough. This discrepancy has led to debate between geologists working respectively on the foreland and on the MFT belt itself. We point out two main differences: (1) a large amounts of volcanic rocks are found within the pericratonic troughs, as a result of the considerable influence of the extensional tectonic regime prevailing during their sedimentary infill; (2) the base of these troughs is not always filled with glaciogenic sediments like those characterizing the Supergroup 2 deposits found within the cratonic basins. The Mauritanide thrust belt. West of the MFT, the structures are very complicated, with a mixture of metamorphosed and non-metamorphosed complexes. Therefore, each section requires specific study. We describe here seven cross-sections, which provide an overall insight into the geological structure along profiles (see location in Fig. 3) through the Mauritanide belt in various areas from north to south. These cross-sections are shown in Figure 7. (1) The Anti-Atlas section. West of the MFT, the formations affected by the Hercynian tectonic regime belong to the Late Neoproterozoic or Cambrian (Hoepffner et al. 2005). No specific metamorphism has been detected in this section. (2) The Dhlou –Sekkem section (profile 1, Fig. 7). This is located in the southwestern part of the Tindouf basin, west of the city of Smara. It has been investigated by Dacheux (1967) and

Fig. 6. (Continued ) Bafat.gr., Bafata Group; Kolo.gr., Kolente Group; Souti gr., Mont Souti (¼ Guemeta formation) group; Tab.fm., Tabe formation; Mak. fm., Makani formation; Tey.fm, Teye formation; Mbl. Fm, Mabole formation; KWH. Fm, Kasewe hills formation; Tala fm, Taia formation. 1, shales; 2, sandstones; 3, crossbedded sandstones; 4, pelites; 5, conglomeratic shales; 6, limestones; 7, cherts, jaspers; 8, tillites, glacial deposits; 9, basic volcanic formations; 10, rhyolitic formations; 11, sandstones with Scolithus; 12, metamorphic basement; 13, brachiopods; 14, graptolites; 15, stromatolites; 16, microfossils.

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Fig. 7. Geological cross-sections in the Mauritanide belt (location of the profiles shown in Fig. 3). Legend: 1, Mesozoic–Cenozoic of coastal basins; 2, Carboniferous and Upper Devonian deposits; 3, Silurian–Devonian; 4, Cambrian –Ordovician; 5, Youkounkoun Group in the foreland; 6, Youkounkoun Group in the folded belt; 7, Late Neoproterozoic and Cambrian in the folded belt; 8, tillites or mixtites; 9, Late Neoproterozoic and Cambrian in the foreland and in the para-autochthonous units; 10, Oumachoueima Group of the Akjoujt area; 11, metamorphosed Late Neoproterozoic to Devonian in the western units of the Mauritanides (Wa-Wa group); 12, Late Proterozoic cover (Supergroup 1); 13, Termesse group; 14, intrusions of calc-alkaline Pan-African granites in the Mauritanides belt; 15, basic and ultrabasic rocks of the belts (Guinguan Group and equivalents); 16, calc-alkaline Pan-African basement; 17, West African craton basement; 18, thrust; 19, HFT, Hercynian front thrust; 20, HML, high metamorphism limit.


Ratschiller (1970). It shows the Mauritanian frontal thrust (MFT) separating the Zemmour syncline (eastern unit) from the ‘Dhlou belt’ (central and western units). The lower part of the sedimentary succession corresponds to the ‘El Tlethyate’ limestones that crop out in the western unit and was correlated (Sougy 1969) to the Supergroup 1 deposits of the Taoudeni basin. Cambrian– Ordovician shales and sandstones overlap these ‘El Tlethyate’ stromatolitic limestones and also the metamorphic Reguibat basement. These Cambro-Ordovician sediments are capped by Upper Ordovician ‘tillites’ and Silurian and Devonian fossiliferous deposits. (3) The Adrar Souttouf section (profile 2, Fig. 7). This section is located in the western part of the Reguibat uplift. Cross-section 2 (Fig. 7) displays four main allochthonous units thrust over the Palaeozoic sedimentary cover, which remained attached to its Archaean basement. The main thrust plane separating the metamorphic units from this Palaeozoic sedimentary cover deposited within the foreland area belongs to the MFT structural framework. West of the MFT we identify, from east to west, the following units (Villeneuve et al. 2006). (a) The Matallah unit is, according to Villeneuve et al. (2006) a part of the Archaean Reguibate uplift basement capped with a very low-grade metamorphic complex of unknown age (around 500 Ma?). (b) The Dayet Lawda unit comprises amphibolitic, gabbroic and basaltic rocks dated by the K/Ar method on whole-rock samples and mineral separates (Villeneuve et al. 2006). Two measurements yield ages around 1200 and 1050 Ma (whole rocks and minerals), and three other experiments provide ages close to 500 Ma. A 733 Ma PanAfrican age is also recorded on a non-deformed basaltic sample (Villeneuve et al. 2006). (c) The Gezmayet unit contains metamorphic (mylonites, gneisses, micaschists) and granitic rocks. K –Ar dating results yield apparent ages ranging from 670 to 633 Ma, which probably relate to the age of the metamorphic episode that affected these rocks. An eclogite from Tasiast (southern Mauritanian part of the Gezmayet unit) dated by Le Goff et al. (2001) at 590 Ma (U –Pb on zircon) and reworked at 330 Ma (Sm–Nd method) provides another temporal constraint for the metamorphism that affected the Mauritanide belt. (d) The Fadrat al Garod unit has not yet been dated. It contains metamorphic rocks (micaschists, gneiss, mylonites and deformed granites) together with dykes of rhyolitic lava. In contrast to the previous interpretation, the Adrar Souttouf belt does not correspond to a Hercynian klippe (Dallmeyer & Lecorche´ 1991) but to

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a pile of different units stacked on top of each other, from east to west, as a result of a compressional regime induced during the Pan-African and Hercynian orogens. However, the evidence of ‘Grenvillian’ ages (1200 and 1050 Ma) obtained on magmatic and metamorphic rocks in the Mauritanide belt is reported for the first time in West Africa. This issue requires further investigation. (4) The Akjoujt Section (profile 3, Fig. 7). This section is one of the best studied, because of the presence of a copper mine, but it is also one of the most complicated and debated section. First investigations by Tessier et al. (1961), Giraudon (1963), Giraudon & Sougy (1963), Michaud (1964), Rippert (1973) and Lecorche´ (1980) allowed the identification of several nappes thrust over the Palaeozoic sedimentary cover, specially in Guelb el Hadej (Fig. 7, cross-section 3). According to Lecorche´ (1980), at least five thrust sheets have been identified; from the top to the base, these are as follows. (a) A metamorphosed volcano-sedimentary thrust sheet (Agualilet unit), which consists of two units: a quartzitic unit and a volcanic unit with basalts, gabbros, prasinites and silicic tuffs. (b) A granitic thrust sheet (Hajar Dekhem and Kleouat Massifs), which contains migmatitic and porphyroid granites. (c) A metamorphosed volcano-sedimentary thrust sheet called ‘Se´rie d’Akjoujt’ with two volcanic series (upper and lower) separated by a schistoquartzitic formation (Michaud 1964). According to dating results presented by Dallmeyer & Lecorche´ (1990b) and Clauer et al. (1991) the ‘Se´rie d’Akjoujt’ yields radiometric ages ranging between 320 and 301 Ma. Recently, Martyn & Strickland (2004) have identified a disconformity between the two volcano-sedimentary sequences. (d) A quartzitic thrust sheet called the ‘Nappe des quartzites’ composed of quartz–sericite- and muscovite-bearing quartzites and sericite-bearing quartz-schists. (e) A sedimentary thrust sheet compositionally very similar to the sedimentary cover sequences in the foreland area but comprising volcanic material and called the ‘Unite´ des regs’. This unit contains various rock elements also found within the Cambro-Ordovician deposits in the foreland area, but the original depositional sequence has been totally dismembered. New observations by Martyn & Strickland (2004) led to a new hypothesis in favour of a model that considers the existence of a Pan-African I ophiolitic unit that occurs beneath the Cambro-Ordovician series. However, many questions remain unsolved in this area, particularly with respect to the origin of the materials observed in the quartzite thrust sheet.

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(5) The Aouker–Kidira section (profiles 4 and 5, Fig. 7). This was investigated by Lille (1967), Chiron (1973), Lepage (1983), Dia (1984), Remy (1987), Ould Souelim (1990), Lafrance (1996) and Diop (1996). The contact between the paraautochthonous deposits and the foreland deposits (i.e. the MFT) has been mapped at several locations close to the Assaba Cliff, by Lafrance et al. (1993). Cross-sections 4 and 5 (Fig. 7) show several units from the eastern foreland to the western ‘Mounts Wa-Wa’ unit. According to Lepage (1983), we can distinguish, from east to west, the following units. (a) The ‘Anietir unit’ (probaly a paraautochthonous unit) comprises more or less identical material to that found in the foreland area but includes some volcanic layers generally interbedded within the basal glacial sequence. Such volcanic layers (basaltic and rhyolitic rocks), have been studied by Lafrance (1996) in the Massar area. BIF and glacial conglomerates correlated with green formations including tillites, shales and siltstones occur there. D. Lahondere (pers. comm.) obtained two U –Pb zircon ages of 635 and 610 Ma on volcanic rocks belonging to this unit. (b) The ‘Gadel unit’ is an ophiolitic and metamorphic complex with quartzites and amphibolites. This complex is linked to the Pan-African I oceanic stage. Despite several attempts to date these volcanic rocks, no consistent age has been obtained. (c) The ‘M’Bout unit’ contains calc-alkaline granites and metamorphic and sedimentary series, which include quartzites as well as rhyolitic tuffs found in the Toukobra zone. The origin of this metamorphic complex can be related to the PanAfrican I active margin and the uppermost volcanosedimentary layers may be correlated to the Palaeozoic sedimentary cover. According to Choubert & Faure-Muret (1971), datings by the Rb/Sr method on the Guidimakha granite provided several ages between 800 and 620 Ma. (d) The ‘Mount Wa-Wa unit’ (suprastructural allochthonous) includes several formations and complexes, including the ‘Wa-Wa quartzites’. This unit also contains prasinites, sericite schists and micaschists. The Palaeozoic sedimentary cover is strongly suspected to be the source rock of this internal metamorphic unit. Dallmeyer & Lecorche´ (1990a) reported several 40Ar/39Ar ages on muscovites from the Wa–Wa quartzites or schists, ranging between 312 and 267 + 2 Ma. These ages were interpreted as the Hercynian tectonothermal imprint, but U/Pb dating on zircons from the Wa-Wa quarzites gives an age around 600 Ma (D. Lahondere, pers. comm.). Therefore, we conclude that the Hercynian tectonic event strongly affected both the Neoproterozoic basement and the Palaeozoic deposits.

(6) The Koulountou–Bove´ section (profiles 6 and 7, Fig. 7). This area, studied mainly by Bassot (1966) in Senegal and Bechennec (1980) in Guinea-Bissau, has also been investigated by Villeneuve (1984), mainly in Senegal and GuineaBissau. The southern part of the Mauritanides is structurally different from the northern part (crosssections 6 and 7, Fig. 7). Here, the Hercynian and Pan-African belts are geographically separated. The MFT now crops out on the eastern side of the belt whereas the main part of the Hercynian Mauritanide belt is concealed underneath the MesoCenozoic Senegalo-Mauritanian basin. The Hercynian belt is turning to the west, the Pan-African belt on the other hand extends to the south. West of the HFT, the Koulountou group, which includes several formations, is very similar to the ‘Mount Wa-Wa’ unit and partly to the M’Bout unit. The Hercynian slaty cleavage (dipping west) is stronger to the west. The ‘Koulountou unit’ consists of rhyolitic tuffs, sericite schists, micaschists, gneisses and mylonitized granites. The mylonitized schists within the Simenti Formation yielded several radiometric ages between 270 and 280 + 0.8 Ma (Dallmeyer & Villeneuve 1987). Discussion and interpretation. The discussion presented here on the major tectonostratigraphic elements that are found within the Mauritanide orogenic belt takes into account several elements retained from numerous structural, metamorphic, geochemical, geochronological, geophysical and geodynamical studies. From the east to the west, we distinguish three main parts of the belt. (1) An eastern volcano-sedimentary paraautochthonous part very similar to the sedimentary series found in the Taoudeni basin but with more abundant volcanic rocks in the lowest part of the sequence, which rests upon the West African basement. (2) A central volcanoclastic to granitic part, comprising volcanic formations and related to an ophiolitic dismembered environment, and a granite-gneiss complex with large amounts of rhyolitic or dacitic tuffs. In these units two volcanic formations have been distinguished, basic and calc-alkaline. (3) A western allochthonous part, which corresponds to a metamorphic unit including gabbros, gneissic and quartzitic rocks. It corresponds to the ‘Agualilet’ and ‘Mount Wa-Wa’ units. However, the Adrar Souttouf section does not correspond exactly to the central Mauritanides sections, because of the paucity of the sedimentary cover in the inner belt and the large amount of basic formations in the central unit.


Tectonic and metamorphic characteristics may be locally complicated, but in general the basement recorded two or more tectonic events (with different cleavages) whereas the Neoproterozoic to Palaeozoic formations underwent only one tectonic phase revealed by only one strong cleavage. The thrusting seems to be generally post-cleavage. In the tectonic thrust sheets comprising mainly Pan-African I basement rocks, the quartzites and micaschists exhibit a paragenesis with staurolite and garnets (Dia (1984) found an eclogitic facies in pyroxenites and amphibole-pyroxenites), but in the tectonic thrust sheets composed of Neoproterozoic or Palaeozoic rocks the metamorphic parageneses remain limited to the greenschist facies. However, in the westernmost ‘Mount Wa-Wa’ unit (central Mauritanides), where quartzites and schists exhibit a paragenesis with staurolite and garnets, or in the Gezmayet unit (northern Mauritanides), where an eclogitic facies related to the Hercynian tectonic event was found in the volcanic formations (Le Goff et al. 2001), the Hercynian metamorphism is more intensive. Therefore, we consider that two main metamorphic domains exist in the Palaeozoic formations: an eastern domain where the metamorphism seems to be moderate and a western one where it seems more intensive. The limit between the two domains is called the HML (high metamorphism line). Only two sets of magmatic rocks have been studied: the ‘ophiolitic assemblage’ and the ‘granitic batholith’ of the central Mauritanides. The ophiolitic assemblage exhibits a mid-ocean ridge basalt (MORB) affinity but Remy (1987) considered a ‘Red Sea type’ ophiolitic assemblage for the central Mauritanide complex of Selibaby. Basalts from the ‘Akjoujt series’ present an island arc volcanic affinity (Kessler 1986; Pouclet et al. 1987). The central Mauritanide granitic batholith (M’Bout unit) is comparable with calc-alkaline active margin magmatism. However, the large amount of rhyolitic tuffs associated with this unit suggests a continent–continent collisional environment. Geochronological data are rather few in number and precarious. Lille (1967), Bassot et al. (1963), Bonhomme & Bertrand-Sarfati (1982) presented some results using both the K/Ar and Rb/Sr method. Clauer et al. (1982) dated the sediments using the K/Ar method and Dallmeyer & Villeneuve (1987) and Dallmeyer & Lecorche´ (1989) dated the metamorphic formations using the 40Ar/39Ar method. Other radiometric data were provided by Blanc et al. (1986) and Le Goff et al. (2001) on zircons (U/Pb method). Based on geochronological data four main geological events can be retained: (1) 680– 620 Ma (first Pan-African orogen); (2) 600– 590 Ma

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(opening of troughs in the western margin of the WAC); (3) 530–500 Ma (metamorphic event and gabbroic intrusions in the Adrar Souttouf are a); (4) 330–270 Ma (Hercynian orogeny) with two peaks: one around 330–300 Ma and another one around 280–260 Ma. A metamorphic event recorded ages of 1200– 1000 Ma (Grenvillian age) in the Adrar Souttouf (northern part of the Mauritanides). However, the Taconian event (440 Ma) supported by Lecorche´ (1980) in the central Mauritanides is not recorded by any geochronological data. Very few geophysical data are available for this part of West Africa. Gravimetric investigations were performed in Mauritania and Senegal by Creen & Rechenman (1965) and in Bissau by Amorin-Ferreira (1966). The gravimetric scheme (Fig. 8a) shows a large north –south ‘Bouguer anomaly’ parallel to the central Mauritanide belt. It was interpreted by Guetat (1981) as a ‘mantle uplift’ located c. 160 km (Fig. 8b) to the west of the HFT. Guetat et al. (1982) connected this huge gravimetric anomaly to a fossil subduction zone. The main question is to decide which is the orogenic event related to this anomaly. Other geophysical investigations using magnetotelluric methods (Ritz & Robineau 1986) or seismological investigations (Dorbath et al. 1983) also argued for a suture but did not provide any further constraint on the age of the incipient orogen. Beyond the local interpretations and the divergent point of views, we consider, at this stage of knowledge, that the Mauritanide belt is a Hercynian tectonic belt that includes metamorphosed sediments belonging to Neoproterozoic and Palaeozoic volcanoclastic formations that infilled several north –south-trending troughs and basins and also included parts of the basement. This basement was incorporated in the belt by the strong Hercynian tectonic phases around 300 and 270 Ma. This simplified geodynamical model explains the occurrence of Middle Precambrian and Pan-African terranes covered with Neoproterozoic or Palaeozoic sediments. Diop (1996) argued for a ‘thin skin tectonic model’. However, an island arc environment was suspected by Kessler (1986) and Pouclet et al. (1987), and suggests another scenario. Conclusions. Because of the large variety of formations observed along the various cross-sections through the Mauritanide belt and because of the different tectonic events that affected these different parts of the belt asynchronously, this area remains the most disputed in West Africa. Our model includes all previous interpretations and reconciles those researchers favouring a unique Neoproterozoic to Palaeozoic basin tectonized during the Hercynian tectonic phases (e.g. Sougy, Lecorche´,

184 M. VILLENEUVE Fig. 8. Gravimetric data for the western fold belt. (a) Gravimetric scheme (after Villeneuve et al. 1990, modified). Legend: 1, zero Bouguer contour line; 2, negative contour line; 3, positive contour line; 4, main positive anomaly; 5, gravimetric sector boundaries; 6, number of gravity sectors (1, West-African crustal block; 2, Senegalese block; 3, Bassaride block; 4, Rokelide block). (b) Gravimetric profile across the central Mauritanide belt and interpretation of the gravimetric anomaly (after Guetat et al. 1982, modified).


Lepage) and the supporters of a Pan-African belt included in the Hercynian belt (e.g. Chiron, Dia, Villeneuve).

The Bassaride belt Introduction. The Bassaride belt, located in the southernmost part of the Mauritanide belt (Fig. 5), has been considered by Villeneuve (1984) as different from the Mauritanide and the Rokelide belts. Previously, Sougy (1962a) and Bassot (1969) correlated it with the Rokelide belt in Sierra Leone. This part of the West African belt has long been controversial despite reliable radiometric data (Dallmeyer & Villeneuve 1987). Unfortunately, only a small part of this belt is exposed; the main part is concealed underneath the Palaeozoic cover, so that the contact with its foreland is never observed. The sedimentary rocks covering the Bassaride belt crop out mainly in the southern part of the Fale´me´ trough, in the Komba trough filled by sediments belonging to the Mali Group (at the base of Supergroup 2), in the Youkounkoun basin (middle part of Supergroup 2) and in the Bove basin (Supergroups 2, 3 and 4). Previous work. Bassot (1966, 1969) considered that the Bassaride formations were affected by the ‘Caledonide’ orogen. He distinguished a western ‘Koulountou Branch’ (considered as the southern end of the Hercynian Mauritanide belt) and an eastern ‘Bassaris Branch’, separated by the Youkounkoun basin. The Soviet–Guinean teams that produced the geological map (1:200 000) for northern Guinea-Bissau (Torchine 1976a, b) assigned the ‘Bassaris ridges’ to the Birimian basement cropping out in the Kedougou inlier. Villeneuve (1982, 1984) identified a major unconformity (the Nadiebary disconformity) between the two groups (Termesse and Guinguan Groups) of the ‘Bassaris ridges’ and the overlying Mali Group. He suggested a Neoproterozoic age for the two Bassaride groups and correlated the overlying Mali Group (Neoproterozoic to Cambrian) to the earlier defined Supergroup 2 of the Taoudeni basin. An 40Ar/39Ar dating on micaschists belonging to the Guinguan Group gave a Neoproterozoic age of 660 + 0.7 Ma (Dallmeyer & Villeneuve 1987). Thus, Villeneuve & Dallmeyer (1987) considered that the deposits belonging to the Guinguan Group were affected by the Pan-African I event. The presence in these two groups of a large amount of volcanic rocks, whereas the overlying Mali Group is devoid of volcanic rocks in this area (Villeneuve et al. 1991), reinforces this interpretation. However, Brinckman et al. (2003), who worked on the Bassaris ridges, correlated the Guinguan Group to the Birimian basement and the Termesse Group to the Neoproterozoic and Palaeozoic cover (as we did for the Mali Group). A new

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interpretation, put forward in this paper, considers the Guinguan Group as representing the main part of the lithostratigraphic sequence affected by the Pan-African I orogen. In this model, the Termesse Group is considered to be coeval with the Bakoye Group (top of Supergroup 1 in western Mali) and related groups. In this new hypothesis the Bassaride belt’s stratigraphy is restricted to two major groups, the eastern ‘Guinguan Group’ and the western ‘Niokolo Koba Group’. The lithostratigraphic succession of this area is shown in Figure 9. The two groups belonging to

Fig. 9. Lithostratigraphic succession in the Bassaride belt (after Villeneuve et al. 1991, modified). 1, sandstones and shales of the Bafata Group; 2, shales of the Telime´le´ Group; 3, cross-bedded sandstones; 4, sandstones and conglomerates of the Youkounkoun Group; 5, argillites and shales of the Mali, Oussekiba and Batapa Groups; 6, tillites and mixtites; 7, upper part of the Termesse Group; 8, Termesse Group; 9, Kemberra sandstones; 10, Niokolo–Koba Group; 11, Guinguan Group; 12, siltsones and sandstones of the Madina– Kouta basin; 13, shales of the Madina–Kouta basin; 14, crystalline basement; D, disconformity. Bft.g, Bafata Group; Tlg, Telime´le Group; Pt.g, Pita Group; Ygr, Youkounkoun Group; MGr., Mali Group; T.gr., Termesse Group; NKbgr, Niokolo– Koba Group; G.g, Guinguan Group, MKgr., Madina–Kouta Group; SGR, Segou Group; BS, Basement; Eb orog, Eburnian (Birimian) orogen; Bass orog., Bassaride orogen (Pan-African 1); Rok.orog., Rokelide orogen (Pan-African II); Herc.orog., Hercynian orogen.

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the Bassaride belt are capped successively by the Termesse Group, the Mali Group, the Youkounkoun Group and finally by the Palaeozoic Bove´ Basin Group. These groups are all separated by unconformities (D3 –D7). The Bassaride frontal thrust. A schematic crosssection of the Bassaride belt is shown in Figure 7 (cross-section 6). Unfortunately, the supposed ‘Bassaride frontal thrust’ does not crop out in this area. The contact between the eastern Guinguan Group and its contemporary foreland (the Birimian Kedougou basement and the Madina –Kouta basin sediments) is concealed underneath the Termesse and Mali Groups. The contact between the Guinguan and the Niokolo–Koba Groups is also concealed underneath the Cambrian –Ordovician sediments of the Youkounkoun basin. The Bassaride foreland. The contemporary foreland consists of the Kedougou basement (Birimian, c. 2000 –1900 Ma) and the sedimentary Madina– Kouta basin. This basin is located in a ‘depression’ between the Kedougou inlier and the Leo uplift. According to Villeneuve (1989), the lithostratigraphic succession consists of two groups: the Segou Group (500 m) at the base and the Madina–Kouta Group (700 m) on top. The Bassaride thrust belt. Following our new hypothesis correlating the Termesse Group to a more recent orogenic event, we are here considering only two groups: the Guinguan Group and the Niokolo-Koba Group. The Guinguan Group. The Guinguan Group crops out mainly in the north–south elongated Bassaris ridge (SE Senegal and northern GuineaBissau). We notice a thrust sheet belonging to the ‘Mali Group’ wedged between the Guinguan and Termesse Groups. The cross-section shows a gently folded belt but detailed tectonic investigations show two superimposed tectonic phases. The first phase (P1) produces recumbent folds with an axial schistosity, whereas the second phase (P2) is expressed as open folds with a brittle schistosity. Villeneuve (1984) distinguished four formations consisting of amphibolites, micaschists, chlorite-schists, prasinites, metabasalts, metadolerites, serpentinites, red jaspers, quartz –biotite and quartzophyllitic rocks. According to Dupont et al. (1984) these serpentinites would be derived from the alteration of peridotites. Grade of metamorphism ranges from greenschist facies to amphibolite facies with the presence of synkinematic green hornblende, epidote and garnet. The Niokolo Koba Group. This group corresponds to the eastern part of the Koulountou branch. Bassot (1966) and Villeneuve (1984)

distinguished three main types of rocks: granitoids, and basic and acidic lavas. The granitic rocks constitute a broad batholith (the Linki –Kountou massif ) and some small windows along the Bollore´ River. The basic lavas consist of aluminous basalts and andesite, and the acidic lavas consist of dacites, rhyodacites, rhyolites and ignimbrites associated with grey, red and green jaspers, breccias and conglomerates. The Niokolo– Koba Group shows gentle folding (NE–SW axis). These open folds are hectometre to kilometre scale with a limb of only 20 –408. The cleavage is usually of fracture type except along shear zones, which are 20 –50 cm wide and dip at 50 –708 to the west. Rb/Sr datings (Bassot & Vachette 1983) provided an age of 683 + 18 Ma for the batholith emplacement. However, rhyolites and tuffs are probably more recent than 645 Ma (Rb –Sr age on biotite in the Linki–Kountou granite) according to Bassot et al. (1963). Consequently, a large part of the basic and acidic lavas intruding the Linki–Koutou massif must be related to a post-Bassaride activity. The Niokolo–Koba Group is disconformably overlain by the ‘Oussekiba Group’ (correlated with the Mali Group) and by the Youkounkoun sandstones. Formations similar to the Niokolo –Koba Group (‘Panampou Group’) crop out in several windows piercing the Youkounkoun basin in northern Guinea-Bissau. The ‘Panampou Group’ is unconformably covered by the ‘Loumbaloumbito conglomerate’ correlated with the basal Mali Tillite, and by green shales of the ‘Batapa Group’ correlated with the Mali Group. Finally, the red sandstones of the ‘Youkounkoun Group’ cap this area. The Telimele window has a granitic basement dated at 755 + 20 Ma and a mylonitic formation dated at 534 + 10 Ma by the K/Ar method (Seliverstov 1970). The sedimentary cover. Four volcanosedimentary or sedimentary formations cover the ‘Bassarides belt’; from the base to the top these are (Fig. 6, logs 3 and 4): (1) the lower volcanosedimentary groups including the Termesse and Panampou Groups and probably the upper part of the Niokolo– Koba Group; (2) the upper sedimentary groups including the Mali, Oussekiba and Batapa Groups; (3) the Youkounkoun Group; (4) the Bove´ basin groups including the Pita, Telimele and Bafata Groups. The Termesse Group probably caps the Madina Kouta basin. The contact between the Termesse Group and the Birimian Kedougou inlier is mapped near Soukouta, but the contact with the Madina –Kouta basin is concealed underneath the sediments of the Mali Group along the Fale´me´ and Komba troughs. The Termesse Group consists of volcanic tuffs, silexites, greywackes, conglomerates of possible glacial origin (mixtites), basalts,


dolomitic carbonates, (red, green and blue) jaspers, shales and basaltic lava flows interbedded with the volcano-sedimentary deposits. East of Termesse, a large (1– 5 km wide) basaltic (with vertical breccias) formation called the ‘Koubia formation’ intrudes the volcano-sedimentary formations of this Termesse Group. Except for the Koubia formation, the Termesse rocks are strongly folded, but are not metamorphosed. According to Bassot (1966), the sedimentary environment of the Termesse Group corresponds to a passive margin. The upper part of the Niokolo –Koba and the Panampou Groups, consisting of pelites, tuffs, red and green jaspers, dacites, rhyolites and felsic rocks, and even aluminous basalts and andesites, are now classified within the Termesse Group. Bassot & Vachette (1983) assumed these rocks to be differentiated from the Linki –Kountou massif. They are grouped into the Niokolo–Koba volcanics series. The Termesse Group contains more basaltic lavas and the Panampou and Niokolo–Koba Groups contain more acidic lavas such as red rhyolites. The Mali and Batapa Groups overlie the Termesse Group after an unconformity that has been observed in many places, such as at Soukouta (Senegal) by Peronne (1967) and at Nadiebary (Guinea-Bissau) by Villeneuve (1982). This group starts with the tillite of Walidiala, found in many places either in Senegal or in the Fouta-Djalon massif in GuineaBissau (Villeneuve 1984). These basal levels are capped by at least 1000 m of shales, pelites, sandstones and several levels of carbonates. Until now, no volcanic rocks or traces of metamorphism have been found associated with the Mali Group. The folding of the units is very slight and limited to the western part of the Fale´me´ and Komba basins. Early Cambrian microfossils have been found within the ‘cap carbonates’ and within the Walidiala Triad boulders (Culver et al. 1996). However, an age of 600 Ma has been given by Deynoux et al. (2006) for the basal levels of these groups as the average age of this level for the entire Taoudeni basin. Discussion and interpretations. The geochemistry of basalts belonging to the the Guinguan Group (Dupont et al. 1984) indicates a mid-ocean ridge basalt (MORB)-type tholeitic affinity. The serpentinites must be considered as altered peridotites. The Linki–Kountou granite and the Lower ‘Niokolo– Koba Group’ show a calc-alkaline affinity according to Angeli (1983) and Bassot & Caen-Vachette (1983). The volcanic rocks of the ‘Termesse Group’ (Dupont et al. 1984; Villeneuve 1984) also display a MORB affinity but the presence of titano-augite basalts with an alkaline affinity indicates a continental influence. The Upper ‘Niokolo–Koba Group’ has a clear calc-alkaline

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affinity (Angeli 1983). The presence of high K2O value (2.4) suggests a thick continental crust below (Villeneuve et al. 1991). Geophysical investigations have been conducted by Akmetjanov et al. (1976), Ritz (1982), Dorbath et al. (1983) and Ponsard (1982). All indicated a suture between two crustal blocks: the eastern West African block and a western unnamed block. According to Ponsard et al. (1982), these two blocks have different densities and a heavy body is located between them. They interpreted this body (situated some 80–100 km west of the Guinguan Group outcrop) as a relict of an oceanic floor. Very few datings have been performed in this area. The available chronological data can be grouped into four time-frames. The oldest ages, between 1000 and 1050 Ma, have been obtained on rhyolites stratigraphically below the first deposits of the Madina –Kouta basin. A range of ages between 650 and 700 Ma has been measured on samples of the Niokolo–Koba granitic intrusions. These ages may date the metamorphism that affected the Guinguan Group. Three dating results obtained on samples from the Guinguan Group indicate a metamorphism around 660 Ma (656 –666 Ma), which was considered by Dallmeyer & Villeneuve (1987) as the true age of the Bassaride orogeny. Ages ranging between 650 and 500 Ma correspond to the volcanic activity of the upper Niokolo–Koba Group and to the Rokelide metamorphism. Ages between 250 and 300 Ma correspond to the Hercynian metamorphism in the groups of the Koulountou branch (Southern Mauritanide belt). Two hypotheses may explain this collisional model: (1) the opening and closure of a ‘Red Sea type’ basin on the western edge of the WAC (with a west-dipping slab); (2) a ‘back-arc basin’ with an east-dipping slab. Conclusion. The Bassaride belt presents the characteristics of a continental collisional belt. Geophysical data indicate a suture between two blocks of different density corresponding to the amalgamation of a western block with the WAC. The metamorphism related to this collision was dated around 660 Ma and this collision was followed by an intense post-collisional volcanism (basalts, rhyolites, dacites and ignimbrites of the upper Niokolo –Koba and related groups) and by the opening of troughs (Termesse Group). The Bassaride belt forms the basement of the Mauritanide and Rokelide belts.

The Rokelide belt Introduction. The Rokelide belt corresponds to a NNW –SSE folded zone that extends from the southern part of the Bove´ basin to the southern

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part of Liberia, through the western part of Sierra Leone (Fig. 5). It faces the western edge of the Leo uplift and cuts the NNE–SSW Bassaride belt. Both belts are overlain by the Bove´ basin sediments. The name ‘Rokelides’ was proposed by Allen (1967), who described a belt deformed at about 550 Ma, during the Pan-African tectonothermal event (Kennedy 1964). This belt was studied separately in each country: by the French and Soviet–Guinean teams in southern Guinea-Bissau, by British and Sierra Leonian geologists in Sierra Leone, and by a USA–Liberian team in Liberia. As in our discussion on the geology of the Mauritanide and Bassaride belts, we distinguish a foreland and a thrust belt separated by the Rokelide frontal thrust (RFT). All are overlain by the Bove basin sediments. Previous work. In Guinea-Bissau, the most recent studies were undertaken by Boufeev (1968), Torchine (1969) and Villeneuve (1984). Datings of the post-orogenic granite of Coyah (which intruded the Forecariah complex) presented by Dallmeyer et al. (1987) yielded ages of 530 Ma. Whereas Delor et al. (2002) provided three U –Pb zircon ages around 580 –550 Ma for the Forecariah granulitic gneisses. In Sierra Leone, Andrew-Jones (1966, 1968) and Allen (1966, 1967, 1968, 1969) carried out geological investigations on the Rokelide belt. The sedimentary Rockel River Group was studied by Reid & Tucker (1972), Culver et al. (1978, 1980), Culver & Williams (1979) and by Latiff et al. (1997), who clearly distinguished the undated Marampa Group from the Archaean basement. Bonvalot et al. (1991) provided the first gravimetric map of Sierra Leone. In Liberia, the Liberian Geological Survey, in co-peration with the US Geological Survey, produced 10 new geological maps between 1976 and 1983 summarized in the 1/1 000 000 geological map of Liberia (Tysdal & Thorman 1983). Behrendt & Wotorson (1970) published the first aeromagnetic and gravity maps of western Liberia. The Rokelide frontal thrust (RFT). The location of the Rokelide frontal thrust depends on the geological interpretation of the Rokelide belt. Previously, Torchine (1969) considered the eastern part of the western basement as a metamorphic equivalent of the Kolente sedimentary formations (equivalent of the Rokel River formations) and consequently considered thrusting of the entire western basement over the Kolente formations. However, Villeneuve (1981) did not agree with this interpetation. On the other hand, Williams & Culver (1988) and Culver et al. (1991) proposed thrusting of the entire western metamorphic belt over the Rokel River trough. For Culver et al. (1991), the Rokelide metamorphism involves the western part of the

Rokel River troughs. However, further investigations by Latiff et al. (1997) did not support this hypothesis. Latiff et al. (1997) considered a reverse fault (but dipping to the east) between the western basement and the Rokel River formations. To the south, in Liberia, the situation of the RFT is well defined. For Thorman (1976) and Tysdal & Thorman (1983), the contact between the Rokelide granulitic zone and the Archaean WAC corresponds to the ‘Todi shear zone’. The granulitic nappe thrust over remnants of the Rokel River sediments, 30 km to the west of the Todi shear zone, represents the maximum thrusting of the Rokelide inner belt. To conclude, we consider the RFT as a NNW–SSW limit represented by the ‘Todi shear zone’ in Liberia, the ‘Mylonite zone’ in Sierra Leone and the ‘Kolakoure´ fault’ in Guinea-Bissau. The Rokelide foreland. South of the Fouta –Djalon massif, the foreland corresponds to the West African crystalline basement (Leo uplift) and comprises belts and basins cropping out on the eastern side of the ‘RFT’. The Neoproterozoic basins and troughs. The sedimentary Neoproterozoic to Palaeozoic formations are located in several basins or troughs from Guinea-Bissau to Liberia, as follows. The Kolente´ basin (Fig. 6, log 5) is very similar to the Komba basin. From the base to the top we see the following formations (Villeneuve 1981, 1984): (1) a basal conglomerate with glaciogenic elements, which was previously ascribed to the Late Ordovician glacial event by Reid & Tucker (1972) in the Sayona-scarp Mountain; (2) a green and red argillitic formation with several levels of carbonates; a thick sandstone (200 m) formation with several conglomeratic beds; (3) a black and green shaly formation with conglomeratic levels (glaciogenic aspect); (4) a volcano-sedimentary formation called the ‘Bania formation’ consisting of spilites, breccias, basaltic lavas with argillites and interbedded jasper levels. It is folded and perhaps metamorphosed on its western side. The Rokel River trough (Fig. 6, log 6) which crops out mainly in Sierra Leone (Allen 1968; Culver & Williams 1979; Culver et al. 1991; Latiff et al. 1997), contains, from the base to the top: the Tabe –Makani formation (180 m) with basal polymict conglomerates overlain by thin shaly sandstone; the Teye formation (200 m) with purple shales, variegated silty shales with quartzite bands; the Kasewe Hill formation (200 m) with grey –green augite andesites, dacitic lavas and tuffs; the Taia formation with grey –green shales and mottled reddish white shales with feldspathic sandstone bands.


According to Culver et al. (1991), the Rokel River trough is an ‘aulacogen’. However, for Latiff et al. (1997) it is a half-graben structure controlled by the Western Kukuna fault. Latiff et al. (1997) favoured an east-dipping fault whereas Culver et al. (1979, 1991) favoured a west-dipping reverse fault. In either case, the southern end of the Rokel River trough is a syncline structure underlined by a basal tillitic level. Folding and metamorphism seem to increase westward but the lack of crops out does not allow us to conclude on a possible thrusting of the central metamorphic formations over the Rokel River trough. The Gibi Mountains Group, which was studied by Magee & Culver (1986), exhibits a basal conglomeratic level overlain by arkosic siltstones and sandstones, which in turn are overlain by shale. The basal conglomeratic level has been related to the glaciogenic basal level of the Rokel River trough by Magee & Culver (1986). The Taban Group is located in several separate troughs located in the southernmost part of the Bove´ basin. From east to west there are three north–south-trending narrow troughs limited by faults: Guemedi, Taban and Bofon troughs. According to Boufeev (1968), they are filled with 2200 m of sandstones and conglomeratic sediments with a great number of rhyolitic clasts. At first, these sediments were considered as younger than those stacked within the Rokel River trough and Kolente basin (Renaud & Delaire 1955; Allen 1968, Boufeev 1968; Villeneuve 1984). However, Culver & Williams (1979) proposed to correlate them to the basal tillitic formations of the Kolente Basin or Rokel River trough. This hypothesis was finally adopted by Latiff et al. (1997) and in the present study. The Taban sandstone beds have an average dip of 458 to the NE. The central metamorphic basement. This comprises the Moussaya complex, the Tabouna complex, the Ouankifondi complex and the Marampa Group. The Moussaya complex corresponds to the western part of the Leo uplift Archaean basement and is composed of granito-gneisses, gneisses, micachists and granites. The Tabouna complex crops out 10 km to the NE of the city of Kindia. Granito-gneisses are sourrounded by the Kolente sediments and caped by the Bove´ basin sandstones. The gneisses are intruded by magmatic rocks dated at 820 + 25 Ma by the Rb/Sr dating method on whole rocks (M. Vachette, pers. comm.). Muscovites from the granite yield an age of 719 Ma and an isochron (muscovite þ chlorite þ whole rock) yields an age of 728 Ma. The Ouankifondi complex (Villeneuve 1984) crops out on both sides of the Moussaya complex

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and consists of amphibolites, gneisses and biotite– muscovite micaschists. Torchine et al. (1969) considered this complex as a metamorphic equivalent of the Kolente Group but Villeneuve (1984) proposed to correlate the Ouankifondi complex, disconformably overlain by the Kolente Group, to the Pan-African I event. Recently (Villeneuve unpubl. data) dated a mylonitic gneiss from Ghemba (5 km NW of Moussaya) by K/Ar on whole rock. This mylonitic gneiss yielded an age of 653 + 15 Ma, similar to that attributed to the metamorphic event that affected the Guinguan Group. The Marampa Group, studied by Junner (1930) and Allen (1969), crops out in many places in tectonic contact with the Archaean basement. It is unconformably overlain by the Rokel River Group. The Marampa Group is the most enigmatic in this part of West Africa because it was successively ascribed to the Archaean klippe of the Kasila Group (Williams & Williams 1979), to a Pan-African klippe of the Kasila Group (Umeji 1988), to a lateral equivalent of the Rokel River Group (Allen 1969) or to an in situ ‘Archaean greenstone belt’ similar to the Kambui greenstone belt of central Sierra Leone (McFarlane et al. 1981). Latiff et al. (1997) identified three tectonic deformational events, and proposed a more ancient history than the Rokel River event. According to the radiometric datings on both the Tabouna granite and the Moussaya gneisses, we propose a Pan-African I age for the Marampa Group and we correlate it with the Guinguan Group of the Bassaride belt. This interpetation is supported by Lytwyn et al. (2006), who pointed out the geochemical similarities between the basalts of the Guinguan and Marampa Groups. The Rokelide thrust belt. The ‘Rokelide thrust belt’ is reduced to the Forecariah Group in GuineaBissau, the Kasila Group in Sierra Leone, and a thin band limited by the ‘Todi shear zone’ in western Liberia. The Forecariah Group (Boufeev 1968; Torchine 1969; Villeneuve 1984) has been divided into four formations: the Kounsouta, Mahera, Kissi-Kissi and Forecariah formations. The main rocks encountered in these formations are garnet–sillimanite – cordierite –biotite gneisses, hypersthene gneisses, kyanite– biotite–garnet gneisses, gneisses and micaschists with amphiboles, two-pyroxenes– garnet biotite gneisses and quartz–hematite (BIFs?). Three stages of metamorphism have been highlighted: a granulite-facies metamorphism resulting from migmatization, an amphibolitefacies phase and a retrograde metamorphism resulting in albitization of the rocks. The post-tectonic leucocratic granite near Coyah has been dated at

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530 Ma on the basis of an Rb/Sr isochron and by 40 Ar/39Ar on biotite concentrates (Dallmeyer et al. 1987). Initially, this Forecariah Group was ascribed to the Archaean basement, but recent datings on single zircon grains (U –Pb method) from syntectonic granites and granulitic gneisses provide an age around 572 + 8 Ma and 558 + 7 Ma for the granulites (Delor et al. 2002). The Kasila Group, which was studied by Williams (1988), includes basic granulites, finegrained granular gneisses (with garnet, hypersthene, brown hornblende), schistose amphibolites, leucograbbros, metasedimentary gneisses and pegmatites with garnetiferous and kyanite schists, and coarsegrained amphibolites with migmatitic folds. Banded iron formations are included in leucogabbros and metasedimentary gneisses and granulites. Williams (1988) described metamorphism at about 10 kbar and 856 + 18 8C. Despite an important number of radiometric data indicating a tectonometamorphic event between 500 and 500 Ma (Allen 1969; Dallmeyer 1989), Williams (1988) maintained his first hypothesis favouring an Archaean suture between the West African and the Guyana shields. In western Liberia, this thrust belt contains melanocratic gneisses locally bearing sillimanite – hypersthene–garnet and two micas, and leucocratic gneisses with kyanite and sillimanite. The southern part of the Bove basin. The Rokelide belt is unconformably overlain by the flat-lying sediments of the Bove´ basin. The three Bove basin groups have been described above. However, the Lower Pita Group includes two formations (Villeneuve 1984): the Guemeta formation at the base and the Kindia formation at the top. Both have a similar (or slightly different) dip but we found (Villeneuve 1984; Villeneuve & Da Rocha Araujo 1984) a cartographic unconformity between them. Discussions and interpretations. I propose that that a Marampa orogenic cycle would precede the Rokelide orogenic cycle, the latter starting with the deposition of the Rokel River basal tillite. According to the radiometric datings quoted above, this cycle is compatible with the Pan-African 1 cycle (670– 650 Ma). Thus, the Rokelide belt would have been affected by the Pan-African 2 tectonic event (570–500 Ma). I also propose that the ‘Guemata formation’ reflects the molassic stage during the build-up of the Rokelide belt. This formation can be compared with the Northern Youkounkoun Group. However, I am not confident regarding the position of the Kasewe Hill (Sierra Leone) and ‘Bania (Guinea-Bissau) volcanic’ formations.

Should they be considered as being interbedded in the Rokel River Group or deposited directly onto the basement? Latiff et al. (1997) confirmed the eastward vergence of the Marampa Group (Pan-African I tectonic event?) and of the Kasila thrusting (Pan-African 2 tectonic event) but also found a westward thrusting of the Rokel River Group (Pan-African 2 tectonic event) over the Marampa Group. However, to bound the westward vergence of the Rokel River Group with the eastward vergence of the Kasila Group it is necessary to link the two areas by a listric fault joining the Rokel River trough to the Kasila thrust. Unfortunately, very few geochemical studies have been performed on the volcanic rocks of the Marampa basalts or the Kasewe hill basaltic formations, or on the Bania basalts and andesites. However, Lytwyn et al. (2006) argued for a geochemical similarity between the volcanics rocks of the Guinguan (Bassaride belt) and Marampa Groups (Rokelide belt). More geochemical information is needed to interpret their geodynamic context. The interpretation by Bonvalot et al. (1991) of their gravimetric data in the Bassaride area is very similar to that of Ponsard (1984). Consequently, I believe that the string of Bouguer anomalies from Senegal to Liberia could be linked to the Bassaride belt. Bonvalot et al.’s gravimetric interpretation is more in accordance with the Bassaride structures than with the Rokelide ones. Radiometric data are limited in number and come from different workers using various methods. In Guinea-Bissau, Boufeev (1968) provided K/Ar datings on whole rocks mainly from the Forecariah Group. Recalculated by M. Vachette (pers. comm.) they display ages between 599 and 437 Ma. Dallmeyer et al. (1987) dated the postorogenic Coyah granite at 537 + 10 Ma by both Rb/Sr and 40Ar/39Ar. The Tabouna granite gives an age of 850 Ma on whole rocks and an isochron age of 719 and 728 Ma on minerals. Delor et al. (2002) dated single zircon grains from the Forecariah complex to around 570 Ma by the U/Pb dating method. In Sierra Leone, Allen (1967) published several ages (K/Ar dating method) for minerals (muscovite, biotite) from the Kasila and Marampa Groups, which are mostly around 530–550 Ma. Dallmeyer (1989) dated the same samples by the 40 Ar/39Ar dating method and obtained similar ages. Rb/Sr isochron datings in western Liberia (Hurley et al. 1971) give ages between 510 and 550 Ma in the Monrovia district and 660 and 700 Ma for the Monrovia leucocratic gneisses. The maximum radiometric dates obtained, between 500 and 600 Ma, correspond to the Rokelide tectonothermal event. However, several ages


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Fig. 10. Three-dimensional geological cross-sections in the western fold belts. (a) Central Mauritanides (Akjoujt and Aouker–Kidira sections); (b) Bassarides (Senegal and northern Guinea); (c) Rokelides (southern Guinea-Bissau and northern Sierra Leone); (d) Rokelides (Liberia). Legend: 1, upper part of the Palaeozoic deposits (Cambrian– Ordovician to Devonian); 2, Youkounkoun Group and Guemeta Formation; 3, Neoproterozoic to Palaeozoic deposits (Mali Group and equivalents); 4, Taban Group; 5, reworked Pan-African 2 calc-alkaline groups (Kasila and Forecariah complex); 6, reworked Pan-African 1 calc-alkaline groups (lower Niokolo–Koba Group); 7, Marampa and Ouankifondi Groups; 8, Termesse Group and equivalents; 9, basic and ultrabasic Pan-African I groups (Guinguan Group and equivalents); 10, Madina– Kouta basin groups; 11, metamorphic rocks of the Mount-Wa-Wa Group; 12, West African craton basement.

192 M. VILLENEUVE Fig. 11. Comparisons between the lithostratigraphic successions of the various sections and belts. Circled numbers: 1, amphibolites; 2, Tabouna complex; 3, El Tlethyate Group; 4, gabbros of Sebkha Gezmayet unit; 5, the lower complex (Bou-Naga intrusive rocks and Eizzenne and Hajar Dekhem Groups); 6, Gadel and M’bout Groups; 7, Guinguan and Lower Niokolo–Koba Groups; 8, Marampa and Ouankifondi Groups; 9, Oumachoueima Group and ‘quartzites’ unit; 10, Selibaby Group; 11, Termesse Group; 12, Kemberra sandstones; 13, Oued Jenne Group; 14, Tichilit-al-Beida and Anietir Groups; 15, Mali Group and equivalents; 16, Rokel River Group; 17, upper parts of the Matallah complex; 18, Mejeria and Ndoumeli Groups; 19, Youkounkoun Group; 20, Guemeta Formation; 21, 22, Pita Group; 23, Palaeozoic formations of Bou-Leriah (Dhloat Ansour Group); 24, Hamdallaye basaltic complex; 25, 26, Telimele´ and Bafata Groups. A– E, differents parts of the sedimentary cover on the Pan-African I (Bassaride) belt; D1–D8, unconformities in the western belts. Circled letters: Sot, Souttoufide tectonic event; PI, Pan-African I tectonic event; PII, Pan-African II tectonic event; H1, early Hercynian tectonic event; H2, late Hercynian tectonic event. Legend: 1, crystalline basement; 2, Souttoufide schists; 3, volcanic formations; 4, stromatolitic limestones; 5, reworked gneissic complex; 6, Kemberra sandstones; 7, volcano-sedimentary rocks of the Termesse Group; 8, tillites and mixtites; 9, shales and sandstones of the Mali Group and equivalents; 10, sandstones of the Youkounkoun Group and equivalents; 11, cross-bedded sandstones; 12, glacial conglomeratic formations of the Late Ordovician (Hirnantian glacial event); 13, black shale formations; 14, limestones and shales of the Dhloat –Ensour Group (northern Mauritanides).


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Fig. 12. Tectonograms showing the western belt’s geodynamic evolution. MKgr, Madina– Kouta basin; LNK, Lower Niokolo–Koba Group; Gui gr, Guinguan Group; Guem, Guemeta Formation; BFT, Bassaride front thrust; UNK, Upper Niokolo– Koba group; Pan.gr, Panampou Group; Tmgr, Termesse Group; Kdg i, Kegougou inlier; Bky, Bakoye Group; Ka/Fo cpx, Kasila– Forecariah complex; RFT, Rokelide front thrust; WA-WA, Wa-Wa Group; Hamd, Hamdallaye volcanic rocks; MFT, Mauritanide front thrust. Legend: 1, crystalline basement; 2, granitic intrusions; 3, syenitic intrusions; 4, Guinguan Group and equivalents; 5, Marampa Group; 6, Pan-African I reworked calc-alcaline complex; 7, Madina–Kouta basin deposits; 8, Bakoye Group sediments; 9, Termesse Group volcano-sediments; 10, tillites and mixtites with shales of the Mali Group and equivalents; 11, conglomerates and sandstones of the Youkounkoun Group and equivalents; 12, volcanic material in the inner troughs (Mount Wa-Wa Group); 13, Siluro-Devonian deposits; 14, thrust fault; 15, compressional or extensional stress directions.

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Fig. 13. Schematic illustrations of the geodynamic and palaeogeographical evolution of the West African margin, from Pan-African to the Late Hercynian times. (a) Crustal block faulting on the West African margin (details and legend as for Fig. 8a). Circled numbers; 1, West African crustal block; 2, Senegalese block; 3, Bassaride block; 4, Rokelide block; 5, Gambia block. GA, Gambia. (b) Opening of the Bassaris rift and connected aulacogen. TB, Taoudeni basin; AO, Atlantic Ocean; RA, Richat aulacogen; MK, Madina–Kouta basin. (c) Closure of the Bassaris rift


between 600 and 750 Ma reflect a later Bassaride tectonothermal event. Conclusions. This synthetic review of the Rokelide belt supports the new hypothesis but some new uncertainties also emerge. The Rokelide belt (c. 550 –500 Ma) consists of two separate parts probably connected to a common listric fault: the Kasila thrust belt to the west and the Rokelide trough, 30 km to the east. These are separated by a metamorphic band composed of Archaean basement overlain by the Marampa thrust belt. The Kasila thrust belt is probably composed of an Archaean basement (possibly the Niokolo –Koba Group) but was strongly remobilized during the Pan-African 2 orogeny. The Rokel River trough has a foreland setting, far from the RFT. It is flat on its eastern margin and slightly folded on its western margin. Considering the unpublished dating results of Ghemba gneisses (653 Ma), the presence of equivalents of the Niokolo–Koba Group near Kindia, and the similarities between the deformations in the Ouankifondi complex and those observed in the Guinguan Group, we think that the Marampa thrust belt could be related to the southern Bassaride belt’s extension. Many questions remains unsolved; for example, the geodynamic model of the Kasila thrust, the tectonic linkage of the Kasila thrust with the Rokel River trough, the statigraphic position of the Taban Group (previously recorded as belonging the Youkounkoun Group and now proposed as an equivalent of the basal tillite of the Rokel River Group). More investigations and, of course, more radiometric datings are needed to improve our knowledge of this younger Pan-African belt.

The West African orogens: structures, correlations and discussion Structure Four block diagrams in Figure 10 display the overall structures of the West African belts along

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different west –east profiles crossing the Mauritanides (block A), the Bassarides (block B) and the Rokelides in Sierra Leone (block C) and Liberia (block D). The central Mauritanide diagram exposes major thrust decreasing to the east (Fig. 10a); The Bassaride diagram for GuineaBissau shows gentle folding of the Cambrian basins and undeformed deposits in the Bove basin (Fig. 10b); the Rokelide diagram for Sierra Leone shows folding of the Rokel River trough and the eastern vergence of the ‘Rokelide thrust belt’ (Fig. 10c); the Rokelide diagram for Liberia shows the ‘Todi Thrust’ (TSZ) and the Gibi mountain klippe (Fig. 10d).

Main orogens and their corresponding tectonothermal events The Pan-African I orogen (Bassaride belt). It is very difficult to date the beginning of this orogenic cycle because of the lack of information about the relationships between the Bassaride oceanic stage and the Madina–Kouta foreland basin. The oldest rocks related to this orogen are probably the Farkaka amphibolites, in the central Mauritanides (Dallmeyer & Lecorche´ 1989). They yield several ages between 870 and 734 Ma, on hornblende concentrates (40Ar/39Ar dating method). The tectonothermal event around 650– 660 Ma (Dallmeyer & Villeneuve 1987) is related to the collision between a western block including the Niokolo Koba Group and the WAC. The molasse stage is not evident. In my opinion it was largely obscured by major rhyolitic eruptions and then the deposition of the Termesse and upper Niokolo– Koba volcano-clastic material. The Pan-African II orogen (Rokelide belt). This concerns all formations included between the Pan- African I disconformity and the Youkounkoun – Pita Group disconformity. A collisional model involving two or several blocks is suspected but only the collision between the Leo uplift and the Guyana shield is well documented. Elsewhere, indices of the Pan-African II event

Fig. 13. (Continued ) during the Pan-African I tectonic event (Bassaride belt). NK, Nouakchott; CNK, Conakry; DKR, Dakar; DK, Dahkla. (d) Opening of several troughs on the western margin with deposition of volcano-sedimentary rocks (Termesse, Wa-Wa, Oumachaueima Groups, etc.) and closure in the Bassaris area (folding of the Termesse Group). BMK, Bamako. (e) Opening of new troughs or enlargement of the previous ones (Fale´me´, Komba, Kolente, Rokel River troughs, etc.) RRT, Rokel River troughs; ND, Nouadhibou. (f) Formation of new belts, in Sierra Leone (Rokelide belt) and in the Adrar Souttouf area(?). Possible new belts in central Mauritania. Folding of the ‘Kounsitel sheet’ in northern Guinea-Bissau. (g) Opening of new troughs or ‘back-arc basin’ in northern and central Mauritania. FR, Freetown. (h) The first Hercynian tectonic event, with imprint of the Reguibat uplift in the Appalachians and southward extrusion of the Senegalese crustal block (thrusting of the Senegalese block onto the Bassaride block, in Bissau). (i) The second Hercynian tectonic event, with eastward motion of the unknown western block and probably emplacement of the main eastward Mauritanide thrusts. Possible opening of the ‘Gambia rift’ (5) by ‘dextral’ strike-slip motion of the Senegalese block along the Bissau– Kidira– Kayes fault zone (pull-apart rifting).

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are suspected (central Mauritanides) or poorly argued (Adrar Souttouf ). The Hercynian orogen (Mauritanide belt). The thrusting of the inner Mauritanide units over the Palaeozoic foreland formations (until Devonian time) has been reported in many places. Metamorphic ages on minerals and whole rocks between 330 and 270 Ma favour a strong Hercynian remobilization. However, very little is known about the Palaeozoic sediments on top of this belt because this Palaeozoic cover was largely destroyed by the Hercynian orogen. Devonian MORB-like basalts are found at Hamdallaye (Lafrance 1996). Other tectonothermal events. Apart from these three main tectonothermal events related to the main orogens, five others are reported. (1) The Souttoufide event. This has ages of 1200– 900 Ma in the Adrar Souttouf. This time span is that of the ‘Grenvillian orogen’ in the eastern USA. (2) The 750 –700 Ma event. Many researchers have reported radiometric ages around 700– 750 Ma. The geodynamic significance of these events is not known. (3) The 600 –580 Ma event. Several ages around 600– 590 Ma are found in many places, especially in the central Mauritanides. (4) The 510 –480 Ma event. Many ages close to 500 Ma have been foundin southern Guinea-Bissau, Sierra Leone and Liberia, and similar ages have been recorded in the central Mauritanides (Dallmeyer & Lecorche´ 1989) and in the northern Mauritanides (Villeneuve et al. 2006). Their significance is not yet clear. (5) The 450 –380 Ma event. Lafrance (1996) reported several Silurian and Devonian ages from the basaltic rocks of the central Mauritanides. Similar ages from basalts have been found in the northern Mauritanides (Villeneuve et al. 2006). No geodynamic explanation has been given for them.

Correlations within the belts Figure 11 attempts to correlate the formations belonging to the different orogenic cycles. Disconformities can be correlated except with the northern Mauritanide area. The Pan-African I orogen. The legend of Figure 11 gives the name of the main lithostratigraphic groups and formations. The Pan-African 2 orogen. This is divided in to three parts (A, B and C) limited by regional unconformities (see the names of groups and formations in the legend of Fig. 11). The Hercynian orogen. This orogen involves all formations included between the Pita Group unconformity and the first Hercynian thrusting occurring around 330 Ma (H1). Rifting or partial oceanization must be envisaged during the Palaeozoic. A

380 + 4 Ma plateau age (40Ar/39Ar) from a pillow basalt sampled in the Hamdallaye transitional (T)-MORB indicates Middle Devonian sea-floor spreading (Lafrance 1996).

Geodynamic evolution Based on field observations, the geodynamic evolution of these fold belts appears to be a succession of compressional and extensional events. Figure 12 distinguishes nine stages in the evolution of these mobile belts (not including the Souttoufide event) and Figure 13 shows the possible palaeogeography at each stage. At least four blocks are distinguished: the West African block, the Senegalese block, the Bassaride block and the Rokelide block. I distinguish (Figs 12 and 13), from the oldest stage to the youngest one, the following phases. (1) Opening of the Bassaride rift around 850– 800 Ma led to the extrusion of the basic volcanic rocks belonging to the Guinguan Group and to the calc-alkaline volcanic units of the Lower Niokolo– Koba Group. (2) Closure of this Bassaride rift led to the formation of the Bassaride belt around 660 Ma. (3) This closure was followed by the deposition of the ‘Kemberra’ sandstones and the spreading and deposition of the volcanoclastic deposits. We notice the folding of the Termesse Group before the deposition of the Mali Group in Senegal and Guinea-Bissau. This event occurred around 620–610 Ma. (4) Deposition of tillites and shales in several troughs parallel to the belts confirms a new extensional stage between 610 and 550 Ma. (5) Folding of the Rokelide belt around 550– 500 Ma was followed by another tectonothermal event in the northern Mauritanides at about 530– 480 Ma. (6) A new extensional event around 510– 480 Ma led to the deposition of the Youkounkoun Group and Guemeta formation. (7) After the deposition of the Late Ordovician tillite, a new phase of opening began within the Central Mauritanides with the eruption of volcanic rocks, whereas the southern Bassaride and Rokelide belts are covered only by marine transgressive sediments. (8) The Final stage, which involved two compressional stages, conclusively fixed the West African belt in its present structural configuration.

General conclusions Nearly 90 years after the first investigations by Hubert (1917) and since the second review by Sougy (1969), many discoveries have been made


on these belts. The distinction into three fold belts, the Mauritanides, Bassarides and Rokelides, is now well established (Villeneuve 1987; Villeneuve et al. 1990) and their main characteristics have been highlighted. Geochemical, geophysical and geochronological data allow us to propose a geodynamic model. In the present review I explain the main divergences between the different researchers and the discrepancies between their interpretations. However, many questions remain unsolved in many places. These are mentioned into this review. Solutions to these questions will require more field investigations. It should be noted that the progress in the understanding of the geodynamic complexities inherent the West African fold belts cannot be compared with that attained in the study of the youngest Late Cenozoic mountain belts because their lifetime is much longer, and because the geologists making field observations in these remote areas are far fewer than in other belts. This situation explains the poor quality of the geodynamic models, which are limited by the small number of sophisticated investigations (geodynamic, radiometric or seismic). Correlation with adjacent areas (Morocco to the north, Brazilian Araguay and Paraguay belts to the south and Appalachians to the west) must be made, to link the geodynamic evolution of the West African belts to that of Rodinia (Murphy & Nance 1991; Murphy et al. 2000) and Gondwana (Von Raumer et al. 2002). However, this is another story. I thank very much the reviewers (K. Attoh, A. Boven and J. P. Lie´geois) for their corrections on a paper that has been reduced to half of its original length. I also thank L. Laugero for the drawing of more than 24 figures. Only 13 are included in the final draft.

References A FFATON , P., R AHAMAN , M. A., T ROMPETTE , R. & S OUGY , J. 1991. The Dahomeyide Orogen: tectonothermal evolution and relationships with the Volta Basin. In: D ALLMEYER , D. & L E´ CORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 107– 122. A KMETJANOV , B. S, L OUTSENKO , V. F. & D IALLO , H. 1976. Carte des anomalies de Bouguer 1/500 000) et re´sultats des travaux geóphysiques en Guineé (Conakry) occidentale. In: Notice de la carte geólogique de la Re´publique de Guineé. Feuille de Gaoual, Technoexport, Conakry, 184 –205. A LLEN , P. M. 1966. Summary Description of the Geology of Western Sierra Leone. 10th Annual Report, Research Institute of African Geology, University of Leeds, 23–24. A LLEN , P. M. 1967. The Geology of Part of an Orogenic Belt in Sierra Leone. PhD thesis, University of Leeds.

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A LLEN , P. M. 1968. The stratigraphy of a geosynclinal succession in western Sierra Leone. Geological Magazine, 105, 62–73. A LLEN , P. M. 1969. Geology of part of an orogenic belt in western Sierra Leone, West Africa. Geologische Rundschau, 58, 588– 620. A MORIN -F ERREIRA , H. 1966. Observacoes Gravimetricas no Territorio da Guinea Portugese. Servicio Meteorologica Nacional, Lisbon. A NDREW -J ONES , D. A 1966. Geology and mineral resources of the Northern Kambui schist and adjacent granulites. Geological Survey of Sierra Leone Bulletin, 6, 100. A NDREW -J ONES , D. A. 1968. Petrogenesis and Geochemistry of Rocks of the Kenema District, Sierra Leone. PhD thesis, University of Leeds. A NGELI , H. 1983. Le magmatisme prote´rozoı¨que de l’ensemble Niokolo–Koba Koulountou, Te´moins d’une marge continentale active au panafricain. DEA, Nancy. B ASSOT , J. P. 1966. Etude geólogique du Seńe´gal oriental et de ses confins guineó-maliens. Me´moires du BRGM, 40. B ASSOT , J. P. 1969. Aperçu sur les formations precambriennes et paleozoiques du Senegal oriental. Bulletin de la Socie´te´ Geólogique de France, 7, 160– 169. B ASSOT , J. P. & V ACHETTE , M. 1983. Donneés nouvelles sur l’aˆge du massif de granitoı¨des du Niokolo–Koba (Seńe´gal oriental); implications sur l’aˆge du stade prećoce de la chaıˆne des uritanides. Journal of African Earth Science, 1, 159–165. B ASSOT , J. P., B ONHOMME , M., R OQUES , M. & V ACHETTE , M. 1963. Mesures d’aˆges absolus sur les se´ries prećambriennes et paleózoı¨ques du Seńe´gal oriental. Bulletin de la Socie´te´ Geólogique de France, 5, 401–405. B ECHENNEC , F. 1980. Etude geólogique du Nord-Est de la Guineé Bissau. BRGM Rapport 79. B EHRENDT , J. C. & W OTORSON , C. S. 1970. Aeromagnetic and gravity investigations of the coastal area and continental shelf of Liberia West Africa and their relations to continental drift. Geological Society of America Bulletin, 81, 3563–3574. B LACK , R., C ABY , R., M OUSSINE -P OUCHKINE , A. ET AL . 1979. Evidence for Late Precambrian plate tectonics in West Africa. Nature, 278, 223–227. B LANC , A., C ARUBA , C., C ARUBA , R., D ARS , R., O HNENSTETTER , D. & P EUCAT , J. J. 1986. Age archeén du socle de la feneˆtre de Bou-Naga (Mauritanie); age panafricain des massifs intrusifs alcalins. 11e`me Reúnion des Sciences de la Terre, ClermontFerrand, Nice University Press, 19. B ONHOMME , M. & B ERTRAND -S ARFATI , J. 1982. Correlation of Proterozoic sediments of Western and Central Africa and South America based upon radiochronological and paleontological data. Precambrian Research, 18, 171– 194. B ONVALOT , S., V ILLENEUVE , M. & A LBOUY , S. 1991. Gravity interpretation of western Sierra Leone (West Africa): implications on the structure and evolution of the Rokelide orogenic belt. Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 312, 841– 848.

198

M. VILLENEUVE

B OUFEEV , Y. B. 1968. Notice explicative de la carte geólogique de la Re´publique de Guineé au 1/200 000. Service Geólogique de Guineé, Conakry. B RINCKMAN , J., M EINHOLD , K. D., B AH , N. ET AL . 2003. Carte geólogique au 1/200 000 des chaıˆnes des Bassarides et les re´gions avoisinantes au Nord-Ouest de la Guineé. BGR et Service Geólogique de Guineé. Ministe`re des Mines, de la Geólogie et de l’Environnement, Conakry. B RONNER , G. & S OUGY , J. 1969. Extension de la glaciation fini-ordovicienne a` la re´gion d’Aoussert (Sahara espagnol me´ridional). Annales de la Faculte´ des Sciences, Clermont-Ferrand, 41, 79– 80. M., C ARIGT , S., H ELG , U., B URKHARD , R OBERT -C HARRUE , C. & S OULAIMANI , A. 2006. Tectonics of the Anti-Atlas of Morocco. Comptes Rendus Geóscience, 338, 11–24. C ABY , R., B ERTRAND , J. M. & B LACK , R. 1981. Pan-African closure and continental collision in the Hoggar– Iforas segment, central Sahara. In: K RONER , A. (ed.) Precambrian Plate Tectonics. Elsevier, Amsterdam, 407– 434. C ASTAING , C., T RIBOULET , C, F EYBESSE , J. L. & C HEVREMONT , P. 1993. Tectonometamorphic evolution of Ghana, Togo, and Benin in the light of the pan-African/Brasiliano orogeny. Tectonophysics, 218, 323–342. C HIRON , J C. 1973. Etude geólogique de la chaıˆne des Mauritanides entre le paralle`le de Moudjeria et le fleuve Seńe´gal(Mauritanie). Me´mories du BRGM, 84. C HOUBERT , G. & F AURE -M URET , A. 1971. Tectonique de l’Afrique. UNESCO, Paris. C LAUER , N. & D EYNOUX , M. 1987. New information on the probable isotopic age of the Late Proterozoic glaciation in West Africa. Precambrian Research, 37, 89–94. C LAUER , N., C ABY , R., J EANNETTE , D. & T ROMPETTE , R. 1982. Geochronology of sedimentary and metasedimentary rocks of the West African craton. Precambrian Research, 18, 53– 71. C LAUER , N., D ALLMEYER , R. D. & L ECORCHE´ , J. P. 1991. Age of the Late Paleozoic tectonothermal activity in north central Mauritanide, West Africa. Precambrian Research, 49, 97–105. C REEN , Y. & R ECHENMAN , J. 1965. Mesures gravime´triques et magne´tiques au Seńe´gal et en Mauritanie occidentale. Cahiers ORSTOM, Se´rie Geophysique, 6. C ULVER , S. J. & H UNT , D. 1991. Lithostratigraphy of Precambrian–Cambrian boundary sequence in the southwestern Taoudeni Basin, West Africa. Journal of African Earth Sciences, 13, 407–413. C ULVER , S. J. & M AGEE , A. W. 1987. Late Precambrian glacial deposits from Liberia, Sierra Leone and Senegal, West Africa. National Geographical Research, 3, 69–81. C ULVER , S. J. & W ILLIAMS , H. R. 1979. The Late Precambrian and Phanerozoic geology of Sierra Leone. Journal of the Geological Society, London, 136, 605– 618. C ULVER , S. J., W ILLIAMS , H. R. & B ULL , P. A. 1978. Infracambrian glaciogenic sediments from Sierra Leone. Nature, 247, 49–51. C ULVER , S. J., W ILLIAMS , H. R. & B ULL , P. A. 1980. Late Precambrian glacial deposits from the Rokelide

fold belt, Sierra Leone. Palaeogeography, Palaeoclimatology, Palaeoecology, 30, 65–81. C ULVER , S. J., P OJETA , J. & R EPETSKI , J. E. 1988. First record of Early Cambrian shelly microfossils from West Africa. Geology, 16, 596–599. C ULVER , S. J., W ILLIAMS , H. R. & V ENKATAKRISHNAN , R. 1991. The Rokelide orogen. In: D ALLMEYER , R. D. & L E´ CORCHE , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 123 –150. C ULVER , S. J., R EPETSKI , J. E., P OJETA , J., J R . & H UNT , D. 1996. Early and Middle(?) Cambrian Metazoan and Protistan fossils from West Africa. Journal of Paleontology, 70, 1–6. D ACHEUX , A. 1967. Etude photogeólogique de la chaıˆne du Dhlou (Zemmour– Mauritanie septentrionale). Laboratoire de Geólogie, Universite´ de Dakar, Rapport, 22. D ALLMEYER , R. D. 1989. A tectonic linkage between the Rokelide orogen (Sierra Leone) and the St Lucie metamorphic complex in the Florida subsurface. Journal of Geology, 89, 183– 195. D ALLMEYER , R. D. & L ECORCHE´ , J. P. 1989. 40Ar/39Ar polyorogenic mineral record within the central Mauritanide orogen, West Africa. Geological Society of America Bulletin, 101, 55–70. D ALLMEYER , R. D. & L ECORCHE´ , J. P. 1990a. 40Ar/39Ar polyorogenic mineral record within the southern Mauritanide orogen (M’Bout-Bakel), West Africa. American Journal of Science, 290, 1136– 1168. D ALLMEYER , R. D & L ECORCHE´ , J. P. 1990b. 40Ar/39Ar polyorogenic mineral record in the northern Mauritanide orogen, West Africa. Tectonophysics, 177, 81–107. D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) 1991. The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin. D ALLMEYER , R. D. & V ILLENEUVE , M. 1987. 40Ar/39Ar polyorogenic mineral age record of a polyphased tectonothermal evolution in the southern Mauritanide orogen, southeastern Senegal. Geological Society of America Bulletin, 98, 602–611. D ALLMEYER , R. D., C AEN -V ACHETTE , M. & V ILLENEUVE , M. 1987. Emplacement age of post tectonic granites in southern Guinea (West Africa) and the peninsular Florida subsurface: implications for origin of southern Appalachians exotic terranes. Geological Society of America Bulletin, 99, 87– 93. D ELOR , C., L AFON , J. M., M ILESI , J. P. & F ANNING , M. 2002. First evidence of 560–575 Ma granulites and syn-tectonic magmatism in the Rokelides belt: geology, geochronology and geodynamic implications. 19th CAG, El Jadida, Morocco, 19–22 March 2002, El Jadida University Press. D ESTOMBES , J., H OLLARD , H. & W ILLEFERT , S. 1985. Lower Palaeozoic rocks of Morocco. In: H OLLAND , C. H. (ed.) Lower Palaeozoic of North-Western and West–Central Africa. Wiley, Chichester, 91– 336. D EYNOUX , M. 1980. Les formations glaciaires du Precambrien terminal et de la fin de l’Ordovicien en Afrique de l’Ouest. Deux exemples de glaciation d’inlandsis sur une plate-forme stable. Travaux du Laboratoire de Science de la Terre, St. Jerome, Marseille, (B), 17.

WEST AFRICAN OROGENIC BELTS D EYNOUX , M., M ARCHAND , J. & P ROUST , J. N. 1989. Notice explicative de la carte geólogique du Mali occidental au 1/200 000. Feuilles Kankossa, Kayes, Kossanto. Re´publique du Mali, Direction Nationale de la Geólogie et des Mines, Bamako; Klo¨ckner, Duisburg, 54– 81. D EYNOUX , M., A FFATON , P., T ROMPETTE , R. & V ILLENEUVE , M. 2006. Pan-African tectonic cycle and glacial climate registered in Neoproterozoic to Cambrian cratonic and foreland basins of West Africa. Journal of African Earth Sciences, 46, 397– 426. D IA , O. 1984. La chaıˆne panafricaine et hercynienne des Mauritanides face au bassin Prote´rozoı¨que supe´rieur a` De´vonien de Taoudenni, dans le secteur clef de Mejeria (Taganet, sud RIM). The`se d’Etat, Universite´ Aix–Marseille III, Marseille. D IA , O., S OUGY , J. & T ROMPETTE , R. 1969. Discordances de ravinement et discordance angulaires dans le ‘Cambro-Ordovicien’ de la re´gion de Mejeria (Taganet occidental, Mauritanie). Bulletin de la Socie´te´ Geólogique de France, 7, 207– 221. D IOP , C. B. 1996. Structures et circulations de fluides dans un avant pays synschisteux: le syste`me de chevauchement des Mauritanides du Seńe´gal. The`se d’universite´, Universite´ de Nancy. D ORBATH , C., D ORBATH , L., L E P AGE , A. & G AULON , R. 1983. The West African craton margin in eastern Senegal: a seismological study. Annales Geophysicae, 1, 25– 36. D ROT , J., L ARDEUX , H. & L E P AGE , A. 1978. Sur la dećouverte de Silurien supe´rieur au sommet de la se´rie de Youkounkoun au Seńe´gal oriental: implications paleógeógraphiques et structurales. Bulletin de la Socie´te´ de Geólogie et Mine´ralogie de Bretagne, 10, 7– 30. D UPONT , P. L., V ILLENEUVE , M. & L APIERRE , H. 1984. Mise en e´vidence de reliques oceániques au sein de la chaıˆne panafricaine des Mauritanides dans la re´gion des Bassaris (Guineé –Seńe´gal). Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 299, 65–70. G EVIN , P. 1960. Etudes et reconnaissances geólogiques sur l’axe cristallin Yetti– Eglab et ses bordures se´dimentaires. Bulletin du Servie de la Carte Geologigne Alge´rie, 23. G IRAUDON , R. 1963. Etude du granit des Hajar Dekhen et des schistes cristallins environnants (re´gion Akjoujt, Mauritanie occidentale). Rapport Ine´dit, BRGM, Dakar, A13. G IRAUDON , R. & S OUGY , J. 1963. Position anormale du socle granitise´ des Hajar Dekhen sur la se´rie d’Akjoujt et participation du socle a` le´dification des Mauritanides hercyniennes (Mauritanie occidentale). Comptes Rendus de l’Acade´mie des Sciences, 257, 937–940. G UETAT , Z. 1981. Etude gravime´trique de la bordure occidentale du craton Ouest africain. The`se 3e`me cycle, USTL, Montpellier. G UETAT , Z., L ECORCHE´ , J. P. & R OUSSEL , J. 1982. Interpre´tation des anomalies gravime´triques de la marge occidentale du craton Ouest Africain. Bulletin de la Socie´te´ Geólgique de France, 7, 763–776. H OEPFFNER , C., S OULAIMANI , A. & P IQUE´ , A. 2005. The Moroccan Hercynides. Journal of African Earth. Sciences, 43, 144–165.

199

H OUDRY , F. 1990. Caracte´risation structurale des tectoniques Panafricaines et Hercyniennes dans les Mauritanides centrales sur la transversale de Barkeol. DEA, Universite´ Nice Sophia-Antipolis. H UBERT , H. 1917. Sur la geólogie du Seńe´gal et des re´gions voisines. Bulletin de la Socie´te´ Geólogique de France, 17, 103–108. H URLEY , P. M., L EO , G. W., W HITE , R. W. & F AIRBAIRN , H. W. 1971. Liberian age province (about 2700 Ma) and adjacent provinces in Liberia and Sierra Leone. Geological Society of America Bulletin, 82, 1004– 1005. J EANNETTE , D. & S CHUMACHER , F. 1976. L’Infracambrien volcano-de´tritique de la bordure occidentale du Massif du Siroua (Anti-Atlas central). Comptes Rendus de l’Acade´mie des Sciences, 282, 823– 826. J EANNETTE , D., B ENZIANE , F. & Y AZIDI , A. 1981. Lithostratigraphie et datation du Prote´rozoı¨que de la boutonnie`re d’Ifni (Anti-Atlas, Maroc). Precambrian Research, 14, 363– 378. J UNNER , N. T. 1930. Geology and mineral resources of Sierra Leone, Mining Magazine, 49, 73– 82. K ENNEDY , W. Q. 1964. The structural differentiation of African in the Pan-African (+500 Ma) tectonic episode. In: 8th Annual Report of Scientific Results University of Leeds, 48–49. K ESSLER , S. F. 1986. Etude structurale et petrographique sur les nappes internes des Mauritanides dans la region d’Akjoujt (RIM). The`se d’universite´, Universite´ Aix– Marseille III, Marseille. K LEIN , E. L., H ARRIS , C., G IRET , A., M OURA , C. A.V & A NGELICA , R. S. 2005. Geology and stable isotopes (H, O, C, S) constraints on the genesis of the Capoeira gold deposits, Gurupi belt, Northern Brazil. Chemical Geology, 221, 188–206. L AFRANCE , A. 1996. La zone frontale des Mauritanides me´ridionales. The`se d’universite´, Universite´ Aix– Marseille III, Aix-en-Provence. L AFRANCE , A., H OUDRY , F., L E P AGE , A., L ECORCHE´ , J. P., V ILLENEUVE , M. & C ARUBA , R. 1993. Reáctivation de la zone frontale d’une chaıˆne polyphaseé: l’exemple de la chaıˆne des Mauritanides (Afrique de l’Ouest). Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 316, 785– 790. L ATIFF , R. S. A., A NDREWS , J. R. & W RIGHT , L. I. 1997. Emplacement and reworking of the Marampa group allochthon, Northwestern Sierra Leone, West Africa. Journal of African Earth Sciences, 25, 333– 351. L ECORCHE´ , J. P. 1980. Les Mauritanides face au Craton Ouest africain: structure d’un secteur clef: la re´gion d’Ijibite`ne (R.I.M). The`se d’Etat, Universite´ Aix – Marseille III, Marseille. L E G OFF , E., G UERROT , C., M AURIN , G., J OHAN , V., T EGYEY , M. & B EN Z ARGA , M. 2001. Dećouverte d’ećlogites hercyniennes dans la chaıˆne septentrionale des Mauritanides (Afrique de l’Ouest). Comptes Redus de l’Acade´mie des Sciences, 333, 711–718. L EPAGE , A. 1983. Les grandes unite´s des Mauritanides aux confins du Seńe´gal et de la Mauritanie. The`se d’Etat, Universite´ Aix– Marseille III, Marseille. L ILLE , R. 1967. Etude geólogique du Guidimaka (Mauritanie). Me´mories du BRGM, 55.

200

M. VILLENEUVE

L YTWYN , J., B URKE , K. & C ULVER , S. 2006. The nature and location of the suture zone in the Rokelides orogen. Sierra Leone, Geochemical evidence. Journal of African Earth Sciences, 46, 439– 454. M AGEE , A. W. & C ULVER , S. J. 1986. Recognition of Late Precambrian glaciogenic sediments in Liberia. Geology, 14, 920–922. M ARTYN , J. & S TRICKLAND , C. 2004. Stratigraphy, structure and mineralisation of the Akjoujt area (Mauritania). Journal of African Earth Sciences, 38, 489–503. M C F ARLANE , A., C ROW , M. J., A RTHURS , J. W., W ILKINSON , A. F. & A UCOTT , J. W. 1981. The Geology and Mineral Resources of Northern Sierra Leone. Overseas Memoir, Institute of Geological Science, London, 7. M ICHAUD , J. G. 1964. Contribution a` l’e´tude geólogique et metallogeńique des environs d’Akjoujt (R.I de Mauritanie). The`se 3e`me cycle, Universite´ de Paris. M URPHY , J. B. & N ANCE , R. D. 1991. Supercontinent model for the contrasting character of Late Proterozoic orogenic belts. Geology, 19, 469–472. M URPHY , J. B., S TRACHAN , R. A., N ANCE , R. D., P ARKER , K. D. & F OWLER , M. B. 2000. ProtoAvalonia: A 1.2– 1 Ga tectonothermal event and constraints for the evolution of Rodinia. Geology, 28, 1071–1074. O ULD S OUELIM , M. 1990. Les roches mafiques et ultramafiques du Guidimaka (Mauritanie) et les gisements de chromite associe´s. The`se d’universite´, Universite´ de Nancy. P ERONNE , Y. 1967. Prospection de 3 secteurs dans le bassin de la haute Gambie. Rapport Annual BRGM, A4. P ONSARD , J. F. 1984. La marge du craton ouest africain du Seńe´gal a` la Sierra Leone: interpre´tation geóphysique de la chaıˆne panafricaine et des bassins du Prote´rozoı¨que a` l’actuel. The`se d’universite´, Universite´ Aix– Marseille III, Marseille. P ONSARD , J. F., L ESQUER , A. & V ILLENEUVE , M. 1982. Une suture panafricaine sur la bordure occidentale du craton ouest africain? Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 295, 1161–1164. P OUCLET , A., G UILLOT , P. L. & B A -G ATTA , A. 1987. Nouvelles donneés lithostructurales, pe´trographiques, mine´ralogiques et geóchimiques du gisement de cuivre d’Akjoujt et son environnement geólogique (R.I de Mauritanie). Journal of African Earth Sciences, 6, 29–43. R ATSCHILLER , L. K. 1970. Lithostratigraphy of the Northern Spanish Sahara. University of Trieste, Institute of Geology, Memoir, 88. R EID , P. C. & T UCKER , M. E. 1972. Probable Late Ordovician glacial marine sediments from Northern Sierra Leone. Nature, Physical Sciences, 238, 38–40. R EMY , P. 1987. Le magmatisme basique des Mauritanides centrales: une ouverture oceánique limiteé d’aˆge Prote´rozoı¨que supe´rieur en Afrique de l’Ouest. The`se d’universite´, Universite´ de Nancy 1. R ENAUD , L. & D ELAIRE , L. 1955. Notice explicative de la carte geólogique de reconnaissance au 1/500 000. Feuille Conakry Est (NC-28, SE-E-12). Direction des Mines AOF, Dakar. R IPPERT , J. C. 1973. Le Tamkarkart: la chaıˆne des Mauritanides contre la bordure du bassin de Taoudeni.

Etude structurale d’un bord de craton. The`se 3eme cycle, Marseille. R ITZ , M. 1982. Etude re´gionale magneto-tellurique des structures de la conductivite´ e´lectrique sur la bordure occidentale du Craton Ouest africain en Re´publique du Seńe´gal. Canadian Journal of Earth Sciences, 19, 1408– 14016. R ITZ , M. & R OBINEAU , B. 1986. Crustal and upper mantle electrical conductivity structures in West Africa: geodynamic implications. Tectonophysics, 124, 115– 132. R JIMATI , E & Z EMMOURI , A. 2002. Notice de la carte geólogique du Maroc au 1/50 000, Feuille Award. Notes et Me´moires du Service Geólogique du Maroc, 439bis. R OMAN ’ KO , Y. F. 1974. New data on the Paleozoic stratigraphy of West Africa. Doklady Academii Nauk SSSR, 217, 420– 423. R OQUES , M. 1948. Le precambrien de l’Afrique occidentale française. Bulletin de la Socie´te´ Geólogique de France, 18, 589–628. S ELIVERSTOV , Y. P. 1970. Notice explicative de la carte geólogique de la Re´publique de Guineé au 1/200 000, feuille C-28– 11. Technoexport, Conakry. S OUGY , J. 1962a. West African fold belt. Geological Society of America Bulletin, 73, 871– 876. S OUGY , J. 1962b. Contribution a` l’e´tude geólogique des guelbs Bou-Leriah (re´gion d’Aoucert, Sahara espagnol). Bulletin de la Socie´te´ Geólogique de France, 4, 436–445. S OUGY , J. 1964. Les formations paleózoı¨ques du Zemmour noir (Mauritanie septentrionale). Etude stratigraphique, pe´trographique et paleóntologique. The`se d’Etat, Universite´ de Nancy. S OUGY , J. 1969. Grandes lignes structurales de la chaıˆne des Mauritanides et de son avant-pays (socle prećambrien et sa couverture infracambrienne et paleózoı¨que), Afrique de l’Ouest. Bulletin de la Socie´te´ Geólogique de France, 11, 133–149. T ESSIER , F., D ARS , R. & S OUGY , J. 1961. Mise en e´vidence de charriages dans la se´rie d’Akjoujt (RIM). Comptes Rendus de l’Acade´mie des Sciences, 252, 1186– 1188. T HORMAN , C. H. 1976. Implications of klippen and a new sedimentary unit at Gibi Mountain, Liberia, West Africa, in the problem of Pan-African– Liberian age province boundary. Geological Society of America Bulletin, 87, 851–856. T ORCHINE , N. S. 1969. Notice explicative de la carte geólogique de la Re´publique de Guineé au 1/200 000, feuille C-28– 12 (Sierumba). Technoexport, Conakry. T ORCHINE , N. S. 1976a. Notice explicative de la carte geólogique de la Re´publique de Guineé au 1/ 200 000, feuilles D 28– 36 et D 28– 29 (Kedougou). Technoexport, Conakry. T ORCHINE , N. S. 1976b. Notice explicative de la carte geólogique de la Re´publique de Guineé au 1/ 200 000, feuilles D 28– 35 et D 28–29 (Youkounkoun). Technoexport, Conakry. T ROMPETTE , R. 1973. Le Prećambrien supe´rieur et le Paleózoı¨que infe´rieur de l’Adrar de Mauritanie (bordure occidentale du bassin de Taoudeni, Afrique de l’Ouest). Un exemple de se´dimentation de craton. Etude stratigraphique et se´dimentologique. Trav aux

WEST AFRICAN OROGENIC BELTS du Laboratire de Science de la St. Je´roˆme, Marseille, (B), 7. T YSDAL , R. G. & T HORMAN , C. H. 1983. Geological Map of Liberia. Map I 1840, 1/1 000 000. US Geological Survey, Reston, VA; Liberian Geological Survey, Monrovia. U MEJI , A. C. 1988. Late Proterozoic to Early Paleozoic supracrustal succession of Sierra Leone: an aulacogen at the western margin of the West African Craton. Geologische Rundschau, 77, 429– 437. V ILLENEUVE , M. 1981. Re´sultats pre´liminaires d’une e´tude geólogique du sud du Fouta Djalon (Guineé). Comptes Rendus Sommaire de la Socie´te´ Geólogique de France, 2, 55–59. V ILLENEUVE , M. 1982. Sche´ma lithostratigraphique des Mauritanides au Sud du Seńe´gal et au Nord de la Guineé d’apre`s les donneés actuelles. Bulletin de la Socie´te´ Geólogique de France, 7, 249– 254. V ILLENEUVE , M. 1984. Etude geólogique de la bordure SW du Craton Ouest africain. The`se d’Etat, Universite´ Aix-Marseille III. V ILLENEUVE , M. 1987. Mise en e´vidence de deux e´ve`nements tectoniques panafricains dans les Bassarides (Guineé–Seńe´gal). Le diachronisme des trois chaıˆnes de la bordure occidentale du Craton Ouest Africain. Comptes Rendus de l’Acade´mie des Sciences, 304, 49– 54. V ILLENEUVE , M. 1989. The geology of the Madina– Kouta Basin (Guineé–Seńe´gal) and its outcomings on the geodynamic evolution of the western part of the West African craton during the upper Proterozoic period. Precambian Research, 44, 305–322. V ILLENEUVE , M. 2005. Paleozoic basins in West Africa and the Mauritanide thrust belt. Journal of African Earth Sciences, 43, 166–195. V ILLENEUVE , M. & D ALLMEYER , R. D. 1987. Geodynamic evolution of the Mauritanides, Bassarides and Rokelides orogens (West Africa). Precambrian Research, 37, 19–28.

201

V ILLENEUVE , M. & D A R OCHA A RAUJO , P. R. 1984. Lithostratigraphie du bassin paleózoı¨que de Guineé (Afrique de l’Ouest). Bulletin de la Socie´te´ Geólogique de France, 1033–1039. V ILLENEUVE , M., B ONVALOT , S. & A LBOUY , Y. 1990. L’agencement des chaıˆnes (Panafricaines et Hercynienne) de la bordure occidentale du craton ouest africain. Comptes Rendus de l’Acade´mie des Sciences, 310, 955–959. V ILLENEUVE , M., B ASSOT , J. P., R OBINEAU , B., D ALLMEYER , R. D. & P ONSARD , J. F. 1991. The Bassaride Orogen. In: D ALLMEYER , R. D. & L E´ CORCHE´ , J. P.(eds) The West African Orogens and Circum Atlantic Correlatives. Springer, Berlin, 151– 185. V ILLENEUVE , M., B ELLON , H., E L A RCHI , A., S AHABI , M., R EHAULT , J. P., O LIVET , J. L. & A GHZER , A. M. 2006. Evenements Panafricains dans l’Adrar Souttouf (Sahara marocain). Comptes Rendus, Geoscience, 338, 359– 367. V ON R AUMER , J. F., S TAMPFLI , G. M., B OREL , G. & B USSI , F. 2002. Organization of pre-Variscan basement areas at the North Gondwanan margin, International Journal of Earth Sciences, 91, 35–52. W ILLIAMS , H. R. 1988. The Archean Kasila group of Western Sierra Leone: Geology and relations with adjacent granite–greenstone terrane. Precambrian Research, 38, 201– 213. W ILLIAMS , H. R. & C ULVER , S. J. 1988. Structural terranes and their relationships in Sierra Leone. Journal of African Earth Sciences, 7, 473– 477. W ILLIAMS , H. R. & W ILLIAMS , R. A. 1976. The Kasila group, Sierra Leone: an interpretation of new data. Precambrian Research, 3, 505– 508. Z IMMERMANN , M. 1960. Nouvelle subdivision des se´ries ante´gothlandiennes de l’afrique occidentale (Mauritanie, Soudan, Seńe´gal). 21st Session, International Geological Congress, Copenhagen, Part VIII, 26– 36.

Neoproterozoic garnet-glaucophanites and eclogites: new insights for subduction metamorphism of the Gourma fold and thrust belt (eastern Mali) R. CABY1, F. BUSCAIL2, D. DEMBE´LE´3, S. DIAKITE´3, S. SACKO3 & M. BAL4 1

Geosciences, Universite´ de Montpellier 2, Place E. Bataillon, F-34095 Montpellier-cedex, France (e-mail: [email protected]) 2

GEOTER, rue Jean Monnet, Clapiers, France

3

Direction Nationale de la Geólogie et des Mines, BP 223, Bamako, Re´publique du Mali 4

Hansa Geomin Consult Gmbh Africa, BP 24257 Dakar, Senegal

Abstract: The Neoproterozoic Gourma fold and thrust belt exposed in eastern Mali includes in its inner part high-pressure, low-temperature metasediments and scarce metabasites. This highpressure metamorphic unit is characterized in the Ansongo region by garnet– glaucophane– paragonite assemblages and eclogites of basaltic derivation, whereas phengite – garnet–rutile mineral assemblages characterize the metapelites. Thermobarometric estimates on the metabasites suggest peak pressure around 13–15 kbar and temperature of 500 + 50 8C for the Seyna Bela garnet glaucophanite and glaucophane-bearing eclogite, and 16 kbar at 600 + 50 8C for the Tin Hama phengite eclogite, values indicative of palaeogeothermal gradients of about 10 8C km21 typical of subduction settings. The high-pressure unit may represents a giant allochthon emplaced on top of very low-grade metasediments. It represents the southern extension of the ultrahighpressure rocks. Garnet pyriclasites from the Amalaoulaou massif, which represents the roots of a c. 800– 730 Ma Neoproterozoic island arc, underwent a medium-temperature metamorphic overprint characterized by barroisite–paragonite assemblages; that is, of same grade as the decompression that affected the eclogites and garnet-blueschists. The Gourma high-pressure metamorphic belt formed as a consequence of the east-dipping subduction of the Neoproterozoic passive palaeo-margin of the West African craton. The presented P– T estimates suggest that subduction-related palaeogeothermal gradients during the late Neoproterozoic period along the main Pan-African suture were similar to those reconstructed for Tertiary Alpine-type belts.

Most occurrences of Alpine-type high-pressure, low-temperature (HP– LT) metamorphic rocks are known from peri-Pacific and peri-Mediterranean late Cretaceous to Tertiary fold belts generated in subduction settings. Several occurrences of HP– LT rocks including glaucophanites have been reported from the Variscan belt in western Europe and Asia. Rare Proterozoic eclogitic assemblages have also been reported from the southern extension of the Pan-African belt in Togo (Agbossoumonde´ et al. 2001) and in Brazil (Beurlen et al. 1992). Besides the present study area in eastern Mali, some occurrences of blue amphiboles have been reported from the Pan-African belt in northern Mali at Timetrine (Fig. 1; Caby & Buscail 2005) and the Anti Atlas in Morocco (Hefferan et al. 2002). However, the only available chemical data on blue amphiboles reported from the two latter areas point to ferro-glaucophane and winchite compositions that are not symptomatic of high-pressure metamorphism.

We describe here the first occurrence of Precambrian glaucophane sensu stricto and glaucophanebearing eclogites so far reported from the Precambrian shield of West Africa. These new findings occur within the Pan-African Trans-Saharan belt along the suture in the Gourma area (eastern Mali), where other occurrences of HP–LT and ultrahigh-pressure (UHP) metamorphic rocks have been described c. 200 km to the NW (Caby 1994; Jahn et al. 2001). We present thermobarometric data compatible with an HP –LT prograde P– T path of blueschist facies culminating in garnet blueschist and eclogitic conditions in the Ansongo region (Figs 1 and 2). We then discuss the possible relationships between these medium-temperature HP occurrences and the HP mafic granulites that delineate the Pan-African suture (Caby 1989; Attoh 1998; Agbossoumonde´ et al. 2001), and present an overview of eclogitic rocks reported along this major suture between the Saharan regions and the Gulf of Guinea.


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Fig. 2. Simplified geological map of the Gourma fold and thrust belt. Location shown in Figure 1. SB, Seyna Bela; TH, Tin Hama; ZH, Zobar Hills.

Fig. 1. Schematic geological map of northeastern Mali (Adrar des Iforas active palaeomargin and Gourma passive margin). 1, Late Cretaceous to Recent cover; 2, Permian to Jurassic rift; 3, Neoproterozoic continental passive palaeo-margin metasedimentary formations of the Gourma basin and Time´trine area (Tim); 4, external nappes; 5, internal nappes with HP–LT (a) and UHP (b) metamorphism; 6, Palaeoproterozoic rocks of the WAC and Bourre´ inlier; 7, Pan-African granitoids; 8, Neoproterozoic active continental palaeo-margin; 9, undifferentiated Pan-African gneisses (inliers of reworked Palaeoproterozoic basement are not shown); 10, Palaeoproterozoic granulite facies microcontinent; 11, intra-oceanic arc terrane (Tilmemsi þ Amalaoulaou); 12, slices of oceanic lithosphere (Time´trine); 14, major fault; 15, thrust; 16, normal fault.

Geological and structural setting The Gourma area represents the deformed passive palaeo-continental margin of the eastern West African craton (WAC) (Reichelt 1967, 1972). The eastern Gourma fold and thrust belt represents the more external domain of the Neoproterozoic Trans-Saharan belt west of the suture zone (Caby 1979, 1994; Sacko 1985; Buscail & Caby, 2005; Caby et al. 2005; Figs 1 and 2). Regional field mapping and classical petrostructural studies allow us to distinguish in eastern Gourma three domains displaying different lithostratigraphic sequences, structural styles and metamorphic evolutions. The external nappes mainly include greenschistfacies monotonous phyllites displaying a flat-lying

cleavage. Chlorite, white mica, albite, epidote and carbonate are the common mineral assemblages in the metapelites and quartz-schists, with the local occurrence of Mn-rich garnet in metapelites and of pale brown biotite in ferro-magnesian metapelites. In the south, the nappes were transported southwestward above parautochthonous metasediments that represent the deformed inner part of the Gourma basin. The Bourre´ domain encompasses reworked Palaeoproterozoic basement overlain unconformably by a metasedimentary cover. The calc-alkaline granitoids of the basement dated by the U – Pb method at 2080 Ma (De La Boisse & Lancelot 1977) have intruded a supracrustal sequence that is strikingly similar to the Palaeoproterozoic (‘Birimian’) formations in Burkina Faso (Lompo 2001). The Palaeoproterozoic metasediments include quartz-free chlorite–albite schists, metabasalts and metatuffs, manganesiferous metasediments (‘gondite’), aluminous quartzite and black phyllites. The unconformable monometamorphic cover labelled the Fafa formation encompasses conglomerates, fluviatile and aeolian quartzites, quartz-arenites (2500 m) and terrigeneous turbidites (1500 m). Basement and cover have been deformed by SW-verging and -plunging folds. The Palaeoproterozoic basement was reworked at high greenschist-facies conditions (Caby & Moussine-Pouchkine 1978). Pan-African mineral assemblages are characterized in orthogneisses by phengite þ green biotite þ albite þ actinolite þ epidote þ carbonates, suggesting regional temperatures around 400 8C. The coexistence of garnet þ phengite þ epidote þ microcline in aluminous orthogneiss may suggest high-pressure conditions. However, no blue amphibole, but actinolite, is observed in metabasites. The Internal Domain (here labelled the internal nappe, Fig. 2) differs from the two other units of the region by a refolded recumbent metamorphic

SUBDUCTION METAMORPHISM, EASTERN MALI

foliation and by a high-pressure, low- to mediumtemperature regional metamorphism. Rock-types comprise coarse-grained garnet – phengite micaschists, metaquartzites, calcareous quartz-schists and rare occurrences of metabasic rocks. The terrigeneous metasedimentary sequence was deposited in a quiet marine environment of the assumed innermost part of the passive palaeo-margin of the West African craton located to the east of the Bourre´ massif. The ubiquitous occurrence of phengite þ garnet þ rutile mineral assemblages symptomatic of high-pressure affinity in this unit was first pointed out by Caby (1979), although evidence for local very high-pressure metamorphism in the In Edem area (Fig. 2) had already been noted by Reichelt (1972) and De La Boisse (1981). More recently, the discovery of coesite in eclogitic metasediments and of mafic eclogites c. 200 km to the NW of the study area at In Edem (Caby 1994) has led to the recognition at this locality of a typical UHP metamorphism formed under low geothermal gradient of c. 10 8C km21. The geodynamic context of this metamorphism dated at 620 Ma (Jahn et al. 2001) is an east-dipping subduction of the West African passive continental palaeomargin underneath late Neoproterozoic oceanic lithosphere. The southeastern extremity of the Gourma fold and thrust belt in the Ansongo region is in contact to the east with mafic rocks of the suture zone exposed in the Amalaoulaou metabasic complex (Fig. 2). This massif coincides with a strong positive gravimetric anomaly, the asymmetry of which is consistent with the existence of a NE-dipping unrooted body c. 10 km thick that allows us to define precisely the Pan-African suture (Bayer & Lesquer 1978). The Amalaoulaou massif essentially consists of mafic granulites (garnet pyriclasites) derived from tholeiitic layered gabbros, pyroxenites, stocks of quartz-gabbro, mafic tonalite and mafic dykes. De La Boisse (1979) obtained U– Pb zircon ages (conventional method) around 810 Ma on a pyriclasite, 730 Ma on a tonalite and 633 Ma on a calcic pegmatoid. These poorly defined ages, however, imply that the granulitic metamorphism took place prior to 730 Ma. The massif is in tectonic contact above greenschist-facies metaquartzites of the Talde´ unit (Fig. 2). Its sole thrust consists of retrogressed amphibolites, chlorite schists, jaspers and serpentinites derived from harzburgites and dunites that contain badly preserved orthopyroxene and chromite clasts, an assemblage that portrays a piece of Neoproterozoic oceanic crust and mantle. The Amalaoulaou complex may represent the roots of a granulitized intraoceanic arc similar to and roughly synchronous with the 730 Ma Tilemsi intra-oceanic arc (Caby et al. 1989; Dostal et al. 1994).

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Main lithology of the HP nappe in the Ansongo region Metaquartzites displaying thin mica-rich and plagioclase-rich layers form the main reliefs of the Ansongo area. Rare occurrences of fresh kyanite are observed in aluminous quartzites. Sedimentary features in orthoquartzites such as curved oblique bedding are locally preserved. Some impure calcmagnesian quartzite layers are poorly exposed and deeply weathered. The flat areas correspond to regularly layered micaschists displaying abundant synmetamorphic quartz veins enriched in phengite megacrysts and rutile prisms occasionally up to several centimetres long. Garnet-, plagioclase-, amphibole- and carbonate-rich layers can be traced as regular gently dipping layers involved in isoclinal folds (metres, decimetres or kilometres in scale). The mean trend of stretching or mineral lineations is east–west. A thinly layered mafic bed up to 50 m thick has been mapped 7 km south of Ansongo on both sides of the Niger River over a distance of 5 km (Buscail & Caby 2005). This mafic layer is exceptional, as it has entirely recrystallized into a garnet-blueschist massive rock containing eclogitic pods. One eclogitized mafic sill possibly connected with this basaltic volcanism has also been observed c. 6 km SE of Tin Hama (Fig. 2). This latter eclogitic occurrence has the major element composition of a K-poor tholeiitic basalt (Table 2) and compares well with the two eclogite occurrences described from the In Edem area (Caby 1994).

Paragenetic analysis and mineral chemistry of selected samples The following petrographic study focuses on the less retrogressed rocks. Mineral analyses (Table 1; Figs 4–6) have been performed at Montpellier II University using a Cameca Datanim microprobe; at operating conditions of 20 kV and 10 nA; natural silicates were used as standards.

Metasediments Based on observations on about 70 thin sections of micaschists and quartz-schists containing garnet, white micas and quartz as chief minerals, it is possible to distinguish plagioclase-, carbonateand amphibole-rich varieties. Early HP assemblages are best preserved in mica-poor, garnet-rich samples in which rutile is the only primary Ti-bearing phase. In many rock types, the poorly expressed high-pressure foliation defined by phengite, occasionally with a random fabric, was diversely overprinted by a low-temperature

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Table 1. Selection of microprobe analyses and structural formulae used for thermobarometric calculations Amphibole

Garnet

S614A 05.9C 05.8A 05.8C core core core core

05.8C core

S614

57.50 0.06 11.30 12.25 7.91 0.01 1.79 6.30 0.03 97.15

25.06 0.01 41.10 5.16 21.63 0.13 0.04 0.02 0.01 93.12

05.9C core

05.9C rim

05.8A core

05.8A rim

05.8C rim

05.9C core

05.9C core

Phengite Paragonite 05.8A core

05.8C core

S.618 core

05.9C core

37.63 38.28 38.63 38.02 37.55 38.38 56.34 55.35 56.03 55.85 0.04 0.12 0.08 0.09 0.05 0.02 0.07 0.04 0.06 0.05 21.45 21.25 21.72 21.34 21.07 21.87 10.92 7.35 8.80 10.47 4.03 0.27 0.62 0.91 1.80 0.71 9.68 8.94 3.03 13.11 29.62 2.29 4.32 2.42 2.57 4.56 3.46 4.95 9.05 3.54 1.09 33.10 27.66 29.82 34.32 30.08 3.66 6.02 2.79 1.21 6.85 1.22 0.30 1.32 0.86 0.19 0.03 0.04 0.16 0.01 0.00 5.81 8.19 7.78 3.95 5.83 6.28 9.89 14.40 6.33 0.00 0.00 0.02 0.02 0.02 0.00 10.67 8.33 6.36 11.03 100.95 102.33 101.54 101.72 102.20 101.66 101.12 100.92 100.68 101.59

54.16 0.18 23.25 4.86 2.63 0.00 0.01 0.31 10.13 95.57

48.85 0.07 39.30 0.09 1.20 0.02 0.14 7.15 0.45 97.26

58.05 0.01 11.30 10.14 10.35 0.02 0.57 7.34 0.02 97.79

47.62 0.29 12.25 10.06 14.86 0.15 6.63 4.96 0.28 99.15

58.47 0.05 10.76 10.44 10.77 0.02 0.80 7.17 0.02 97.11

7.941 0.001 1.823 0.146 2.069 1.038 0.002 0.083 1.947 0.004 15.051

6.842 0.031 2.074 0.893 2.154 0.893 0.019 1.021 1.381 0.052 15.360

7.957 0.005 1.726 0.144 2.119 1.082 0.003 0.117 1.893 0.003 15.050

1.011 0.000 1.954 0.311 0.730 0.004 0.002 0.002 0.000 2.000 6.011

2.948 0.003 2.002 0.097 0.471 1.832 0.072 0.000 0.000 0.000 8.000

3.003 0.007 1.965 0.016 0.268 2.171 0.081 0.488 0.001 0.000 8.000

2.989 0.004 1.981 0.036 0.498 1.790 0.019 0.679 0.004 0.000 8.000

2.983 0.005 1.973 0.054 0.283 1.957 0.088 0.654 0.004 0.000 8.000

2.965 0.003 1.961 0.107 0.302 2.266 0.058 0.335 0.004 0.000 8.000

2.978 0.001 2.001 0.042 0.528 1.952 0.013 0.485 0.000 0.000 8.000

2.008 0.002 0.459 0.260 0.184 0.109 0.001 0.240 0.738 0.01 4.000

2.013 0.001 0.315 0.245 0.268 0.183 0.001 0.385 0.588 0.01 4.000

1.993 0.002 0.369 0.081 0.480 0.083 0.005 0.549 0.439 0.01 4.000

0.251 0.334 0.293 0.338

0.701

0.795 0.62 0.16

0.890 0.72 0.09

0.782 0.60 0.17

0.874 0.66 0.09

0.882 0.77 0.10

0.787 0.66 0.18

0.372

0.406

0.147

Si 7.861 Ti 0.006 Al 1.820 Fe3þ 0.068 Mg 2.497 Fe2þ 0.837 Mn 0.002 Ca 0.263 Na 1.670 K 0.004 Sum 15.026 XFe

Clinopyroxene

Xalm Xpyr

1.985 0.001 0.438 0.350 0.187 0.036 0.000 0.241 0.760 0.01 4.000

Si4þ 0.161 XNa

3.590 0.009 1.816

3.056 0.003 2.898

0.480 0.146 0.000 0.001 0.039 0.857 8.940 3.59 0.04

0.09 0.063 0.001 0.009 0.868 0.036 8.943 3.056 0.960

R. CABY ET AL.

SiO2 TiO2 Al2O3 MgO FeO MnO CaO Na2O K2O Sum

Chloritoid


Table 2. Chemical composition of Tin Hama eclogite (sample S.618) Oxide

wt%

SiO2 TiO2 Al2O3 MgO FeO MnO CaO Na2O K2O P2O5 Sum

49.52 2.46 13.84 6.08 15.26 0.25 10.46 1.88 0.26 0.26 100.30

retrogression. The latter was synchronous with a late metamorphic microfolding and spaced cleavage parallel to the steep axial planar surfaces of the NW –SE- to east– west-trending open folds of any scale that refold the metamorphic foliation in the Ansongo region. Garnet micaschists contain c. 20 vol% of garnet, which may reach 1 cm in diameter. Two stages of growth separated by a quartz ring with disseminated phengite are clearly identified in many samples. Albitic micaschists contain albite poikiloblasts up to 1 cm across (commonly 20– 35 vol%). Albite displays numerous minute inclusions of quartz, white mica, euhedral bluish acicular calcic amphibole, garnet, carbonate, rutile and paragonite, and therefore does not represent a late phase. Mica-rich samples (c. 30 vol%) contain phengite, paragonite, Mg chlorite, garnet (15 vol%), quartz, rutile and blue amphibole. Pressure shadows of albite from the more felsic rocks are formed by paragonite, whereas those of garnet are filled by Mg-chlorite. Fe chlorite is seen replacing amphibole, whereas Mg-chlorite replaces garnet and also forms unoriented rosettes. Rare brown biotite to phlogopite is observed in very few mica-rich samples, either as thin lamellae hosted by phengite, or as a discrete rim around garnet. Up to 20 vol% of carbonate (ferroan dolomite) is observed in Ca–Mg layers that are also rich in epidote (clinozoisite with a zoisite core), grossular-rich poikilitic garnet, white micas, quartz, albite and titanite overgrowing relict rutile. Euhedral garnet occasionally included in dolomite contains a dense network of oriented rutile needles that probably formed as a result of TiO2 exsolution. Amphibole– garnet micaschists to massive quartz – amphibole-schists exposed south of Zobbar hills (Fig. 2) contain abundant bluish amphibole prisms of actinolite composition locally up to 1 cm long (up to 10 vol%) in part replaced by chlorite. Sample S.610.E, a rather massive rock

207

from this locality, has the primary mineral assemblage quartz þ phengite þ paragonite þ albite þ garnet þ actinolite þ carbonate þ rutile. The euhedral garnet core has the composition alm63 pyr8 gro27 spe2. It is sharply separated by a ring having composition alm72 pyr11 gro17 spe1. Phengite with Si ¼ 3.30 –3.36 is surrounded by paragonite (XNa ¼ 0.92). Mg chlorite, titanite and minor epidote represent secondary minerals.

Mafic rocks The Seyna Bela garnet glaucophanite exposed 2 –3 km east of the village (Fig. 2; coordinates 233005 1727654 WGS84 UTM 31 N) is a regular horizon of homogeneous massive dark blue rock with a thin-layered structure (a few millimetres in scale) defined by variable abundance of white micas, amphiboles, garnet and carbonate. This uncommon layer is intercalated between metaquartzites below and silvery quartz-poor paragonite schists on top, which contain abundant unretrogressed glaucophane prisms. The precursor of the layered mafic rock may derive from a basaltic tuff or epiclastite. Typical unaltered samples of this conspicuous layer contain c. 5– 30 vol% of blue glaucophane prisms; these represent a relict phase most frequently mantled by a thin rim of barroisite. Replacement features as well as patches of barroisite and symplectites of barroisite–albite are observed towards the glaucophane rims. The mutual arrangement of secondary amphibole suggests that prismatic glaucophane in textural equilibrium with garnet was the chief mineral (50 vol%) in some dark blue–green samples. In sample S614A, collected from the Seyna Bela, the chemistry of the pale blue glaucophane (Fig. 3a) characterized by Al-, Mg-rich and Ca-poor compositions (Table 1 and Fig. 4a) is similar to those from Alpine white schists (Kieńast et al. 1991). Minute garnet (2– 3 mm) has homogeneous composition alm65 pyr16 gro19 spe02 (Table 1, Fig. 5). Some cores contain primary inclusions of rutile, glaucophane, epidote and paragonite. Rims are in contact with barroisite but a few garnet–glaucophane joints free of barroisite reaction rims are also observed (Figs 3a and 4b). Paragonite (XNa ¼ 0.91– 0.98) is in contact with clinozoisite that frequently forms a thin reaction rim. Prismatic epidote that nucleated around an allanite nucleus shows Fe enrichment from core to rim (XFe ¼ 0.46–0.66). Ferroan dolomite (c. 5 vol%) is in contact with glaucophane and paragonite, suggesting its immobile behaviour during retrogression. Rutile is included in most minerals and has been replaced by ilmenite only when surrounded by very rare chlorite–albite– quartz symplectites.

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R. CABY ET AL.

Fig. 3. Microphotographs of the most representative studied rocks. (a –g) Seyna Bela. (a) Garnet glaucophanite. The textural equilibrium between glaucophane and ferroan dolomite and the barroisite reaction rim between garnet and glaucophane should be noted (sample S. 614). (b –d) Omphacite-bearing glaucophanite schist (sample 05.8C). Note worthy features are the thin symplectite rim around prismatic omphacite (b), the textural equilibrium between paragonite and garnet (c) and the chloritoid inclusion in garnet (d). (e–g) Glaucophane-bearing eclogite.


(a)

S.610 quartz-schist

0

* S 614.A

0.2

S614A Seyna Bela Gt glaucophanite

05.9C

Fe-Glaucophane

S.618 Tin Hama eclogite

*

XMg

05.8A 0.4 0.6 0.8

209

Crossite

Alm + Sps

05.8C

** *

** * * * Glaucophane

05.9C Gl eclogite 05.8A Barroisite eclogite 70

50

50

1 8.0

7.5

(b)

Pyr

Gro

7.0

10

Si in formula

05.8C Gt-Cpx glaucophanite

30

50

Fig. 5. Composition of garnets.

0.00

Barroisite S 603.2 Winchite

* XMg

*

S 618

*

0.50

1.00 8.00

*

**

S 614.A

Ferrowinchite

Ferrobarroisite

7.50

7.00

6.50

Si in formula

Fig. 4. (a) Composition of Na-amphiboles of the glaucophane group from the Seyna Bela garnet glaucophanites and eclogites. (b) Composition of Ca–Na amphiboles (after Leake 1978).

Quartz forms about 3 vol% and poikilitic albite 3 vol%. A few decimetre-scale pods of clinopyroxenebearing rocks occur towards the core of the glaucophanitic layer 2 km to the east of the Seyna Bela village. Owing to the patchy, or vein-type concentration of garnet and omphacite, these quartz-poor (1 vol%) pods contain abundant syn-eclogitic segregates. Sample 05.8C, a mica-rich garnetglaucophanite collected from the cortex of a pod, highlights the transition between slightly retrogressed garnet glaucophanite and eclogite. The rock contains c. 1 vol% of pale green Na clinopyroxene prisms that are free of retrogression when armoured in quartz, or are otherwise rimmed by

minor symplectite (Fig. 3b). This clinopyroxene plots in the augite –aegirine field (Fig. 5). The foliation is defined by the shape fabric of glaucophane (c. 40 vol%), the planar disposition of paragonite (c. 15 vol%) and quartz (15 vol%). Euhedral garnet (c. 25 vol%) is in contact with glaucophane and paragonite (Fig. 3c). A few inclusions of chloritoid (XFe ¼ 0.70) are observed in garnet cores (Fig. 3d) with composition alm72 pyr09 gro16 spe03 and in rims with composition alm66 pyr18 gro16. Fe-dolomite includes all HP minerals. Inclusions of minute rutile grains are observed in all minerals. This unretrogressed sample is cut by some quartz veinlets containing clinopyroxene prisms up to 1 cm long. Sample 05.9C is a non-foliated quartz-free glaucophane-bearing eclogite with millimetre-thick Fe-dolomite-rich layers. It contains c. 40 vol% of garnet, a similar amount of Na pyroxene, about 5 vol% relict glaucophane and similar amounts of paragonite. Back-scattered electron images reveal that the primary Na pyroxene prisms, up to 1 cm long, consist of two immiscible phases with the compositions of chloromelanite (Jd48) and augite– aegirine (Table 1 and Fig. 5). Glaucophane is commonly rimmed by barroisite but is occasionally in contact with garnet without a reaction rim (Fig. 3e). Garnet composition varies from alm72 pyr09 gro16 spe03 to alm60 pyr17 gro23 spe01 from core to rim. Sample 05.8A, another eclogite collected from the core of the thickest lens, contains 50 vol% of greenish omphacite (Jd36) and irregularly distributed garnet (c. 20 vol%, Fig. 3f). Barroisite (c. 20 vol%) is preferentially localized between garnet and omphacite, and is therefore interpreted as the reaction product of garnet þ Na-pyroxene.

Fig. 3. (Continued) (e) Textural equilibrium between glaucophane, omphacite and garnet (sample 05.9C). (f) Quartz-poor equigranular barroisite-bearing eclogite (sample 05.8A). (g) Early paragonite crystal displaying an inner corona of epidote and barroisite towards omphacite (sample 05.8A). (h–j) Tin Hama eclogite. (h) Barroisite-bearing eclogite with a random structure. (i) Early phengite crystal in textural equilibrium with garnet and omphacite (sample S. 618). (j) Amphibolite-facies overprint: relict phengite surrounded by atoll garnet and overgrown by biotite. The matrix includes amphibole, epidote and oligoclase.

210

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Very few glaucophane relicts surrounded by barroisite are nevertheless observed in this sample. Large flakes of paragonite (c. 5 vol%) display a ,50 mm thick inner rim of epidote (Fig. 3g) and of albite towards clinopyroxene. The above-described mineral assemblages suggest that the primary paragenesis of the garnet blueschist was glaucophane þ paragonite + Na pyroxene þ quartz þ garnet þ ferroan dolomite þ zoisite þ rutile, prograde chloritoid being observed in garnet cores. The occurrence of the eclogitic boudins enclosed in garnet glaucophanite may suggest that they derive from water-poor protoliths represented by unaltered basaltic flow or dolerite. The primary paragenesis of eclogites is regarded as coeval with that from enclosing garnet glaucophanite but with much lower XH2O. Invariable pyrope enrichment from garnet cores to rims reflects temperature increase. However, it is not clear if glaucophane breakdown to form barroisite and resulting from the reaction between Na-pyroxene and garnet occurred at peak temperature within or outside the glaucophane stability field. Albite, epidote and calcite formed during retrogression that was minor in this conspicuous layer. It should be noted that euhedral very pale blue glaucophane entirely devoid of retrogression occurs in the adjacent garnet-free paragonite schists. The Tin Hama eclogite (Fig. 2: 277005 1719460) is a new eclogite occurrence and represents a c. 5– 10 m thick mafic sill that has rather sharp contacts with enclosing garnet –phengite micaschists. The mafic rock can be followed over a few hundred metres. It is a fine-grained mafic eclogite with a random fabric. In sample S.618 omphacite (Jd34 – 41) (c. 10 vol%) is polycrystalline (Figs 3h and 5), thus suggesting that it crystallized on the site of former magmatic pyroxene. Garnet is frequently surrounded by omphacite and in contact with ferroan dolomite, phengite, quartz and amphibole. Garnet cores have the composition alm60 pyr12 gro26 spe02 and the rims alm51 pyr27 gro22 spe01. Phengite (Fig. 3i) has Si ¼ 3.55– 3.59, the highest values being measured from cores (Table 1). Blue–green amphibole of barroisitic composition surrounding garnet and omphacite is considered as secondary. Epidote forms euhedral grains with optical zoning and displays Fe enrichment (8–12 wt% FeO) from core to rim. Rutile is the only titaniferous phase included in all minerals. Neither symplectites nor alteration of garnet and omphacite is observed.

In sample S.621, amphiboles have compositions of magnesio-hornblende and ferro-tchermakite (Fig. 4c). They represent a late phase, having possibly replaced omphacitic pyroxene and/or glaucophane. Relict phengite (c. 5 mm) surrounded by atoll garnet has Si values ranging from 3.30 to 3.40. Garnet cores far from biotite and amphiboles have the mean composition alm71 pyr12 gro15 spe02. The garnet corona close to biotite displays Ca enrichment, with a mean value alm65 pyr09 gro24 spe02. Late blastic plagioclase An24 is superimposed on the other minerals, as is the coarse brown biotite (XMg ¼ 0.52 –0.55). Ilmenite has replaced most of the prismatic rutile. The An content of plagioclase in this sample indicates that intervening late amphibolite-facies conditions affected a high-pressure mineral assemblage, with no sign of the low-temperature retrogression as found elsewhere in the Ansongo region. In summary, we interpret the mineral assemblage from the Seyna Bela garnet-glaucophanite as coeval with the glaucophane-bearing eclogitic parageneses occurring in enclosed boudins. Garnet from the Seyna Bela eclogite displays lower pyrope content than garnet from the Tin Hama eclogite. The latter occurrence is therefore considered to record higher peak temperatures.

Thermobarometric estimates Mineral compositions of selected minerals used for calculations are given in Table 1 and Figures 4–6. All analysed garnets show significant zoning characterized by Fe, Ca and Mn decrease and Mg enrichment towards rims, a classical feature generally interpreted as reflecting growth zoning. From microscopic observations, it is not clear, however, to what metamorphic stage the different compositions of minerals relate, as no zoning is observed in the garnet-glaucophanite (sample S.614A), and as the occurrence of barroisite is controlled first by fluids that may have

*

05.8C Gt-Cpx glaucophanite 05.9C Gl eclogite 05.8A Barr eclogite

Aegirine

25

S618 Tin Hama eclogite 30

augite aegirine

* chloromelanite

15

15 omphacite

Amphibolite-facies overprint An occurrence of coronitic garnet –amphibole rock is conspicuous by the late blastesis of randomly oriented brown biotite up to 1 cm in size (Fig. 3j).

Diopside 10

25

50

60 Jadeite

Fig. 6. Composition of clinopyroxenes from garnet blueschist and eclogites.


promoted the reaction between Na-pyroxene and garnet. The lack of lawsonite in mafic rocks implies temperatures on the high-temperature side of the reaction Lws þ Jd ¼ Pg þ Zo þ Qtz þ H2O. The stability of glaucophane in sample 05.9C suggests temperature 550 8C according to Maresch (1977). The coexistence of glaucophane and paragonite implies a pressure 14 kbar at 550 8C according to Guiraud et al. (1990), a similar value being estimated by those workers for the chloritoid –glaucophane assemblage observed in a few garnet cores. Conventional thermometry based upon the Fe– Mg exchange between garnet and clinopyroxene according to Powell (1985) gives temperatures of about 450 + 50 8C for sample 05.8.C; this low temperature is in agreement with the preservation of chloritoid relicts and glaucophane armoured in garnet cores. A temperature of about 500 + 50 8C is obtained for the Seyna Bela eclogites, and 660–680 8C for the Tin Hama eclogite. Geobarometric estimates based upon Si content in phengite and Jd content in omphacite indicate for such temperatures peak pressure bracketed between 13 and 15 kbar for Seyna Bela samples and 16 kbar for the Tin Hama eclogite. Provisional calculations using the Thermocalc program of Powell & Holland (1988) for the system Na2O –FeO–MgO –Al2O3 –SiO2 –H2O (Guiraud et al. 1990) have given the same values. However, it is not clear to what extent garnet chemistry actually relates to peak pressure. The occurrence of oligoclase, biotite and amphiboles of the hornblende–Tchermakite series in sample S.621 clearly indicates amphibolite-facies overprint to the east possibly resulting from an intervening heating event of unknown significance that affected a high-pressure assemblage in this area, as garnet – phengite thermometry on this sample suggests a fossilized temperature of 495 8C (Krogh & Ra˚heim 1978) for P fixed at 16 kbar.

Low-temperature, high-pressure overprint of the Amalaoulaou metagabbros Overview of the Amalaou massif This massif, which was first investigated by De La Boisse (1979), includes foliated garnet pyriclasites derived from a layered sequence of tholeiitic gabbros and lesser pyroxenites that crystallized around 840 Ma (U –Pb zircon, conventional method, De La Boisse 1979). It overlies a sole of serpentinites, metabasalt and jaspers containing blue Fe-rich amphibole with winchite composition. The sheared serpentinites derive from mantle spinel –orthopyroxene peridotites. This slice underlines the southwestward thrusting of the

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Amalaoulaou massif on top of the Talde´ unit with greenschist-facies metamorphism, itself on top of the Ansongo HP–LT rocks (Fig. 2). The granulitic assemblage of the less retrogressed metagabbros includes clinopyroxene, pyrope-rich garnet (up to pyr55 in cores) with abundant rutile exsolutions, rare hypersthene, calcic plagioclase, late pargasitic hornblende and rutile. Peak temperatures of the granulitic metamorphism calculated by De La Boisse (1979) for the slightly retrogressed garnet pyriclasites from the northeastern part of the massif are around 850–900 8C and pressure is around 10 kbar. Textural relationships suggest that already solidified gabbros recrystallized under near-isotropic HP conditions into granulites. This early stage was followed by pervasive HT shearing. Alternatively, the gabbroic magma may have crystallized directly under HP granulitic conditions, followed by pervasive HT deformation. Plastic deformation of garnet- and pyroxene-rich rocks was followed by the development of steep mylonitic to ultramylonite bands defined by minute clasts of unaltered brown pargasitic amphibole and green magnesio-hornblende. Garnet breakdown took place in selected amphibolitic bands subjected to fluid circulations. All plagioclases were invariably replaced later by low-temperature cryptocrystalline minerals (mainly zoisite þ albite?), except for rare samples of spinel-rich norite and garnet-poor to garnet-free pyroxenite, which contain unaltered labradorite.

Late Pan-African overprint Our sampling has been focused on some garnet-rich pyriclasites that look like eclogites in hand specimen, which form 0.5 –15 cm thick layers exposed some 300 m above the basal sole thrust. Sample IC937 contains c. 35 vol% of 1– 2 cm cataclastic garnet grains and a high content of low-temperature matrix minerals (c. 50 vol%) including epidotes, paragonite, barroisite, carbonate and rutile. Larger garnet fragments from this sample are unzoned and belong to the granulite-facies assemblage, with composition alm33 – 35 pyr44 – 46 gro19 – 21 spe01 similar to that reported by De La Boisse (1979). Indeed, garnet has preserved some inclusions of euhedral clinopyroxene with Na2O 1 wt%. Garnet –clinopyroxene (inclusion) pairs give temperatures of 750–850 8C (Ferry & Spear 1978). Sample S603.3, an adjacent sample, contains similar cataclastic garnet clasts up to 2 –3 cm in diameter, but also millimetre-scale ribbons up to 1 cm thick that delineate shear bands. Some shear bands are formed by 60 vol% of minute euhedral garnet, and others are rich in clinopyroxene porphyroclasts mantled by and grading to a fine-grained mosaic of

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sub-grains. Bluish amphibole with compositions between magnesio-hornblende and Tschermakite (Fig. 4c) has gradually replaced both clinopyroxenes and the HT pargasitic amphibole; this replacement is complete in some shear bands. The site of the former calcic plagioclase is recognized through the blastesis of aggregates of zoisite, paragonite and carbonate. Rutile grains are present as inclusions in both the HT and LT minerals, but coexist with ilmenite and titanite. Rare chlorite is associated with paragonite. The mean composition of the minute garnet clasts is alm42 – 45 pyr32 – 39 gro18 – 22 spe01, but rims in contact with barroisite have the composition alm48 – 54 pyr24 – 30 gro21 – 22 spe01. Minute garnet from this sample and the rims of larger clasts are therefore considered as entirely recrystallized and in equilibrium with the matrix minerals. It is concluded that these garnet-rich samples that resemble eclogites derive from diversely sheared and recrystallized garnet-rich pyriclasite layers in which the granulitic assemblage older than 730 Ma was thoroughly overprinted by the secondary LT assemblage garnet 2, barroisite, epidote–zoisite, rutile, paragonite and carbonate of Pan-African age; that is, around 620 Ma (Jahn et al. 2001). As chloritoid is also present in some samples (De La Boisse 1979), the temperature of the second metamorphism does not exceed temperatures of about 550 8C and approaches that of the Ansongo HP– LT rocks described above. No attempt at thermobarometric estimates was made on such partly overprinted rocks.

Discussion and conclusions P – T evolution of the Ansongo high-pressure, low-temperature rocks The discovery of glaucophane-bearing assemblages in schists, basalt-derived detrital metasediments and metabasalt, and of glaucophane-bearing eclogites nearly free of symplectites allows us to better document the metamorphic evolution of the 400 km long high-pressure metamorphic belt that runs through the Gourma region along the Pan-African suture zone. Regional metamorphism culminated in the northwestern part of this belt, with coesitebearing eclogitic metasediments (white schists) and kyanite eclogites in the In Edem area (Caby 1994). The Ansongo glaucophane-bearing rocks have been slightly overprinted by the growth of calcic amphiboles replacing glaucophane and grown in the matrix. As the In Edem eclogitic rocks, the Ansongo eclogites are nearly free of symplectites. Thermobarometric estimates suggest peak pressure around 13 –15 kbar and temperature

of 500 + 50 8C in the Ansongo area, and 16 kbar at 600 + 50 8C in the Tin Hama area. These values indicate palaeogeothermal gradients of about 10 8C km21 typical of subduction settings (Peacock 1992). These occurrences compare well with those from the North Himalayan belt (de Sigoyer et al. 1997; Guillot et al. 1997). The poorly constrained decompression path characterized by the overprint of barroisite–albite in the Ansongo area is suggestive of cooling during decompression. Biotite–oligoclase –hornblende assemblages in the east may indicate amphibolitefacies overprint, with a possible slight temperature increase, as reported from other HP –LT terranes (ie. the North Himalayan belt; de Sigoyer et al. 1997).

Late Neoproterozoic subduction and occurrences of eclogites in the Trans-Saharan and Dahomeyan segments of the Pan-African belt The occurrences of eclogitic belts in the Tuareg shield have been interpreted to delineate several Neoproterozoic fossil subduction zones (Caby 2003; Caby & Monie´ 2003). The main Pan-African suture zone recognized in central Sahara immediately to the east of the West African craton is delineated by a string of positive gravimetric anomalies (Bayer & Lesquer 1978). The identification of the 730 Ma intraoceanic arc terrane in the Tilemsi area, northern Mali (Fig. 1; Caby et al. 1989; Dostal et al. 1994) and the reinterpretation of the Amalaoulaou complex as a piece of oceanic lithosphere (De La Boisse 1979) resulted in the reinterpretation of the Pan-African belt as an Alpine-type subduction- and collision-related belt in continuity with the Dahomeyan belt up to the Gulf of Guinea (Caby 1987, 1989). The recognition of eclogitic metasediments that host some eclogitic bodies in the Gourma region and the thermobarometric estimates on these HP to UHP rocks in Mali (Caby 1994; Jahn et al. 2001) and in Togo (the inner Atacora nappe, Meńot & Seddoh 1985; Agbossoumonde´ et al. 2001) allow a better evaluation of the significance of this HP belt. This 1500 km long eclogitic belt indeed only consists of passive margin Mesoproterozoic(?) to Neoproterozoic metasediments deposited on the eastern passive margin of the west African craton, which were intruded by very few tholeiitic sills. The Nd model ages of these eclogitized mafic rocks in Togo are younger than 1.15 Ga (Bernard-Griffiths et al. 1991). The Ansongo region exposes glaucophane-bearing rocks dragged down during the subduction to depths of 45 km, whereas In Edem rocks that contain coesite were subducted to


a depth of about 100 km. Eclogites from southern Togo record maximum pressure around 19 kbar (Agbossoumonde´ et al. 2001). The occurrence of a kyanite-bearing mafic eclogite identified in northern Togo at the root of the Kabye´ complex at Yade´ (R. Caby, unpublished result) also requires such minimum pressure.

Gross structure and exhumation tectonics Two cross-sections of the studied area are presented in Figure 7. These interpretative sections, inspired by the studies of Caby (1979, 1994) and Sacko (1985), integrate new detailed geological mapping at the scale of 1:50 000 associated with the measurement of about 250 foliation dips and a similar number of stretching or mineral lineations in the Ansongo area. The structural data are available from the database attached to the Ansongo– Amalaoulaou map (Buscail & Caby 2005). In section A, the NE-dipping thrust between the Bourre´ basement and the HP nappe is inferred from the concordant NE-dipping foliations bearing NE-dipping stretching –mineral lineations observed c. 20 km SE of Ansongo. This contact is elsewhere reworked as a steeply dipping sinistral strike-slip fault. The existence of the Ouattagouna klippe and the thin slice of orthogneiss above the turbidites further illustrates the large nappe structure already proposed by Caby (1979) and Sacko (1985). A northeastward progressive deformation of the buried rigid Palaeoproterozoic basement is inferred from the structure of the Bourre´ horst and its cover, which underwent severe ductile Pan-African deformation (Caby & Moussine-Pouchkine 1978). Section B is rather speculative as far as the deep structure is concerned. An estimated horizontal shortening of several tens of kilometres necessarily implies the existence of a major dećollement at the base of the Gourma basin and of blind thrusts. The antiform that exposes in the east the slice of UHP rocks (c. 2– 3 km thick?) is interpreted as a large anticlinal stack. Reconnaissance work in the Guemri area (c. 15 km SW of Gao, Fig. 2) has identified a major low-angle basal sole thrust delineated by greenschist-facies mylonites to the base of the HP nappe. At this locality, metaquartzites and quartz-schists containing kyanite relicts rest in tectonic contact above sediments of the Guemri half-window, which consists of virtually nonmetamorphic quartzites and siltstones free of slaty cleavage in a upside-down position (Caby 1979). This superimposition suggests that the entire sequence of HP rocks and the slice of In Edem UHP rocks actually represent a giant allochthonous unit. Agbossoumonde´ et al. (2001) have shown in southern Togo that the higher pressure eclogites

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(about 19 kbar) overlie the lower pressure eclogites (about 13 kbar), both emplaced above less metamorphosed metasedimentary units that underwent only a prograde medium-pressure greenschistfacies evolution (Meńot & Seddoh 1985). The entire package of Togo –Benin nappes of the Dahomeyides rests above non-metamorphic units (the ‘Buem’) of the Volta basin (Caby 1989; Affaton et al. 2000). NW of Gao, however, the HP metamorphic foliation of the western limb of the antiform is concordant with the slaty cleavage of the phyllites and sediments of the basin. This suggests that the HP –UHP rocks plunge underneath the parautochthonous –autochthonous metasediments of the Gourma –Mopti basin. The northward progressive flip of the basal contact of the external nappes c. 70 km SW of Gao (Fig. 2) may relate to large-scale post-nappe back-folding. Alternatively, such a geometry change along strike is possibly the result of late underthrusting of the HP metamorphic wedge underneath the rather thick autochthonous turbiditic sediments deposited in the northeastern part of the basin (Moussine-Pouchkine & Bertrand-Sarfati 1978). As in many other UHP terranes (i.e. Western Alps, Chopin et al. 1991, Caby 1996), it is suggested that exhumation of the Gourma HP– UHP rocks from the steeply dipping suture zone requires an extrusion process.

Significance of the high-pressure granulite belt Recent data obtained on the metabasic massifs affected by high-pressure granulite-facies metamorphism to the east of the eclogitic belt highlight their distinctive metamorphic evolution (Duclaux et al. 2006). The possible age of the gabbroic precursors in Mali (Amalaoulaou massif) is around 800 Ma and the granulitization occurred before 730 Ma (De La Boisse 1979). In contrast, ages around 610 Ma interpreted to date the high-pressure granulite-facies metamorphism have been obtained in northern Togo (Bernard-Griffiths et al. 1991; Affaton et al. 2000) and in Ghana (Attoh et al. 1997; Hirdes & Davis 2002). The low-temperature metamorphic overprint evidenced in the garnet pyriclasites close the root of the Amalaoulaou complex by paragonite–zoisite –barroisite– rutile + chloritoid assemblages and the recrystallization of garnet suggest P–T conditions similar to the retrograde evolution of the Ansongo eclogites linked with the exhumation of the HP nappe. It is therefore concluded that the granulitized metagabbros were first uplifted prior to 730 Ma and buried again slightly before 620 Ma in a Pan-African

214 R. CABY ET AL.

Fig. 7. Cross-sections across the Gourma fold and thrust belt. After Caby (1994), modified (see Fig. 2 for location of the sections).


cold tectonic regime within or adjacent to the east-dipping subduction channel from where the Ansongo high-pressure rocks were expelled.

Concluding remarks As suggested previously (Caby 1979) a metamorphic continuity is found between the Ansongo eclogites described here and the In Edem white schists, kyanite eclogites and coesite-bearing UHP rocks crystallized at P 27 kbar and T ¼ 750 8C (Caby 1994). Within the metasediments belonging to the Proterozoic passive palaeomargin of the West African craton, the Ansongo garnet – paragonite glaucophanites and glaucophanebearing eclogites reported here document, as in Togo, an exceptional LT –HP metamorphic regime characteristic of subduction settings. The thermobarometric estimates (T 500– 680 8C, P 13 –16 kbar) indicate minimum pressures, as is frequently the case for eclogitic rocks containing mineral phases that can be stable at much higher pressures (i.e. paragonite and glaucophane; Kieńast et al. 1991). The age of 620 Ma for the eclogitization obtained by several methods at In Edem (Jahn et al. 2001) is identical within error limits to the U –Pb zircon age of the oldest synkinematic calc-alkaline plutons exposed c. 70 km to the east of the suture in the Adrar des Iforas massif (Caby & Andreopoulos-Renaud 1989; Caby 2003). This further illustrates the geodynamic link between the LT –HP metamorphism and the eastward subduction of the West African continental crust under oceanic lithosphere and intra-oceanic arcs, and their subduction under the Iforas palaeocontinent. From the south Saharan regions to the Gulf of Guinea, the main Pan-African suture is delineated by some eclogitic occurrences and by the high-pressure granulitic metabasic massifs derived from arc roots. This work has been undertaken as part of the project ‘Inventaire minier et cartographie geólogique de l’Adrar des Iforas et du Gourma oriental’ supported by the Banque Europeénne d’Investissement (BEI). We are grateful to the Direction Nationale de la Geólogie et des Mines (Bamako) and to GEOTER for logistical support and facilities. We thank C. Passchier and R. P. Meńot for fruitful discussions in the field and for their constructive criticism of the manuscript.

References A FFATON , P., K RO¨ NER , A. & S EDDOH , K. F. 2000. Pan-African granulite formation in the Kabye´ Massif of northern Togo (West Africa): Pb –Pb zircon ages. International Journal of Earth Sciences, 88, 778– 790. A GBOSSOUMONDE´ , Y., M E´ NOT , R. P. & G UILLOT , S. 2001. Metamorphic evolution of Neoproterozoic

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eclogites from south-Togo (West Africa). Geodynamic implications. Journal of African Earth Sciences, 33, 227– 244. A TTOH , K. 1998. High-pressure granulite facies metamorphism in the Pan-African Dahomeyide orogen, west Africa. Journal of Geology, 106, 236–246. A TTOH , K., D ALLMEYER , R. D. & A FFATON , P. 1997. Chronology of nappe assembly in thze Pan-African Dahomeyide orogen, West Africa: evidence from 40 Ar/39Ar mineral ages. Precambrian Research, 82, 153– 171. B AYER , R. & L ESQUER , A. 1978. Les anomalies gravime´triques de la bordure orientale du craton ouest-africain: Geóme´trie d’une suture pan-africaine. Bulletin de la Socie´te´ Geólogique de France, 7, 863–876. B ERNARD -G RIFFITHS , J., P EUCAT , J.-J. & M E´ NOT , R.-P. 1991. Isotopic (Rb–Sr, U– Pb and Sm–Nd) and trace element geochemistry of eclogites from the Pan-African belt: a case study of REE fractionation during high-grade metamorphism. Lithos, 27, 43–57. B EURLEN , H, S ILVA F ILHO , A. F., G UIMARA˜ ES , I. P. & B RITO , S. B. 1992. Proterozoic C-type eclogites hosting unusual Ti–Fe + Cr + Cu mineralization in northeastern Brazil. Precambrian Research, 58, 195– 214. B USCAIL , F. & C ABY , R. 2005. Notice explicative de la carte geólogique du Gourma oriental au 1/200 000 Ansongo– Amalaoulaou. Direction Nationale des Mines et de la Geólogie, Bamako, Mali. C ABY , R. 1979. Les nappes prećambriennes du Gourma dans la chaıˆne pan-africaine du Mali. Comparaison avec les Alpes occidentales. Revue de Geólogie Dynamique et de Geógraphie Physique, 21, 365– 376. C ABY , R. 1987. The Pan-African belt of West Africa from the Sahara Desert to the Gulf of Benin. In S CHAER , J. P. & R ODGERS , J. (eds) Anatomy of Mountain Ranges. Princeton University Press, Princeton, NJ, 129– 170. C ABY , R. 1989. Precambrian terranes of Benin Nigeria and Northeast Brazil and the Late Proterozoic South Atlantic fit. In: D ALLMEYER , R. D. (ed.) Terranes in the Circum-Atlantic Paleozoic Orogens. Geological Society of America, Special Papers, 230, 145–158. C ABY , R. 1994. Precambrian coesite from northern Mali: first record and implications for plate tectonics in the trans-Saharan segment of the Pan-African belt. European Journal of Mineralogy, 6, 235–244. C ABY , R. 1996. Low-angle extrusion of high-pressure rocks and the balance between outward and inward displacements of Middle Penninic units in the western Alps. Eclogae Geologicae Helvetiae, 89, 229–267. C ABY , R. 2003. Terrane assembly and geodynamic evolution of Central– Western Hoggar: a synthesis. Journal of African Earth Sciences, 37, 133–159. C ABY , R. & A NDREOPOULOS -R ENAUD , U. 1989. Age U– Pb a` 620 Ma d’un pluton synorogeńique de l’Adrar des Iforas (Mali). Conse´quences pour l’aˆge de la phase majeure de l’orogene`se pan-africaine. Comptes Rendus de l’Acade´mie des Sciences, 308, 307– 314. C ABY , R. & B USCAIL , F. 2005. Notice explicative de la carte geólogique de l’Adrar des Iforas au 1/200 000. Direction Nationale de la Geólogie et des Mines, Bamako, Mali.

216

R. CABY ET AL.

C ABY , R. & M ONIE´ , P. 2003. Neoproterozoic subductions and differential exhumation of western Hoggar (southwest Algeria): new structural, petrological and geochronological evidence. Journal of African Earth Sciences, 37, 269 –293. C ABY , R. & M OUSSINE -P OUCHKINE , A. 1978. Le horst birimien de Bourre´ (Gourma oriental, Re´publique du Mali): nature et comportement au cours de l’orogene`se pan-Africaine. Comptes Rendus de l’Acade´mie des Sciences, 287, 5– 8. C ABY , R., A NDREOPOULOS -R ENAUD , U. & P IN , C. 1989. Late Proterozoic arc–continent and continent – continent collision in the pan-African Trans-Saharan belt of Mali. Canadian Journal of Earth Sciences, 26, 1136–1146. C ABY , R., B USCAIL , F., D EMBELE , D. & S ACKO , S. 2005. Field Trip Guide Book, Gourma, 3rd IGCP Symposium, Gao 2005, Mali. Ministe`re des Mines de l’Eńergie et de l’Eau, Bamako, Mali. C HOPIN , C., H ENRY , C. & M ICHARD , A. 1991. Geology and petrology of the coesite bearing terrain, Dora Maira massif, Western Alps. European Journal of Mineralogy, 3, 263–291. D E L A B OISSE , H. 1979. Pe´trologie et geóchronologie de roches cristallophyliennes du bassin du Gourma (Mali): conse´quences geódynamiques. The`se 3e cycle, Universite´des Sciences, Montpellier. D E L A B OISSE , H. 1981. Sur le me´tamorphisme du micaschiste ećlogitique de Takamba (Mali) et ses conse´quences paleógeódynamiques au Prećambrien supe´rieur. Comptes Rendus Sommaires de la Socie´te´ Geólogique de France, 3, 97–100. D E L A B OISSE , H. & L ANCELOT , J. R. 1977. A propos de l’e´ve`nement a` 1000 Ma en Afrique occidentale: 1. Le ‘granite’ de Bourre´. Comptes Rendus Sommaires de la Socie´te´ Geólogique de France, 4, 223– 225. D E S IGOYER , J., G UILLOT , S., L ARDEAUX , J. M. & M ASCLE , G. 1997. Glaucophane-bearing eclogites in the Tso Morari dome (eastern Ladakh, NW Himalaya). European Journal of Mineralogy, 9, 1073– 1083. D OSTAL , J., D UPUY , C. & C ABY , R. 1994. Geochemistry of the Neoproterozoic Tilemsi belt of Iforas (Mali): a crustal section of an oceanic island arc. Precambrian Research, 65, 55– 69. D UCLAUX , G., M E´ NOT , R. P., G UILLOT , S., A GBOSSOUMONDE´ , Y. & H ILAIRET , N. 2006. The mafic layered complex of the Kabye´ massif (north Togo and north Benin): evidence of a Pan-African granulitic continental root. Precambrian Research, 151, 101– 118. F ERRY , J. M. & S PEAR , F. S. 1978. Experimental calibration of the partitioning of Fe and Mg between biotite and garnet. Contributions to Mineralogy and Petrology, 66, 113–117. G UILLOT , S., S IGOYER DE , J., L ARDEAUX , J. M. & M ASCLE , G. 1997. Eclogitic metasediments from the Tso Morari area (Ladakh, Himalaya): evidence for continental subduction during India–Asia convergence. Contributions to Mineralogy and Petrology, 128, 197–212. G UIRAUD , M., H OLLAND , T. J. B. & P OWELL , R. 1990. Calculated mineral equilibria in the greenschist– blueschist–eclogite facies in Na2O– FeO– MgO– Al2O3 –SiO2 –H2O. Methods, results and geological applications. Contributions to Mineralogy and Petrology, 104, 58–98.

H EFFERAN , K., A DMOU , H., H ILAL , R. ET AL . 2002. Proterozoic blueschist-bearing me´lange in the Anti-Atlas Mountains, Morocco. Precambrian Research, 118, 179–194. H IRDES , W. & D AVIS , D. W. 2002. U –Pb zircon and rutile metamorphic ages of Dahomeyan garnet– hornblende gneiss in southeastern Ghana, West Africa. Journal of African Earth Sciences, 35, 445–449. J AHN , B. M., C ABY , R. & M ONIE´ , P. 2001. The oldest UHP eclogites of the world: age of UHP metamorphism, nature of protoliths and tectonic implications. Chemical Geology, 178, 143– 158. K IE´ NAST , J. R., L OMBARDO , B., B INO , G. & P INARDON , L. 1991. Petrology of very-high-pressure eclogitic rocks from the Brossasco–Isasca Complex, Dora Maira Massif, Italian Western Alps. Journal of Metamorphic Geology, 9, 19–34. K ROGH , E. J. & R A˚ HEIM , A. 1978. Temperature and pressure dependence of the Fe– Mg partitioning between garnet and phengite with special reference to eclogites. Contributions to Mineralogy and Petrology, 66, 75–80. L EAKE , B.E. 1978. Nomenclature of amphiboles. Canadian Mineralogist, 16, 501– 520. L OMPO , M. 2001. Le Paleóprote´rozoı¨que (Birimien) du Burkina Faso, Afrique de l’Ouest. E´volution crustale et concentrations aurife`res. Me´moire d’habilitation a` diriger des recherches, Universite´ Paul Sabatier, Toulouse III. M ARESCH , W. V. 1977. Experimental study on glaucophane: an analysis of present knowledge. Tectonophysics, 43, 109 –125. M E´ NOT , R. P. & S EDDOH , K. F. 1985. The eclogites of Lato Hills (South Togo, West Africa): relics from tectonometamorphic evolution of the Pan-African orogeny. Chemical Geology, 50, 313– 330. M OUSSINE -P OUCHKINE , A. & B ERTRAND -S ARFATI , J. 1978. Le Gourma: un aulacoge`ne du Prećambrien spe´rieur? Bulletin de la Socie´te´ Geólogique de France, 20, 851–857. P EACOCK , S. M. 1992. Blueschist facies metamorphism, shear heating and P –T – t paths in subduction shear zones. Journal of Geophysical Research, 97, 17693–17707. P OWELL , R. 1985. Regression diagnostics and robust regression in geothermometer/geobarometer calibration: the garnet–clinopyroxene geothermometer revisited. Journal of Metamorphic Geology, 3, 231–243. P OWELL , R. & H OLLAND , T. J. B. 1988. An internally consistent dataset with uncertainties and correlations: 3. Application to geobarometry, worked examples and a computer program. Journal of Metamorphic Geology, 6, 173–204. R EICHELT , R. 1967. Carte geólogique du Gourma. Me´moires du BRGM, 53. R EICHELT , R. 1972. Geólogie du Gourma (Afrique occidentale). Un ‘seuil’ et un bassin du Prećambrien supe´rieur. Stratigraphie, tectonique, me´tamorphisme. Me´moires du BRGM, 53. S ACKO , S. 1985. Contribution a` l’e´tude structurale du Gourma oriental (chaıˆne pan-africaine du Mali). These 3e cycle, Universite´ de Montpellier.

Tectonic significance of carbonatite and ultrahigh-pressure rocks in the Pan-African Dahomeyide suture zone, southeastern Ghana KODJOPA ATTOH1 & PROSPER M. NUDE2 1

Department of Earth and Atmospheric Sciences, Cornell University, Ithaca, NY 14853, USA (e-mail: [email protected]) 2

Department of Geology, University of Ghana, Legon, Ghana

Abstract: The association of carbonatite and ultrahigh-pressure (UHP) metamorphic rocks in the Dahomeyide suture zone of southeastern Ghana is unique among the Neoproterozoic orogens that surround the West African craton (WAC). Carbonatite occurs in an alkaline complex that decorates the sole thrust of the suture zone and is characterized by high concentrations of incompatible trace elements such as light rare earth elements (LREE), Sr and Ba. Within the suture zone deformed alkaline rocks, including carbonatite, together with mafic granulites form an imbricate stack of thrust panels that involve 2.1 Ga rocks of the WAC basement. The dominant rock unit of the suture zone is composed of mafic granulites in which garnet megacrysts preserve a diagnostic microstructure of UHP metamorphism; it consists of a crystallographically controlled array of exsolved rutile rods in garnet. Metamorphic pressures estimated from Ti concentrations in the inferred precursor garnet indicate P .3 GPa, which requires subduction (and exhumation) of the suture zone rocks to and from mantle depths during collisional orogeny on the WAC margin. Available age constraints on carbonatite magmatism suggest that continental rifting, leading to the formation of the passive WAC margin c. 700 Ma, occurred c. 100 Ma before intrusion of carbonatite, which was preceded by HP and UHP metamorphism at 610 + 5 Ma.

The occurrence of deformed alkaline rocks including carbonatite has been proposed as a reliable indicator of old suture zones by Burke et al. (2003). They used a compilation of deformed alkaline rocks and carbonatite (DARC) in Africa by Woolley (2001) to show that the distribution of such rocks correlates with known suture zones or reveals previously elusive sites of ocean closure. One of the DARC occurrences cited in that compilation was in the Dahomeyide orogen in southeastern Ghana, where a Pan-African suture has long been inferred along the southern extension of the Trans-Saharan belt (e.g. Caby 1987; Affaton et al. 1991). Along this suture zone, high-pressure granulites and eclogites (HIPGE) of basaltic composition are the dominant rock types (Attoh 1998a; Agbossoumonde et al. 2001; Attoh & Morgan 2004); however, the significance of the DARC and HIPGE rock association in the Dahomeyide suture zone has not been investigated. In this paper we (1) describe the field relations between the main rock units in the suture zone, (2) present, for the first time, geochemical data on the carbonatite postulated to record continental rifting along the WAC margin, and (3) evaluate petrological evidence for ultrahigh-pressure (UHP) metamorphism, which indicates subduction of the

suture zone rocks to mantle depths during the Pan-African orogeny. These new data from the Dahomeyide orogen provide significant constraints on the Rodinia to Gondwana supercontinent cycle.

Geological setting and previous studies The assembly of NW Gondwana from various cratonic fragments postulated to be derived from the breakup of Rodinia supercontinent (Hoffman 1991; Cordani et al. 2003) resulted in Pan-African (Neoproterozoic) orogens including the 2000 km long Trans-Saharan orogen (Caby 1987; Trompette 1994) located on the eastern margin of the West African craton (WAC). The southeastern segment of the Trans-Saharan belt exposed in southeastern Ghana and adjoining parts of Togo and Benin comprises the Dahomeyide orogen (Affaton et al. 1991; Castaing et al. 1993; Attoh et al. 1997). The principal tectonic elements of the Dahomeyide orogen are (Fig. 1): (1) the deformed edge of the WAC with its cover rocks consisting of cratonverging nappes and thrust sheets bounded by ductile shear zones; (2) the suture zone representing the eastern boundary of the autochthonous WAC; (3) exotic rocks that form the granitoid gneiss complexes east of the suture zone. The suture zone can


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K. ATTOH & P. M. NUDE

Fig. 1. Dahomeyide orogen showing the location of Figure 2 and distribution of high-pressure granulites and eclogite previously analysed. LT, Mont Lato; HZ, Hodzo Hills; AD, Mt Adaklu; KB, Krobo Hills; SH, Shai Hills.

be traced more or less continuously for c. 1000 km along the Dahomeyides on the basis of a distinctive package of mafic and ultramafic rocks that form the protolith of HIPGE rocks (Attoh & Morgan 2004; Duclaux et al. 2006, and references therein). In southeastern Ghana, the HP mafic granulites have been referred to as Shai Hills gneiss (Attoh et al. 1997), and are tectonically juxtaposed with the alkaline gneiss complex in the suture zone (Fig. 2). Although the occurrence of alkaline rocks in this area has been known for some time (e.g. Holm 1974, and references therein) their tectonic significance was not recognized, in part because the magmatic origin of the carbonatite and its association with the other alkali rocks was not resolved. For example, Holm (1974) described the carbonatite unit as ‘an unusual rock that consists essentially of calcite and biotite . . . that appeared to have formed by metasomatic introduction of calcite along a mylonitized zone that developed after or late during the latest regional metamorphism (Pan-African)’. As such, this paper presents the first data to confirm the magmatic character of the carbonatite and to consider the tectonic implications

Fig. 2. Geological map of the suture zone showing the relations between the alkali rocks (Kpong complex) and high-P granulites (Shai Hills gneiss) and rocks of the WAC margin (Atacora nappes and Ho Gneiss). OY, Osu Yongwa; KH, Kluma Hills.

of the DARC and UHP metamorphic rock association in the Dahomeyide orogen.

Field relations in the suture zone Tectonic stacking Figure 2 is the geological map of a segment of the suture zone displaying the distribution of the main lithotectonic units; it is based on Holm (1974), open-file maps by the Ghana Geological Survey Department (see Kesse 1985, for references to Field Sheets 101, 103 and 104), and recent interpretations (Attoh et al. 1997). From NW to SE (Fig. 2), the principal lithotectonic units are: Atacora nappes unit consisting of quartzites, quartz-schists and phyllonites, followed by a mylonitic granitoid gneiss unit characterized by variably developed feldspar porphyroclasts. This protomylonitic granitoid unit, which represents the 2.1 Ga Ho gneiss, was thrust over the Atacora units as is evident in tectonic windows such as the Kluma Hills, where the tectonic stacking can be inferred from the klippen of the Ho gneiss. Slivers of thrust-bounded garnet amphibolite and alkalic gneiss west of the Kluma Hills are interpreted to

CARBONATITES IN DAHOMEYIDE SUTURE ZONE

Fig. 3. Inferred tectonic stacking of the major units in the suture zone.

be imbricated with the mylonitic granitoid gneiss. Further east, the main suture is a DARC-decorated, north–south zone which turns sharply west, SE of the Kluma Hills; along this zone garnet amphibolites and mafic granulites, which form the HIPGE

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rocks and are hereafter referred to as the Shai Hills gneiss unit (Attoh et al. 1997), occupy the hanging wall. The inferred tectonic stacking in the suture zone (Fig. 3) depicts the 2.1 Ga WAC granitoids overthrust by mylonitic granitoid rocks derived from them at the base and intruded by alkali gneiss. This stack was overthrust by nappes of garnet amphibolites and high-pressure mafic granulites of the Shai Hills gneiss unit. The overall structure of the suture zone is interpreted to have resulted from early east –west compression, which produced the north–south imbricate thrust slices followed by NNW-directed thrusting in the orogen (Attoh et al. 1997). The low-angle dips of the thrust surfaces in the suture zone account for the complex map pattern in this area (Fig. 2), such as the repetition of traces of the thrusts and the lithotectonic units. Asymmetric feldspar porphyroclasts, with well-developed tails in the protomylonitic granitoid (Fig. 4a), provide diagnostic kinematic indicators that allow documentation of displacements in the suture zone.

Fig. 4. Suture zone and WAC margin rocks: (a) mylonitic granitoid showing feldspar porphyroclasts with variably developed tails; (b) Kpong complex rocks showing layering on all scales (exemplified by recessive carbonatite and resistant nepheline syenite), and xenoliths in various stages of digestion); (c) high-P mafic granulites showing discontinuous shear layering consisting of Grt-rich and Hbl-rich layers, veins with hornblende tadpoles and, in detail (d), with Grt megacrysts typical of those with rutile exsolution rods (also note asymmetric pressure fringe).

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Relations within deformed alkaline rocks and carbonatite (Kpong Complex) Rocks of the alkaline gneiss complex extend for over 60 km along the suture zone from south of Somanya to NE of Pore (Fig. 2) and include nepheline syenite and carbonatite, which together are referred to as the Kpong complex. This complex is typically less than 100 m thick, consists of alternating layers of carbonatite and nepheline gneiss, and is intensely deformed by shearing. In the field, the carbonatite, which is composed essentially of calcite and biotite, occurs throughout the alkaline rock complex as recessed layers of varying thickness (Fig. 4b) or as discrete, lensoid bodies infolded with the more resistant nepheline gneiss. West of the Volta River (Fig. 2), the carbonatite is a dark brown, fine- to medium-grained, variably foliated rock, whereas east of the river, coarse-grained, dark grey porphyritic varieties are more common. Grain size and colour variation appear to be related to deformation and are enhanced by weathering. Xenoliths are distributed throughout the carbonatite horizons as disaggregated, angular or rounded blocks of granitoid gneiss, and occasional amphibolites. Many of these have disintegrated into smaller fragments consisting of mineral clusters, ranging in size from pebbles to boulders of up to 50 cm across. The largest xenoliths preserve strong foliation akin to the mylonitic granitoids, suggesting that they were derived from the Ho gneiss. In addition, partially assimilated xenoliths include rounded to drawn-out mechanical inclusions of feldspars as well as nepheline gneiss. These relationships suggest that the carbonatite intruded into the WAC margin after the emplacement of nepheline syenite; a relation similar to that for alkalic complexes elsewhere (e.g. Philpotts 1990, p. 301). Nepheline-bearing gneisses are the dominant rocks in the Kpong complex and have lateral continuity along strike from the SW of Kpong to areas NE of Alabo River. Three main mineralogical varieties of nepheline-bearing rocks have been distinguished based on abrupt to gradational changes in mineral proportions (Holm 1974): leucocratic nepheline syenite, mafic nepheline gneiss and rare feldspar-poor nepheline gneiss. Nepheline syenite, the dominant type, is bluish grey on fresh surfaces (Fig. 4b) with textural varieties that include fineto medium-grained zones with strong foliation, coarser grained gneissose zones that are typically dark grey in weathered surface, and locally porphyritic varieties exposed east of the Volta River. The foliation is generally subparallel to the alternating carbonatite and nepheline syenite layering, as is evident in carbonatite and nepheline syenite layers that are apparently transposed along the foliation.

Mafic nepheline gneiss is a garnet-bearing rock that is restricted to the contact zone with mafic garnet granulites. It occurs in isolated outcrops in the NE of the map area, where it is typically folded with steep axial surfaces, subvertical hinge zones and asymmetrical limbs. The dark colour, coarse texture and significant modal content of garnet and pyriboles make the mafic alkaline rock conspicuous in the bluish grey nepheline gneiss and carbonatite. It is here interpreted as an alkaline facies of the Shai Hills gneiss unit that formed by alkali metasomatism along the tectonized suture zone; the metasomatic origin is supported by nepheline-rich veinlets in the mafic alkalic gneiss.

Relations within the high-pressure granulites (Shai Hills Gneiss unit) Rocks of the Shai Hills Gneiss unit are folded into west- and SW-verging nappes and crop out in inselbergs of the Accra Plains, which include the north– south-trending Shai Hills (SH, Fig. 1). Most of the isolated hills are asymmetrical with steep, west-facing scarp slopes such as at Krobo Hills (KB) and the prominent Osu Yongwa (OY; Fig. 2). Several quarries in Krobo and Shai Hills and in the low hills south of SH provide access to fresh outcrops for detailed sampling. In these quarries the distinctive features of the rocks exposed are the prominent modal layering and extensive veining. The layering is discontinuous and consists of alternating garnet-rich and hornblende-rich zones, which give the rock a streaky appearance (Fig. 4c) and are interpreted to be shear induced. The veins occur in all sizes and orientations to the tectonic layering; some prominent veins attain a thickness of 0.5 m and together with the thin veins, which are only a few millimetres thick, are estimated to make up to 10% of the rock by volume (Burke 1959). The thickest veins, which are typically folded and locally conformable with the modal layering, are composed of plagioclase, hornblende and scapolite. Hornblende porphyroclasts with well-developed s- and d-type tails in the thick veins, together with asymmetrical pressure fringes around garnet porphyroblasts (Fig. 4d) form spectacular sets of kinematic indicators, which indicate SSW-directed thrusting of the suture zone rocks south of the SH locality (Fig. 1) onto the WAC margin. Despite the intense deformation, primary layering can be inferred from quarry floor exposures at SH where an ultramafic layer (composed essentially of garnet and diopside) is in contact with a layered hornblende, and garnet–plagioclase rock. The inferred primary layering strikes NNE, which is strongly oblique to the WNW strike and NE dip of the streaky foliation south of SH. The


interference of this NNE-striking layering and WNW shear-induced foliation may explain the salient and re-entrant boundary in the SH to KB segment of the suture zone (Fig. 1). The suture zone garnet granulites and garnet amphibolites exposed in the OY locality (Fig. 2) appear generally less deformed, as the shearinduced, streaky foliation so characteristic of the garnet mafic granulites in the quarries at SH and KB is less evident. This may reflect the strain gradient within the suture zone, where the most intensely deformed zones are close to the bounding thrust shear zones. The two main lithological varieties that appear to dominate the exposures at OY are white-weathering, hornblende-poor, garnet– plagioclase-rich rocks and garnet amphibolites with a mottled weathered surface, typical of dyke rocks. In the plagioclase-rich rocks, garnet megacrysts are up to 2.5 cm across in the garnet-rich layers (3–4 cm thick), which alternate with thinner (,1 cm) plagioclase-rich layers with finer garnet. The garnet amphibolites are also distinguished by low garnet modes (typically ,10%) and appear to be the dominant lithology in the OY locality. Within the suture zone, but distinctly different from the typical garnet-bearing Shai Hills gneiss unit, are bodies of mafic and ultramafic rocks interpreted as sills and dykes, which apparently represent late intrusions (Fig. 2). They are distinguished by being typically undeformed or only weakly deformed, variably metamorphosed and include metanorites characterized by the development of corona structures around primary orthopyroxene (Attoh 1998b; Agbossoumonde et al. 2004). The mafic and ultramafic intrusive suite includes layered complexes, norites, pyroxenite and hornblendite sills in the Shai Hills area.

Carbonatite intrusion and continental rifting Petrography Modally, the carbonatite consists of calcite (35– 45%), biotite (25–40%) and variable amounts of alkali feldspar (5–20%) and nepheline (2–20%). The modal calcite content is low compared with carbonatites elsewhere, in which calcite may be up to 75%, but the absence of other carbonates minerals is typical of ancient carbonatites (Woolley & Kemp 1989; Philpotts 1990, p. 300). The calcite typically occurs as subhedral to euhedral megacrysts set in a foliated matrix of fine calcite and biotite. Selected minerals in the carbonatite were analysed using the electron microprobe (EMP) facility at Brigham Young University, where

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major element concentrations were determined using wavelength-dispersive spectrometry (WDS), and synthetic and natural mineral standards. Calcite compositions from two representative samples of carbonatite are listed in Table 1; they show little compositional variation with CaO concentrations in the range 49–54 wt%, negligible MgO contents, significant but variable SrO contents (0– 1.4 wt%), and low concentrations of MnO and FeO. FeO contents are highly variable, up to c. 1 wt%, whereas the MnO contents are less variable and are generally ,0.45 wt%. Analysed calcite compositions can be represented by Ca0.98Sr0.014 – 0.015Fe0.004 – 0.006Mn0.003 – 0.006CO3. Biotite is dark brown, locally occurs as overgrowths on nepheline, and is annitic with low values of Al and Mg and high FeO contents yielding an Fe/Mg ratio of c. 3 (Table 1). K2O contents are less variable, between 9.14 and 9.54 wt%, but TiO2 contents range from 2.95 to 4.07 wt%. The limited compositional variation of biotite can be expressed as K0.99 Fe1.5 – 1.7Mg0.6 – 0.7Si2.6Al1.6 – 1.7O10(OH,F0.1 – 0.4). Alkali feldspar occurs as dynamically recrystallized grains characterized by deformed albite twin lamellae, with most grains composed of nearly pure albite (Ab99) but grains with Or content up to 22 mol% also occur. Diagnostic cancrinite alteration occurs around nepheline, which shows limited compositional variation in the range K0 – 0.42 Na1.4 – 2.0Ca0 – 0.3Al1.8 – 1.95Si2.6O8.

Geochemistry of carbonatite Determinations of major and trace element concentrations in carbonatite by X-ray fluorescence (XRF) analysis and inductively coupled plasma mass spectrometry (ICP-MS) were carried out in Utah State University and the results for two representative samples are listed in Table 2. They are distinguished by low SiO2 contents (30–33 wt%) typical of alkaline rocks; however, the silica content of the Kpong complex carbonatites is relatively high, compared with carbonatite lavas from the East African Rift volcanoes (Woolley & Kemp 1989; Dawson et al. 1995). Figure 5a is a plot of Kpong complex carbonatite and nepheline syenite compared with modern carbonatites in the compositional space used to investigate carbonate– silicate liquid immiscibility (Kjarsgaard & Hamilton 1988). It shows that the Dahomeyide samples plot just inside the two-liquid field, between silicate and carbonate liquids, whereas the nepheline gneiss samples, based on a representative analyses (Table 2) together with the dataset from published analyses (e.g. Junner & Harwood 1928, listed by Holm 1974) and Oldoinyo carbonatite lava (Dawson et al. 1995), plot respectively in the Na –K-rich and Si–Al-rich immiscible liquid

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Table 1. Calcite and biotite compositions from carbonatite Calcite Sample: Analysis no.: FeO MnO MgO CaO SrO

PN32

PN37

1

2

3

1

2

3

4

0.03 0.01 0.00 52.64 0.00

0.31 0.24 0.00 53.28 1.38

0.26 0.26 0.00 53.64 0.00

0.40 0.43 0.10 54.93 1.38

0.38 0.41 0.03 52.54 1.30

0.30 0.36 0.05 50.86 1.25

0.30 0.43 0.04 53.00 1.36

Biotite Sample: Analysis no.: SiO2 TiO2 Al2O3 FeO MnO MgO Na2O K2O BaO F Total Si Ti Al Fe Mg Mn K Na Ba F

PN32B

PN37

1

2

1

2

3

33.38 2.95 18.12 23.43 0.33 5.98 0.09 9.63 0.20 0.25 94.35 2.64 0.16 1.69 1.55 0.70 0.02 0.97 0.01 0.012 0.11

33.03 2.91 18.02 24.34 0.41 5.83 0.08 9.69 0.23 0.23 94.84 2.67 0.17 1.69 1.62 0.69 0.03 0.98 0.01 0.014 0.10

33.25 3.76 17.14 25.20 0.35 5.20 0.15 9.21 0.73 0.35 95.32 2.64 0.22 1.60 1.67 0.62 0.02 0.93 0.02 0.046 0.14

33.45 3.93 17.15 25.00 0.43 5.09 0.13 9.26 0.80 0.35 95.30 2.64 0.23 1.60 1.63 0.60 0.03 0.94 0.02 0.048 0.14

32.19 4.23 17.27 24.66 0.36 4.95 0.17 9.31 0.78 0.39 95.04 2.62 0.25 1.62 1.64 0.59 0.02 0.95 0.03 0.047 0.14

fields. Minor and trace element contents provide strong support for the magmatic origin of the carbonatites; for example, the high Sr (4100– 4500 ppm), Ba (3600–3900 ppm), Zr (267 –283 ppm) and Nb (100–135 ppm) contents are typical of alkaline igneous rocks, as are the high TiO2 and P2O5 contents. The plot of Sr v. SiO2 (Fig. 5b) for the carbonatites compared with the other alkaline rocks in the Kpong complex shows clearly their distinct, high Sr concentrations relative to nepheline syenite; however, these are still significantly lower than Sr contents of natro-carbonatite lavas, which may be up to 1– 1.6 wt%. The high total REE concentrations, as well as the strong fractionation of light REE (LREE) (Fig. 5c), provide further, unmistakable support for mantle-derived carbonatite magma as the source of the Kpong complex carbonatites. Evidently, these high LREE values

(LaN .800) and the steep REE pattern (La/Yb 60) are not compatible with sedimentary carbonate or any likely metasomatic origin. Fourcade et al. (1996) have described carbonatites associated with fenites in the Palaeoproterozoic granulites of the Hoggar, where they suggested that crust-contaminated, mantle-derived metasomatic fluids could have resulted in the highly enriched trace element contents of such rocks. In the Hoggar area, the occurrence of marbles provides an obvious source of the carbonic fluids, which is not the case in the Dahomeyides. Also, key trace elements such Ba and Rb are significantly lower in these carbonatites than in the Kpong complex; for example, in the Hoggar carbonatites, Ba content is 290–1200 ppm compared with 3600–4000 ppm in the Kpong complex rocks (Table 2). Similarly, whereas the Rb concentration


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Table 2. Major and trace element analyses of representative carbonatite (PN32, PN37) and nepheline syenite (KP61C)

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 CO2 Total Nb Zr Y Sr Ba Rb Hf Ta Pb Th U Sc V Cr Ni Cu Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

PN32B

PN37

KP61C

33.52 1.64 12.45 11.68 0.30 2.27 19.08 6.25 4.62 0.60 n.d. 92.41 101 267 40 4142 3637 150 0.21 4.05 13.92 2.11 0.53 38 180 12 13 16 107 142.02 274.18 29.40 99.99 15.28 4.55 11.07 1.32 6.56 1.22 2.98 0.37 2.11 0.32

30.42 2.35 11.35 14.19 0.41 3.57 27.07 3.60 5.16 0.55 n.d. 98.67 135 283 47 4482 3975 140 0.43 6.42 4.05 2.94 0.60 36 239 23 7 9 116 204.70 354.85 37.89 122.65 17.87 5.31 13.50 1.61 7.88 1.45 3.58 0.45 2.65 0.40

40.60 1.06 17.10 5.90 0.20 1.68 11.80 4.66 3.51 0.51 9.50 95.99 n.d. n.d. n.d. 3500 2984 104 3.18 4.23 n.d. 2.91 2.37 2.62 n.d. 24.88 17.83 n.d. n.d. 133.05 240.7 n.d. 88.69 11.60 2.75 n.d. 1.86 n.d. n.d. n.d. n.d. 1.86 0.25

n.d., not determined.

in the Dahomeyide carbonatite is c. 155 ppm, those in the Hoggar metasomatic carbonate rock samples are c. 3–28 ppm. Korsakov & Herman (2006) also presented trace element data for carbonate melt inclusions associated with UHP carbonate rocks but detailed comparison with the Dahomeyide carbonatites shows major differences. The chondrite-normalized LREE ratios of the carbonate melts are ,200 compared with 800–1000 for the

Fig. 5. Geochemical plots of carbonatites: (a) SiO2 þ Al2O3 – Na2O þ K2O–CaO, (b) SiO2 v. Sr and (c) chondrite-normalized REE (symbols labelled in figure).

carbonatites in this study, and the Ba contents are significantly lower, generally ,200 ppm.

Age of carbonatite intrusion Agyei et al. (1987) reported Rb–Sr and K –Ar ages for the carbonatite; their whole-rock Rb–Sr analysis, which yielded 590 + 250 Ma, may be interpreted as allowing a maximum crystallization age less than c. 840 Ma. They also reported mineral ages that have smaller analytical errors and indicate younger ages; the biotite K –Ar age of 572 + 15 Ma is older than the Rb–Sr biotite age of 545 + 11 Ma but is close to the 40Ar – 39Ar hornblende ages of 582–587 Ma for the mafic granulites (Attoh et al. 1997). The latter have been interpreted as metamorphic cooling ages. Preliminary U –Pb analysis

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of zircon separated from a sample of nepheline syenite (D. Hawkins, pers. comm.) defined a reverse concordia from 2109 + 13 Ma to 590 + 4 Ma. The 2109 Ma age is clearly that of the WAC basement whereas the younger intercept age of 590 Ma, in the light of the other age data, could be either the metamorphic or magmatic age of the Kpong complex. This younger age has been confirmed by new U– Pb analyses of zircons separated from carbonatite and nepheline syenite samples (Nude et al. 2006). The zircons from the two samples yielded indistinguishable ages of 592– 594 + 4 Ma, which are interpreted as the time of intrusion of the Kpong complex. Thus the available age data are compatible with carbonatite magmatism associated with continental rifting but apparently this occurred later than UHPM in the Dahomeyides.

HP and UHP metamorphism Petrography of garnet mafic granulites The characteristic petrographic feature of the Shai Hills unit in the suture zone is the high modal abundance of garnet, up to c. 25 vol%, coexisting with diopsidic pyroxene and scapolite. Variable mineral proportions allow the following petrographic types to be distinguished (Attoh 1998a): hornblende-rich granulite, garnet-rich granulite and garnet –diopside rock. The mineral assemblage of the hornblende-rich granulite is Hbl þ Pl þ Grt þ Di þ Qz + Sc þ accessory minerals (Sph + Ap + Ilm + Ru) and typical mineral proportions are 42% hornblende, 38% plagioclase, 9% garnet, 4% diopside and 5% quartz. Garnet-rich granulites have the same mineral assemblage but with different mineral proportions of 29% garnet, 26% plagioclase, 20% diopside, 9% hornblende, 10% quartz and 2% scapolite. The diopside-rich granulite sample is an ultramafic rock composed of Di þ Grt + Hbl + Pl þ opaque accessory minerals, which display complex exsolution lamellae of rutile and ulvospinel in ilmenite. Plagioclase-rich garnet granulites from OY are similar to the garnet granulites from SH but with higher plagioclase content, estimated to reach .30%, and with less than 5% Hbl. The garnet amphibolites are inferred to be derived from gabbroic dyke rocks and are characterized by low modal content of garnet in the assemblage Hbl + Pl + Grt + Di þ Sc + accessories.

Thermobarometry and age of HP metamorphism The P –T conditions under which the mafic garnet granulites and eclogites of the Dahomeyide suture

zone recrystallized have been reported in recent studies (Attoh 1998a, b; Agbossoumonde et al. 2001). These estimates were derived from thermobarometric calculations based on mineral compositions and were constrained by mineral assemblage data from selected samples including eclogites with the mineral assemblage Grt þ Di + Zo. The results indicate that peak metamorphic temperatures c. 800 8C were attained at pressures of 14–15 kbar although some workers (e.g. Meńot & Seddoh 1985; Agbossoumonde et al. 2004) have suggested pressures as high as 18 kbar for the suture zone rocks in Togo. These results, and all available geological data, suggest that the Dahomeyides preserve eclogitic high-pressure (EHP) metamorphic rocks, which recrystallized during a collision event related to subduction on the WAC margin. The age of this event is well constrained by mineral ages: these include U –Pb zircon age of 610 + 2 Ma (Attoh et al. 1991), 603 + 5 Ma from U –Pb zircon analyses by Hirdes & Davis (2002) and 613 + 1 Ma age from zircon Pb– Pb data (Affaton et al. 2000). These ages are in good agreement with the c. 620 Ma age determined for eclogites and related UHP rocks in the Trans-Saharan belt of northern Mali (Jahn et al. 2001).

UHP metamorphism The Shai Hills mafic granulites also preserve evidence of UHP metamorphism in the form of rutile exsolution lamellae in garnets (Fig. 6). The garnets that contain this microstructure are megacrystic (1–2 cm across) and are typically subhedral to euhedral. Two samples of garnet granulites with this texture (SH6 and SH23C) are from quarries in the SH locality: SH6 is a hornblende-rich granulite with isolated euhedral garnet megacrysts up to 2 cm across set in the mineral assemblage Hbl þ Pl þ Grt þ Di + Qz + Sc. SH23C is representative of a sample of ultramafic layer with abundant garnet megacrysts set in a matrix with the mineral assemblage Grt þ Di þ Hbl þ Pl + Sc + Qz and secondary calcite. Accessory opaque minerals in both samples are rutile and ilmenite. The garnet megacrysts are distinguished by the exsolved rutile needles arranged in a rectilinear pattern (Fig. 6) but the rims of some garnet grains are devoid of rutile exsolution rods (Fig. 6a). In plane-polarized light, but especially under cross-polarized light, this microstructure is striking because the birefringent rutile needles stand out in contrast to the isotropic garnet background (Fig. 6b and c). Backscattered-electron (BSE) imaging reveals the rutile rods to be ,3 mm in diameter, ,50 mm long and typically spindle shaped, and energy-dispersive spectrometry confirmed that they are composed of


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Fig. 6. Photomicrographs of rutile exsolution in garnets indicating UHP metamorphism: (a) SH23C; (b) detail of SH23C (under cross-polarized light (XPL)), (c) SH6 under XPL; (d) the same area of SH6 in plane-polarized light.

TiO2. Although previous workers (von Knorring & Kennedy 1958) noted this microstructure and described it as ‘lamellar and rod-like intergrowth in garnet’ in samples from this locality they were unable to identify the ‘rod-like intergrowth’ and did not recognize the significance of the microstructure. The key observation in this microstructure is the apparent, crystallographically controlled growth of the exsolution rods as evident in the rectilinear pattern; this suggests transformation in the

garnet crystal structure involving substitution of TiIV in the precursor garnet structure (Zhang et al. 2003). So far, such rutile-in-garnet microstructure has been reported only from UHP rocks (e.g. Zhang & Liou 1999, and references therein) and this observation has led to experimental studies of Ti solubility in UHP phases such as garnet and clinopyroxene (Zhang et al. 2003). For garnet, the basic theory behind the experiments is that majorite

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is the stable component of garnet in UHP rocks and that this transforms to pyralspite –ugrandite garnet components plus rutile needles with decreasing pressure (P) by the reaction M3 ðMTiÞSi3 O12 þ Al2 ðin CaAl2 SiO6 Þ þ SiO2 Majorite Grt

Al in Cpx

¼ M3 Al2 Si3 O12 þ TiO2 þ CaMSiO6 Grt

Rutile

where M ¼ Ca, Mg, Fe. The prograde (reverse) reaction can be expressed by the exchange CaTi ! 2Al, which results in an increase in Ca, Ti and Si, the majorite components, with increasing pressure. From the experimental results, Ti solubility in garnets may be up to 4.5 wt% TiO2 and is strongly dependent on pressure in the range 5–15 GPa. Thus the abundant exsolved rutile rods in the garnet of the Shai Hills samples indicate precursor majoritic garnet and can be used to estimate the peak pressures attained during UHP metamorphism. Compositions of garnets with exsolution microstructures are listed in Table 3. The samples were analysed at the EMP facilities of the University of Chicago and Cornell University, where major element concentrations were determined by WDS. Garnet with exsolved rutile in SH23C shows variable Ti content from minimum of 0.04 wt% in the rim (R) to a maximum of 0.21 wt% near the grain centre (C) but rim compositions as high as 0.14 wt% are also noted. There is no apparent correlation

between these Ti contents and Ca, Al or Si. On the other hand, the TiO2 content of garnet in SH6 suggests some correlation with Ca and Al, as Ti content of 0.002–0.006 moles in the structural formulae corresponds to 0.72–0.63 moles of Ca compared with Ti content of 0.06 moles and corresponding 0.60 moles of Ca, whereas the apparent correlation between Ti and Al is the inverse. The high Ti of 0.06 mole in the garnet mid (M) domain of SH6 is anomalous and probably represents a spot analysis close to a rutile needle, whereas the 0.002 moles of Ti in the garnet grain rim (R) is not different from background. These low and variable Ti concentrations in garnets with exsolved rutile rods are significant because they suggest that if exsolved rutile is integrated, the garnet structural formulae will require major adjustment for the majoritic component. Table 4 lists the rutile content (mode) of the garnet, which was obtained by digital image analysis. From the mode of exsolved rutile we calculated rutile content (wt%) by taking into consideration the density differences between the two phases. The TiO2 concentrations (wt%) in the inferred precursor majorite component of the garnets were estimated from the rutile contents and are listed as well as the pressure estimates based on those concentrations (Table 4). Estimated majoritic component for SH23C is c. 7.5 mol%, corresponding to the garnet formula Fe1.42Mg1.14Mn0.05Ca0.47 [TiSi]0.24 – 0.26Al1.70 – 1.72Si2.98O12. In SH6 the estimated majoritic component of garnet is c. 6 mol%, giving a formula of Fe1.39Mg0.93Mn0.06 Ca0.72[TiSi]0.22 – 0.26Al1.71 – 1.75Si2.96O12. These

Table 3. Compositions of representative garnets with rutile rods Sample: Position: SiO2 TiO2 Al2O3 FeO MnO MgO CaO Total Si Ti Al Fe Mn Mg Ca

SH6

SH23C

R

M

MC

R

M

C

38.90 0.03 21.47 21.53 0.87 8.09 8.75 99.64 2.99 0.002 1.95 1.39 0.06 0.93 0.72

39.45 0.11 21.28 22.11 1.28 8.61 7.77 100.61 3.00 0.006 1.91 1.41 0.09 0.98 0.63

38.73 0.99 21.36 22.28 1.16 8.51 7.38 100.86 2.96 0.06 1.93 1.43 0.08 0.97 0.63

39.50 0.04 22.25 22.60 0.85 10.13 5.82 101.19 2.97 0.002 1.97 1.42 0.05 1.14 0.47

39.67 0.09 22.17 21.99 0.79 10.09 6.55 100.35 2.98 0.006 1.96 1.38 0.05 1.13 0.53

39.70 0.21 21.99 21.66 0.85 10.15 6.31 100.69 2.98 0.012 1.95 1.36 0.05 1.14 0.53

R, rim; M, mid; C, centre.


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Table 4. TiO2 concentrations and P estimates from rutile contents of garnets Sample

Rutile (vol%)

Rutile (wt%)

TiO2 (wt%)

P (GPa)

Majorite (mol%)

SH23C SH23C SH6 SH6

7.8 8.5 5.8 7.1

9.6 10.4 7.2 8.7

2.01 2.2 1.5 1.8

6.0 – 7.2 6.2 – 7.5 4.0 – 6.0 5.0 – 7.0

7.2 7.8 5.4 6.6

majorite contents are similar to the c. 5.2 mol% calculated for the experimental samples at 4–5 GPa and 1000 8C (Zhang et al. 2003; Table 4). Pressure estimates were also obtained directly from the experimental solubility data of Zhang et al. (2003); the results indicate P .4 GPa, which is much higher than the current thermobarometric estimates of 15– 18 kbar from phase compositions and assemblages of the host rocks. This pressure is also higher than the 27 kbar that was inferred for UHPM rocks in the Trans-Saharan belt segment in Mali (Caby 1994). The principal sources of uncertainty in the P estimate include analytical errors in the modal counts of rutile rods in garnet, and the effect of lower temperatures (T ) on Ti solubility. For example, estimates of rutile content of garnets probably represent maximum values because they were determined preferentially for areas densely populated by rutile rods (Fig. 6); as such, the TiO2 contents of the garnets are probably an overestimate, and alternative reasonable values could be much lower if garnet domains with few rutile rods are considered. This would result in lower P estimates. On the other hand, the experimental results of Zhang et al. (2003) suggested that Ti solubility increases with temperature, so the effect of the lower T of recrystallization of the samples is to increase the P estimates (i.e. high Ti in low-T rocks could indicate even higher pressure). These considerations suggest uncertainties of c. 1–2 GPa in the pressure estimate in Table 4, which might indicate maximum pressures of c. 3 GPa.

Qz equilibria calibrations of Ellis & Green (1979) and Eckert et al. (1991)). These pressures are similar to those previously reported and consistent with the significant jadeite component of the diopsidic pyroxene. In these calculations, the garnet grain centre –mid compositions gave slightly higher temperatures and pressures than the garnet rim compositions, suggesting a recrystallization path of decreasing pressure and temperature with garnet growth as indicated at SH (Fig. 7). A similar overall path was inferred for the other samples from the suture zone (Attoh 1998b). If correct, such a path may be a segment of overall P– T trajectory along the UHP decompression path segment (Fig. 7). The other point on the P–T path (LT) is largely based on the eclogite thermobarometry

P – T path The compositional parameters of garnets with rutile exsolution (Table 3) together with published estimates of EHP metamorphism in the suture zone rocks can be used to infer possible P –T paths. For SH23C, garnet compositions in Table 4 are combined with compositional parameters of coexisting diopside (Na0.18Ca0.79Mg0.63Fe0.24Al0.31 Si1.88O6) and plagioclase (Na0.71Ca0.27K0.01Al1.28 Si2.72O8) to calculate pressures and temperatures. These yielded metamorphic recrystallization temperatures of c. 800 8C and pressures of 13 –14 kbars (based on the Grt–Cpx and Grt –Cpx–Pl–

Fig. 7. Inferred P –T path of HP to UHP metamorphism, showing segments corresponding to inferred path for Mont Lato eclogite (LT; see Fig. 1) and Shai Hills HP (SH) and UHP rocks. Phase boundaries from Liou & Zhang (2000).

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reported by Attoh (1998a) and Agbossoumonde et al. (2001, and references therein). Although the inferred path shown in Figure 7 is somewhat speculative, it is, as noted, supported by three reliable pieces of data: (1) an eclogite-facies metamorphic record characterized by low-T and high-P (samples from LT and HZ in Fig. 1); (2) the UHP metamorphic record inferred from rutile-in-garnet microstructure; (3) HP metamorphism typical of the mafic garnet granulites of the suture zone documented in previous studies at SH, KB and AD (Fig. 1). Thus the only new segment of the path is the UHP segment that is proposed here.

Discussion Tectonic implications for Rodinia breakup and Gondwana assembly The palaeogeographical coordinates of the WAC in Rodinia supercontinent reconstructions are unconstrained (e.g. Hoffman 1991; Condie 2003; Cordani et al. 2003) and this is commonly addressed, or rather avoided, by the typical peripheral positions in nearly all the reconstructions in which WAC is attached to Amazonian craton. Cordani et al. (2003) have reconsidered the available data, together with some new data, and concluded that the Amazonian craton (AMC), with WAC tenuously tethered to it, was part of Laurentia-cored eastern margin of Rodinia (see Cordani et al. 2003, fig. 4). They inferred that a large ocean (Brasiliano) separated these NW Gondwana cratons from other African cratons; in particular, the Congo–Saõ Francisco (CSFC) and a postulated Borborema Trans-Saharan (BTS) craton. Thus the vagrancy of the WAC remains an unresolved problem. This is further compounded by the uncertainty regarding the extent of the proposed ocean basin that surrounded the WAC prior to incorporation into Gondwana. For example, Condie (2003) has proposed a rift system surrounding the WAC that evolved into a short-lived, limited-extent Brasiliano ocean during the Rodinia breakup interval. Evidently, key missing information for better constrained reconstructions involving the WAC is documentation of its breakup from Rodinia and the chronology of those events. To date, such information has not been available; therefore, the record of deformed carbonatite as an indicator of rifting on the WAC margin is of crucial significance, which, combined with the preservation of UHP rocks, should provide an important constraint on the vagrancy of the WAC. In this case, however, the age of the deformed carbonatites does not appear to provide such a constraint (Nude et al. 2006).

Dahomeyide record of Gondwana assembly Hypothetical Wilson cycle tectonic scenarios for the WAC margin are shown in Figure 8, where the chronologies are constrained in the Dahomeyides only after c. 650 Ma. Although initial rifting might have been accompanied by alkaline magmatism, such a suite must be older than the Kpong complex and, to date, there is no record of it on the eastern margin of the WAC. Consequently, the duration and extent of rifting are not known (Fig. 8a), but it is speculated that lithospheric breakup resulted in the formation of an ocean basin on the eastern margin of the WAC (Fig. 8b). This is consistent with the current interpretation of the geology in southeastern Ghana, which suggests that, immediately east of the suture zone, 2.1 Ga (Birimian) rocks are not preserved; rather, an inferred juvenile crust now composed of granitoid gneiss complexes (Fig. 3) was accreted to the suture zone (Attoh et al. 1997). These arc rocks may be correlated with the Tilemsi arc in northern Mali (Caby et al. 1989), which has been dated at 726 + 7 Ma by U– Pb zircon data. This is older than U– Pb zircon age of 640 + 50 Ma for eclogites reported by Bernard-Griffiths et al. (1991) and interpreted by them to be the age of the protolith of the suture zone rocks. The age of HP (including EHP and UHP) metamorphism is, on the other hand, well constrained by U –Pb zircon analyses to be between 612 and 605 Ma, which is similar to the 620 Ma reported for the UHP metamorphism in Mali by Jahn et al. (2001). With the age of carbonatite magmatism determined to be 592– 594+4 Ma (Nude et al. 2006), it is evident that the DARCs of the Kpong complex were not emplaced during continental rifting related to the breakout of the WAC from Rodinia. Nd and Hf isotopic analysis of garnet granulites (Attoh & Schmitz 2005) also provides an upper limit for the magmatic arc protoliths of the suture zone mafic rocks; the Hf isotopic data yielded TDM of 970– 770 Ma, which we interpret as indicating a minimum age of rifting on the WAC at c. 770 Ma. Such an age is permissive of the postulated timing of Rodinia breakup (Hoffman 1991) around 750 Ma and suggests that the breakout of the WAC occurred some 100–150 Ma before oceanic closure along the margin. Figure 8c and d suggests that HP metamorphism was followed by carbonatite intrusion and deformation accompanied by exhumation. It also depicts the collision of the WAC with an inferred oceanic island arc complex following the consumption of the oceanic lithosphere. The collision depicted may account for EHP metamorphism but UHP metamorphism requires the initial collision to be accompanied by subduction of the suture zone rocks to mantle depths, and probably collision


229

Fig. 8. Wilson cycle reconstructions based on the geology of the Dahomeyides in southeastern Ghana.

with an eastern continental lithosphere was required (not shown). As the protoliths of the suture zone HP rocks preserve geochemical signatures of MORB and oceanic island arcs (Attoh & Morgan 2004), it is likely that at least some of the material subducted and exhumed from the mantle represented the roots of the island arc (Fig. 8c). In the setting of NW Gondwana, during the late Neoproterozoic, the Dahomeyide margin of the WAC may have collided with an exotic craton from the east, such as the Saharan metacraton (Abdelsalam et al. 2002) or its extension to South America, the BTS (Cordani et al. 2003). As in other UHP metamorphic terranes, the processes leading to the exhumation of the Dahomeyide UHP metamorphic rocks are very uncertain (e.g. Hacker & Peacock 1995; Liou & Zhang 2000).

Significance of the Dahomeyide suture zone in southeastern Ghana The association of carbonatite and UHP metamorphic rocks, which are respectively diagnostic

of continental rifting and continental subduction, in the Dahomeyide suture zone provides the setting to study these processes in Gondwana assembly. The lithotectonic association described here is distinct from the other Pan-African alkaline rock complexes such as those reported from the Iforas (Boullier et al. 1986, and references therein), which were interpreted as products of posttectonic, within-plate magmatism. In this sense, the Dahomeyide segment preserves a crucial record of the Wilson cycle; namely, the potential to document the breakup of continental lithosphere leading to the creation of an ocean basin and the final stage, involving the closure of the ocean basin. Key tests of these ideas have yet to be devised but must include studies of rock suites in the Dahomeyide segment. One such test is the extent of juvenile crust production east of the postulated suture zone in the rocks exposed within the reworked Pan-African zone of the BTS comprising the Benin –Nigerian shield (Ferre´ et al. 2002). In this paper, we have shown that the postulated linkage between alkaline magmatism, inferred to

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herald the breakup of the WAC lithosphere from Rodinia, and the HP– UHP metamorphic record related to the collision of the WAC during Gondwana assembly cannot be readily supported by available geochronological data. Although HP metamorphic rocks in the Dahomeyides have been documented by previous studies, and maximum pressures of metamorphism confirmed to be in excess of 16 kbar, this is the first report of evidence for UHP metamorphism in this segment of the Trans-Saharan orogen. To date, however, the occurrence of diagnostic UHP metamorphic minerals such as coesite has not been reported, but if the interpretation presented is correct, it predicts the future discovery of such minerals. On the other hand, in the absence of such corroborating evidence it is possible that the crystallographically controlled exsolution of rutile in garnet may indicate lower than UHP metamorphic conditions, although this requires further experimental verification. The research collaboration reported here was made possible by grants from the US Fulbright program (K.A.) and University of Ghana grants for field work (P.M.N.). Reviews by A. Korsakov and K. Ouzegane are much appreciated.

References A BDELSALAM , M. G., L IE´ GEOIS , J.-P. & S TERN , R. J. 2002. The Saharan metacraton. Journal of African Earth Sciences, 34, 119–136. A FFATON , P., R AHAMAN , M. A., T ROMPETTE , R. & S OUGY , J. 1991. The Dahomeyide orogen: Tectonothermal evolution and relationship with the Volta basin. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, New York, 95–111. A FFATON , P., K RONER , A. & S EDDOH , K. F. 2000. Pan-African granulite formation in the Kabye massif of northern Togo (West Africa): Pb– Pb zircon ages. International Journal of Earth Science, 88, 778–790. A GBOSSOUMONDE , Y., M E´ NOT , R.-P. & G UILLOT , S. 2001. Metamorphic evolution of Neoproterozoic eclogite from south Togo (West Africa). Journal of African Earth Sciences, 33, 227 –244. A GBOSSOUMONDE , Y., G UILLOT , S. & M E´ NOT , R. P. 2004. Pan-African subduction collision event evidenced by high-P corona in metanorites from Agou massif (southern Togo). Precambrian Research, 135, 1–25. A GYEI , E. K., VAN L ANDEWIJK , J. E. J. M., A RMSTRONG , R. L., H ARAKAL , J. F. & S COTT , K. L. 1987. Rb–Sr and K–Ar geochronometry of southeastern Ghana. Journal of African Earth Sciences, 6, 153–161. A TTOH , K. 1998a. High-pressure granulite facies metamorphism in the Pan-African Dahomeyide orogen, West Africa. Journal of Geology, 106, 236– 246. A TTOH , K. 1998b. Models for orthopyroxene– plagioclase and other corona reactions in metanorites, Dahomeyide orogen, West Africa. Journal of Metamorphic Geology, 16, 345–362.

A TTOH , K. & M ORGAN , J. 2004. Geochemistry of highpressure granulites from the Pan-African Dahomeyide orogen, West Africa: constraints on the origin and composition of lower crust. Journal of African Earth Sciences, 39, 201–208. A TTOH , K. & S CHMITZ , M. D. 2005. Nd and Hf isotopic compositions of Pan-African high-pressure mafic granulites. EOS Transactions, American Geophyical Union, 86, Supplement, V13B-02. A TTOH , K., H AWKINS , D. & B OWRING , S. 1991. U– Pb zircon ages of gneisses from the Pan-African Dahomeyide orogen, West Africa. EOS Transactions, American Geophyical Union, 72, S299. A TTOH , K., D ALLMEYER , R. D. & A FFATON , P. 1997. Chronology of nappe assembly in the Pan-African Dahomeyide orogen: evidence from 40Ar/39Ar mineral ages. Precambrian Research, 82, 153–171. B ERNARD -G RIFITHS , J., P EUCAT , J. J. & M ENOT , R. P. 1991. Isotopic (Rb –Sr, U–Pb and Sm–Nd) and trace element geochemistry of eclogites from the Pan-African belt: a case study of REE fractionation during high grade metamorphism. Lithos, 27, 43–57. B OULLIER , A. M., L IE´ GEOIS , J. P., B LACK , R., F ABRE , J., S AUVAGE , M. & B ERTRAND , J. M. 1986. Late Pan-African tectonics marking the transition from subduction-related calc-alkaline magmatism to withinplate alkaline granitoids (Adrar des Iforas, Mali). Tectonophysics, 132, 233–246. B URKE , K. 1959. Replacement veins in the Dahomeyan of Ghana. Geological Magazine, 96, 353–360. B URKE , K., A SHWAL , L. D. & W EBB , S. 2003. New way to map old sutures using deformed alkaline rocks and carbonatites. Geology, 31, 391 –394. C ABY , R. 1987. The Pan-African belt of West Africa from the Sahara Desert to the Gulf of Guinea. In: S CHAER , J. P. & R ODGERS , J. (eds) Anatomy of Mountain Ranges. Princeton University Press, Princeton, NJ, 129–170. C ABY , R. 1994. Precambrian coesite from northern Mali: First record and implications for plate tectonics in the Trans-Saharan segment of the Pan-African belt. European Journal of Mineralogy, 6, 235– 244. C ABY , R., A NDREOPOULOS -R ENAUD , U. & P IN , C. 1989. Late Proterozoic arc– continent and continent– continent collision in the Pan-African Trans-Saharan belt of Mali. Canadian Journal of Earth Sciences, 26, 1136– 1146. C ASTAING , C., T RIBOULET , C., F EYBESSE , J.-L. & C HEVREMONT , P. 1993. Tectonometamorphic evolution of Ghana, Togo, and Benin in the light of the PanAfrican/Brasiliano orogeny. Tectonophysics, 218, 323–342. C ONDIE , K. C. 2003. Supercontinents, superplumes and continemtal growth: the Neoproterozoic record. In: Y OSHIDA , M., W INDLEY , B. F. & D ASGUPTA , S. (eds) Proterozoic East Gondwana Supercontinent Assembly and Breakup. Geological Society, London, Special Publications, 206, 1– 21. C ORDANI , U. G., D’A GRELLA -F ILHO , M. S., B RITO -N EVES , B. B. & T RINDALE , I. F. 2003. Tearing up Rodinia: the Neoproterozoic paleogeography of South American cratonic fragments. Terra Nova, 15, 350–359.

CARBONATITES IN DAHOMEYIDE SUTURE ZONE D AWSON , J. B., P INKERTON , H., N ORTON , G. E., P YLE , D. M., B RONIN , P., J ACKSON , D. & F ALLICK , A. 1995. Petrology and geochemistry of Oldinyo Lengai lavas extruded in November 1988: Magma source, ascent and crystallization. In: B ELL , K. & K ELLER , J. (eds) Carbonatite Volcanism; Oldoinyo Lengai and the Petrogenesis of Natrocarbonatites. Springer, New York, 47–69. D UCLAUX , G., M E´ NOT , R. P., G UILLOT , S., A GBOSSOUMONDE , Y. & H ILAIRET , N. 2006. The mafic layered complex of the Kabye´ massif (north Togo and north Benin): evidence of Pan-African granulitic continental arc root. Precambrian Research, 116, 101– 118. E CKERT , J. O., N EWTON , R. C. & K LEPPA , O. J. 1991. The DH of reaction of garnet–pyroxene–plagioclase– quartz geobarometers in the CMAS system by solution calorimetry. American Mineralogist, 76, 148–160. E LLIS , D. J. & G REEN , D. H. 1979. An experimental study of the effect of Ca upon the garnet–clinopyroxene Fe– Mg exchange equilibria. Contributions to Mineralogy and Petrology, 71, 13–22. F ERRE´ , E., G LEIZES , G. & C ABY , R. 2002. Obliquely convergent tectonics and granite emplacement in the Trans-Saharan belt of eastern Nigeria: a synthesis. Precambrian Research, 114, 199– 219. F OURCADE , S., K IENAST , J.-R. & O UZEGANE , K. 1996. Metasomatic effects related to channelled fluid streaming through deep crust: fenites and associated carbonatites (In Ouzzal Proterozoic granulites, Hoggar, Algeria). Journal of Metamorphic Geology, 14, 763–781. H ACKER , B. R. & P EACOCK , S. M. 1995. Creation, preservation and exhumation of UHPM rocks. In: C OLEMAN , R.G. & W ANG , Xiaomin(eds) Ultrahigh Pressure Metamorphism. Cambridge University Press, New York, 159–181. H IRDES , W. & D AVIS , D. W. 2002. U– Pb zircon and rutile metamorphic ages of the Dahomeyan garnet– hornblende gneiss in southeastern Ghana, West Africa. Journal of African Earth Sciences, 35, 445–449. H OFFMAN , P. F. 1991. Did the breakout of Laurentia turn Gondwana inside-out? Science, 252, 1409–1412. H OLM , R. F. 1974. Petrology of alkalic gneiss in the Dahomeyan of Ghana. Geological Society of America Bulletin, 85, 1441–1448. J AHN , B. M., C ABY , R. & M ONIE , P. 2001. The oldest UHP eclogites of the world: age of UHP metamorphism, nature of protoliths and tectonic implications. Chemical Geology, 178, 143– 158. J UNNER , N. R. & H ARWOOD , H. F. 1928. Microscopical features and chemical analyses of certain

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representative igneous rocks from the Gold Coast, British West Africa. Gold Coast Geological Survey Bulletin, 4, 10–11. K ESSE , G. O. 1985. Rock and Mineral Resources of Ghana. B ALKEMA , B OSTON , M. A., K JARSGAARD , B. A. & H AMILTON , D. L. 1988. Liquid immiscibility and the origin of alkali-poor carbonatites. Mineralogical Magazine, 52, 43–55. K ORSAKOV , A. & H ERMAN , J. 2006. Silicate and carbonate melt inclusions associated with diamond in deeply subducted carbonate rocks. Earth and Planetary Science Letters, 241, 104–118. L IOU , J. G. & Z HANG , R. Y. 2000. Petrological ad geochemical characteristics of ultrahigh-pressure metamorphic rocks from the Dabie– Sulu terrane, east central China. International Geology Review, 42, 328– 352. M E´ NOT , R. P. & S EDDOH , K. F. 1985. The eclogites of Lato Hills, south Togo, West Africa: Relics from the early tectonometamorphic evolution of the Pan-African orogeny. Chemical Geology, 50, 313– 330. N UDE , P. M., C ORFU , F. & A TTOH , K. 2006. U–Pb zircon ages of deformed carbonatite and alkaline rocks in the Pan-African Dahomeyide suture zone, West Africa. EOS Transactions, American Geophysical Union, 87, Fall Meeting Supplement, V31B-0585. P HILPOTTS , A. R. 1990. Principles of Igneous and Metamorphic Petrology. Prentice Hall, Englewood Cliffs, NJ. T ROMPETTE , R. 1994. Geology of Western Gondwana (2000– 500 Ma); Panafrican Aggregation of South America and Africa. Balkema, Rotterdam. VON K NORRING , O. & K ENNEDY , W. Q. 1958. The mineral paragenesis and metamorphic status of garnet– hornblende– pyroxene– scapolite gneiss from Ghana. Mineralogical Magazine, 31, 846– 859. W OOLLEY , A. R. 2001. Alkaline Rocks and Carbonatites of the World: Part 3. Geological Society, London. W OOLLEY , A. R. & K EMP , D. R. C. 1989. Carbonatites: nomenclature, average chemical compositions, and element distribution. In: B ELL , K. (ed.) CarbonatiteGenesis and Evolution. Unwin Hyman, London, 1–14. Z HANG , R. Y. & L IOU , J. G. 1999. Clinopyroxene from the Sulu ultrahigh-pressure terrane, eastern China: Origin and evolution of garnet exsolution in clinopyroxene. American Mineralogist, 88, 1591– 1600. Z HANG , R. Y., Z HAI , S. M., F EI , Y. W. & L IOU , J. G. 2003. Titanium solubility in coexisting garnet and clinopyroxene at very high pressure: the significance of exsolved rutile in garnet. Earth and Planetary Science Letters, 216, 591–601.

Me´langes and ophiolites during the Pan-African orogeny: the case of the Bou-Azzer ophiolite suite (Morocco) ROMAIN BOUSQUET1, RACHID EL MAMOUN2, OMAR SADDIQI2, BRUNO GOFFE´3, ¨ LLER1,5 & ATMAN MADI4 ANDREAS MO 1

Institut fu¨r Geowissenschaften, Universita¨t Potsdam, Karl Liebknecht Strasse 24, 14476 Potsdam-Golm, Germany (e-mail: [email protected])

2

De´partement de geólogie, Universite´ Hassan II—Aı¨n Chock, Route d’El Jadida, B.P. 5366 Maaˆrif, Casablanca, Morocco

3

Laboratoire de Geólogie, UMR 8538, Ecole Normale Supe´rieure Paris, 24 rue Lhomond, 75231 Paris cedex 05, France 4

Akka Gold Mining, MANAGEM-ONA, Rabbat, Morocco

5

Present address: Department of Geology, Kansas University, 1475 Jayhawk Boulevard, Lawrence, Kansas 66045-5276, USA Abstract: Since the discovery of ophiolite sequences, the Bou-Azzer inlier has been considered a key area for understanding the evolution of the northern margin of the West African craton during the Pan-African orogeny. For about 20 years, it had been commonly accepted that the Bou-Azzer inlier represents an accretionary me´lange accreted onto the West African craton under blueschist metamorphic conditions, similar to the Franciscan Complex and the Sanbagawa facies series. This would imply that a low geothermal gradient was prevalent during the subduction of the PanAfrican oceanic plate, and that the ocean was subducted with a high convergence rate. A reinvestigation of the metamorphic conditions by a thermodynamic approach shows that the ophiolite sequence of Bou-Azzer underwent HT greenschist metamorphic conditions instead of blueschist metamorphic conditions. We propose that the ophiolites of Bou-Azzer are not similar to the Sanbagawa facies series or to the Franciscan Complex, but bear similarities to the Albanian or Cyprus ophiolites, which represent dismembered ophiolite sequences overprinted by greenschist conditions.

Ophiolites are remnants of oceanic lithosphere that have been tectonically emplaced onto continents. Well-preserved ophiolite sections consist of (in descending stratigraphic order) pillow lavas, a sheeted dyke complex, gabbro, cumulate ultramafic rocks, and tectonized mantle. Ophiolites provide valuable information on the nature of sea-floor processes, global heat loss, palaeogeographical reconstructions of the continents, and the subduction processes. However, the mechanisms of ophiolite accretion onto the continental margins are debatable. Did subduction and obduction change through time (van der Velden & Cook 1999) or not (Kusky & Polat 1999)? Did the thermal gradient of subduction and orogenic wedge change through time (Maruyama & Liou 1998) or not (O’Brien & Ro¨tzler 2003)? Since the discovery of the first Precambrian ophiolite sequence in the Bou-Azzer inlier (Leblanc 1976), this inlier has played a critical role in the understanding of the evolution of the northern margin of the West African craton (WAC) during the Pan-African orogeny (Leblanc

& Lancelot 1980; Saquaque et al. 1989; Hefferan et al. 2000). This ophiolite, together with the Khzama ophiolite (Sirwa inlier, Admou & Juteau 1998), represents a unique remnant of a PanAfrican ocean, although the exact localization of the Pan-African suture is in question (see discussion by Ennih & Lie´geois 2001; Bouougri 2003). It has been proposed that the Bou-Azzer ophiolites were buried to depths sufficient to generate lower blueschist-facies mineral assemblages in a scrapedoff me´lange (Hefferan et al. 2002). Such a mechanism of accretion of the Bou-Azzer ophiolite onto the WAC is similar to that for accretion of the Franciscan me´lange along the North American west coast. This mechanism implies that a low geothermal gradient was prevalent during the subduction of the Pan-African oceanic plate or that the oceanic plate was subducted at a high convergence rate. Can we really document such conditions for the ophiolite of Bou-Azzer? This study first reviews the mechanisms of ophiolite accretion, formation of me´langes and the


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evolution of the thermal gradients occurring in subduction zones through time, and then reinvestigates the metamorphic P–T conditions experienced by the ophiolite suite of Bou-Azzer, using thermodynamic methods, to re-evaluate the mechanisms of accretion of the ophiolites onto the WAC.

Geological setting The Bou-Azzer inlier is a critical element of the Anti-Atlas area (Fig. 1) during the Pan-African orogeny for the following reasons: (1) it contains outcrops of the northern margin of the WAC and a succession of arc and oceanic crustal components that record a progressive history of deformation and metamorphism in the region; (2) it is unconformably overlain by only slightly deformed Phanerozoic sedimentary sequences, implying that many of the original Neoproterozoic relationships survive; (3) it is intruded by post-tectonic calc-alkaline intrusions. The Bou-Azzer inlier consists of a complex association of rock units occurring within a series of tectonic blocks surrounded by a latest Neoproterozoic to Early Cambrian cover (Leblanc 1981). The tectonic blocks are separated by oblique slip faults that are parallel to the main suture zone with the WAC to the south

(Saquaque et al. 1989) and are believed to represent dismembered parts of a subduction zone complex. South of the Bou-Azzer inlier a deformed platform sequence of quartzites and stromatolitic limestones, which is overlain by basic lava flows and a volcano-sedimentary pile (Leblanc 1975), rests upon basement previously interpreted as Palaeoproterozoic gneiss (c. 2 Ga) and as part of the WAC. However, D’Lemos et al. (2006) suggested a Neoproterozoic age for the whole basement, based on U –Pb dating and Nd isotopic signature of the Tazigzaout gneiss. The passive margin units occur in close proximity to a variety of igneous, meta-igneous, and metasedimentary rocks (Leblanc 1975). These include augen-gneiss and muscovite pegmatite and leucogranite, which, on the basis of lithological similarity and deformation, have been correlated with rocks of the nearby Zenaga Massif (Leblanc 1975) and have been considered to be Eburnean basement of 2 Ga or older age by all previous workers. Northern parts of the Bou-Azzer inlier expose volcano-sedimentary sequences (e.g. Tichibinine Formation) considered to be part of an arc- or forearc-related sequence that is Neoproterozoic in age. Near the platform sequence crop out an intricately interleaved sequence of tectonic slices including ophiolitic fragments, metavolcanic rocks and metasediments (Leblanc 1975). This sequence

Fig. 1. Geological sketch map of the Anti-Atlas Proterozoic belts in southern Morocco and the location of the study area.

ME´LANGES AND OPHIOLITES, MOROCCO

is dominated by mafic–ultramafic plutonic bodies that had traditionally been viewed as parts of an ophiolite series (Leblanc 1981), but were later interpreted as a me´lange complex (Saquaque et al. 1989; Hefferan et al. 2002), with deformation taking place at blueschist-facies metamorphic conditions (Hefferan et al. 2002). The northern sector of the Bou-Azzer inlier consists of a metasedimentary sequence with subordinate volcanic units bearing calc-alkaline and island arc tholeiitic geochemical signatures (Naidoo et al. 1991). A pervasive greenschist-facies fabric, associated with recumbent tight to isoclinal folds, is variably developed within each of the tectonic blocks in the Bou-Azzer inlier, oriented NNE–SSW (Leblanc 1981). Kinematic analysis of these fabric elements in the southern part of the inlier is consistent with dominantly south-vergent sinistral oblique slip movements interpreted to record thrusting of the Bou-Azzer complex onto the WAC (Leblanc 1975; Saquaque et al. 1989). A clastic sedimentary succession, termed the Tiddiline Formation, unconformably overlies the tectonic blocks (Leblanc 1975). The majority of the above units are unconformably overlain by a thick succession of subhorizontal ignimbrites and conglomerates termed the Ouarzazate Supergroup.

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The earliest identifiable structure within the inlier occurs only in the gneissic basement. It consists of a NW–SE-striking, NE-dipping, upper greenschist-facies ductile fabric and a mineral stretching lineation that dips shallowly to the NW. A lower-grade, greenschist-facies fabric overprint is the dominant structure within the inlier. Much of the main schistosity is represented by a composite foliation and is associated with the development of folds and several generations of subparallel foliations trending WNW –ESE. In the south, structures associated with the main schistosity are dominated by WNW –ESE-striking composite foliations that dip steeply to the north with a common eastward-plunging mineral stretching lineation (Inglis et al. 2005). Greenschist-facies structures are repeatedly overprinted at successively lower temperatures by increasingly brittle fault zones, duplexes and cataclastic shear zones with a general WNW–ESE orientation (Leblanc 1981). This is consistent with faults and shear zones that crosscut both the Tiddiline sedimentary succession and late orogenic intrusions such as the Bleida granodiorite. A new precise U –Pb age of 579.4 + 1.2 Ma for the Bleida granodiorite (Inglis et al. 2004) provides a firm constraint on the latest stages of brittle transcurrent movement in the

Table 1. Representative analyses of Na-amphiboles from Ait-Atman and diabase bulk-rock composition used for the thermodynamic calculations Sample:

Ait021

Analysis no.: SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O Total

2 51.60 0.04 3.24 * 23.31 0.25 7.73 2.18 7.94 0.00 96.29

Si Ti Al Fe3þ† Fe2þ Mn Mg Ca Na K

7.739 0.005 0.573 0.931 1.993 0.032 1.728 0.350 2.309 0.000

Ait023c1 5 52.74 0.11 5.15 * 21.9 0.35 7.47 2.5 5.61 0.01 95.84 7.735 0.012 0.890 1.233 1.453 0.043 1.633 0.393 1.595 0.002

7 52.96 0 2.59 * 22.49 0 7.82 2.83 7.03 0.14 95.89 7.961 0.000 0.459 0.632 2.196 0.000 1.752 0.456 2.049 0.027

Bulk rock composition 9 51.9 0.04 2.11 * 22.11 0.03 8.56 2.98 8.03 0.01 95.85 7.873 0.005 0.377 0.535 2.271 0.004 1.936 0.484 2.362 0.002

*Not analysed. † Calculated. It should be noted that the Na-amphibole analyses are comparable with those of Hefferan et al. (2002).

53.41 0.32 14.68 3.00 6.37 0.20 7.24 7.86 4.02 0.19 97.27

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Bou-Azzer inlier. Block faulting and weak folding during the formation of the Atlas Mountains in Mesozoic times resulted in the uplift and exhumation of the Bou-Azzer basement inlier.

Subduction and thermal gradients throughout Earth’s history Reviews of the formation of the Archaean continents (e.g. Kusky & Polat 1999) show that subduction and collision processes during the Archaean at crustal scale are not well known, and thus hardly discussed. Although many examples of wedge structures are recognized in active mountain belts displaying an HP–LT (blueschist –eclogite conditions) evolution, such as the Central Alps or the Apennines in a Tethyan setting, and Japan (Shikoku island), the Franciscan Complex and the south of Alaska (Aleutians) in a Pacific setting, this mechanism may have also been active during Archaean times (see review by Sengo¨r 1999).

However, mechanisms of accretion within the wedge structure are supposed to have changed through time (Kusky & Polat 1999; Stern 2004). The thermal regime of orogenic belts through time has been the subject of many discussions (see Stern 2005). Considering the fact that the oldest ultrahigh-pressure (UHP) rocks that have been found until now occur only in the Pan-African belt in Mali (Caby 1994), dated at around 620 Ma (Jahn et al. 2001), Maruyama & Liou (1998) proposed that no UHP rocks occur in the Archaean and Proterozoic belts. They assumed that thermal regimes of subduction and collisional processes changed through time (Fig. 2). Other workers (Mo¨ller et al. 1995, 1998; O’Brien & Ro¨tzler 2003; Reddy et al. 2003), knowing that medium-T eclogites and high-pressure granulites are known from both old and young metamorphic terranes (e.g. c. 45 Ma, Namche Barwa, Eastern Himalayas; 400–340 Ma, European Variscides; 620 Ma, African belt, Mali; 1.9 Ga, Snowbird, Saskatchewan (Baldwin et al. 2003,

Fig. 2. Metamorphic facies diagram (after Oberha¨nsli et al. 2004) and subduction gradients occurring at different times. Whereas eclogite and UHP conditions were probably possible at all times during Earth history, blueschist and LT eclogite conditions are documented only since Palaeozoic times.


2004); 2.0 Ga in Tanzania) and that many new occurrences of UHP rocks are found as relicts in HP granulite terranes, argued that thermal and tectonic processes in the lithosphere have not changed significantly since at least the end of the Archaean, and that HP conditions could have existed during Archaean times. Despite apparent contradictions between these two views, discussion about secular changes in the P–T regimes of subduction processes will not constrain geothermal gradients at any time, because the possible temperature range of UHP minerals is very large (see Chopin 2003). In fact, UHP eclogite conditions may have been present even when the Earth was much hotter (Fig. 2). The large temperature range for UHP minerals is supported by recent experiments showing that partial melting of hydrous basalt under eclogitefacies conditions produces granitoid liquids with major- and trace-element compositions equivalent to those of Archaean tonalite–trondhjemite– granodiorite (TTG) (Rapp et al. 2003). Geotherms along subduction zones (referred to below as subduction gradients) in the lithosphere are hence better constrained by occurrences of blueschists in earlier times. Blueschists require unusually cold

X 0.0 0.0

Fe3+

0.2

237

upper mantle geotherms, found only in recent subduction zones (van Keken et al. 2002).

Metamorphic conditions of the Bou-Azzer ophiolites Petrology Hefferan et al. (2002) described and mapped occurrences of sodic amphiboles in the ophiolite suite of Bou-Azzer. These sodic amphiboles occur in a mineral assemblage together with garnet (of grossular-rich composition), epidote, albite, chlorite and quartz. Sodic amphiboles occur also in mafic rocks (diabase and gabbro) always at the contact with the Tiddiline Formation, which is mainly composed of greywackes and sandstones (see Hefferan et al. 2002, fig. 3). As index minerals of blueschist facies for mafic rocks, Na-amphiboles have been very important for the characterization of metamorphic conditions. Whereas the chemical compositions of Na-amphiboles characterizing blueschist facies are generally Al3þ-rich (XAl 0.6) with varying

= Fe3+/(Al3+ + Fe3+) 0.4

0.6

0.8

1.0 1.0

Ferroglaucophane

Riebeckite

Crossite

Fe2+

Greenschist environments

0.4

Blueschist terranes

0.4

X

Mg2+

0.6

0.6

= Fe2+/(Mg2+ + Fe2+)

= Mg2+/(Mg2+ + Fe2+)

0.8

X

0.2

Bou Azzer ophiolites

0.8

Glaucophane

0.2

Mg-Riebeckite

1.0 1.0

0.0 0.8

0.6

X

Al3+

0.4

0.2

0.0

= Al3+/(Al3+ + Fe3+)

Fig. 3. Mineral chemistry of sodic amphiboles in a Mg/(Fe2þ þ Mg) v. Al/(Fe3þ þ Al) diagram regardless of the parageneses and the chemical compositions of the rocks. We note a clear difference between the chemistry of Na-amphiboles from blueschist terranes and those growing under greenschist-facies conditions. The chemistry of Na-amphiboles from the Bou-Azzer ophiolite is compatible with greenschist metamorphic conditions.

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Mg2þ content, ranging from glaucophane to ferroglaucophane composition (Fig. 3; e.g. Oberha¨nsli 1986; Evans 1990; Bousquet et al. 1998; Katzir et al. 2000), the compositions of Na-amphiboles characterizing the greenschist facies are Al3þ-poor (XAl 0.3) with varying Mg2þ content, ranging from crossite to riebeckite and Mg-riebeckite compositions (Fig. 3; Oberha¨nsli 1986; Evans 1990; Frey et al. 1991). Mineral chemistry of Na-amphiboles from Bou-Azzer ranges between crossite and Mg-riebeckite composition (Fig. 3), the typical composition for greenschist-facies conditions.

P– T estimates Although first-order analyses of chemical composition of Na-amphiboles indicate greenschist-facies conditions, we will quantify P–T conditions of diabase samples from the Bou-Azzer ophiolite by thermodynamic methods. Hefferan et al. (2002) suggested greenschist to lower blueschist facies (c. 7 kbar, c. 350 8C) based on the contents of Na v. AlIV (Brown 1977). New developments in thermodynamic studies, however, show that mineral compositions are controlled not only by pressure and temperature conditions but also by whole-rock chemistry (de Capitani & Brown 1987; Powell et al. 1998; Connolly & Petrini 2002; Karpov et al. 2002). Methods. The method used to determine P –T conditions is based on the notion of ‘bulk-rock equilibrium’, which computes stable assemblages, including mode and composition of solution phases, for specific chemical rock compositions using the program DOMINO (de Capitani 1994). The independent variables are any combination of temperature, pressure, activity of a particular phase or compositional vectors. To include highly non-ideal solution models for minerals with potential miscibility gaps, stable mineral assemblages are computed using a Gibbs’ free energy minimization (de Capitani & Brown 1987). In such equilibrium phase diagrams, all phases are considered for each point assuming complete thermodynamic equilibrium for the whole rock. In this case, each field represents the predicted stability field of a particular assemblage. However, the interpretation of the diagrams is limited by the degree of equilibrium reached in rocks at each step of the metamorphic evolution. In none of the studied samples has any relic of the prograde, ‘burial’ evolution been observed; thus we assume that the rocks were fully equilibrated at the pressure peak of their history. In this case, we can use the ‘bulk-rock equilibrium’ method only to constrain the peak pressure. The updated JAN92.RGB thermodynamic database of Berman (1988) was used for all calculations,

supplemented with the following thermodynamic data: the Mg-chloritoid data of B. Patrick (listed by Goffe´ & Bousquet 1997), the Fe-chloritoid data of Vidal et al. (1994), the chlorite data of Hunziker (2003), and the alumino-celadonite data of Massonne & Szpurka (1997). Thermodynamic data for riebeckite were not available in the database used, but we use Cp- and Volume-functions of Holland & Powell (1998) for Na-amphiboles and experimental data from Holland (1988) for glaucophane, from Hoffmann (1972) for ferroglaucophane, and from Ernst (1962) for riebeckite. The solution models for phengite from Parra et al. (2002) and for chlorite from Hunziker (2003) have been used. Results. The equilibrium phase diagram for BouAzzer ophiolites shows that the stability field of the assemblage Na-amphibole, garnet, epidote, albite, chlorite and quartz in a diabase composition is well constrained in pressure between 5 and 9 kbar for temperatures varying between 300 and 600 8C (Fig. 4a). In this stability field, the composition of different minerals, specificially Na-amphibole, varies according to P and T. Whereas riebeckite and ferro-glaucophane components show opposite chemical evolution trends mainly controlled by temperature, with increase in riebeckite (Fig. 4b) and decrease in ferro-glaucophane (Fig. 4c) towards higher temperature conditions, the glaucophane component is controlled by pressure as well as temperature (Fig. 4d). Based on the isopleths of different Naamphibole end-members within the stability field of the Na-amphibole, garnet, epidote, albite, chlorite and quartz association, we can better constrain metamorphic conditions experienced by the diabases of Bou-Azzer. For the mineral assemblage described the composition of Na-amphiboles varies from pure glaucophane compositions at lower temperatures (below 400 8C) to riebeckite composition at higher temperatures conditions between 500 and 600 8C (Fig. 5). At intermediate temperature conditions, the composition is pressure dependent (Fig. 5). The sodic amphibole compositions found in Bou-Azzer rock samples are stable between 4 and 6 kbar for temperatures between 420 and 600 8C. As mentioned above, Hefferan et al. (2002) suggest greenschist to lower blueschist facies (c. 7 kbar, c. 350 8C) based on the content of Na v. content of AlIV (Brown 1977). However, according to many facies definitions (Yardley 1989; Spear 1993; Oberha¨nsli et al. 2004) pressures between 4 and 6 kbar for temperatures between 420 and 600 8C are typical for low-temperature to hightemperature (epidote– ) greenschist conditions.

Th e r i

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6

4

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0.1

0.3 0.4 0.5 0.6 0.7

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Fe-Glaucophane end-member 2 300

400 500 Temperature [°C]

(d) 10 No Garnet

Alb-Gt-Chl Ep-Gln

Th e r i

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i m no

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i m no

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(c) 10

a k Do

Alb-Gt-Chl Ep-Gln

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(b) 10

i m no

8

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i m no

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239

Th e r i

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2 300


600

Fig. 4. Stability field of mineral assemblage composed of Na-Amph þ garnet þ epidote þ albite þ chlorite þ quartz for diabase from the Bou-Azzer ophiolitic suite (a). Computed isopleths of riebeckite (b), ferro-glaucophane (c) and glaucophane (d) in the stability field of the mineral assemblage Na-Amph þ garnet þ epidote þ albite þ chlorite þ quartz. All diagrams are computed using the THERIAK-DOMINO software (http://titan.minpet.unibas.ch/ minpet/theriak/theruser.html). This is a program collection by de Capitani (1994) to calculate and plot thermodynamic functions, equilibrium assemblages and rock-specific equilibrium assemblage diagrams (elsewhere also called pseudo-sections). Based on its approach to equilibrium by means of Gibbs’ free energy minimization (see de Capitani & Brown 1987) rather than solving complex and large equation systems, the THERIAK-DOMINO software calculates and plots without user intervention that might be a source of errors.

Me´langes and ophiolites Tectonic me´langes are one of the hallmarks of convergent margins, yet understanding their genesis and relationships of specific structures to plate kinematic parameters has proven elusive because of the complex and seemingly chaotic nature of these units. Many field workers regard me´langes as too deformed to yield useful information, and simply map the distribution of me´lange type rocks

without further investigation. Other workers map clasts and matrix types, search for fossils or metamorphic index minerals in the me´lange, and assess the origin and original nature of the highly disturbed rocks. Analysis of deformational fabrics in tectonic me´lange may also yield information about the kinematics of past plate interactions (e.g. Le Pichon et al. 1988; Cowan & Brandon 1994; Kusky et al. 1997).

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Fig. 5. Computed Na-amphibole chemistry within the stability field of Na-Amph þ garnet þ epidote þ albite þ chlorite þ quartz for a basaltic composition. Each bold vertical line represents a composition range of Na-amphibole at each point in the stability field. The chemistry of Na-amphiboles occurring in the Bou-Azzer inlier documents HT greenschist metamorphic conditions at 4 and 6 kbar for temperatures between 420 and 600 8C.

Me´lange is a special type of breccia containing local or exotic competent blocks embedded in a less competent matrix in regions where high-level incompletely consolidated and unlithified sediments have been disturbed by imbricate faulting or gravitational gliding (Greenly 1891; Ramsay & Hubert 1987). A me´lange is defined on the basis of the following criteria: (1) a me´lange must be a mappable unit (typically at 1:25 000 scale); (2) it includes blocks of many sizes and diverse lithologies, some of which are ‘exotic’ (i.e. not derived from immediately adjacent units); (3) it has a matrix of fine-grained material, typically shale, slate, or serpentinite, with a tectonic fabric; (4) the matrix supports the blocks, which are not in contact with each other. Ophiolites are commonly associated with underlying continental margin units by me´lange. Me´langes can be subdivided into two general types (Fig. 6): (1) me´langes that are sandwiched between ophiolites and underlying continental margin units; (2) me´langes that form large

exposures and commonly include basic volcanic rocks, cherts and serpentinite, as blocks and broken formations. The latter can be classified as accretionary complexes. Intact ophiolites are not found regionally overlying such me´langes, although ophiolitic me´lange is widespread (Gansser 1974). The me´langes as a whole may be of tectonic, sedimentary or composite origin in different examples. Those of mainly sedimentary origin commonly correspond to the olistostromes (with olistoliths) of the classical literature.

Me´langes beneath ophiolites Me´langes are commonly found between overriding ophiolites and underlying continental margin units (Fig. 6a). Me´langes and broken formations were traditionally seen as deformed thrust sheets of ‘volcanic–sedimentary’ successions in which the present complexity was the result of a pervasive faulting of otherwise coherent units during emplacement (Oman, Eastern Mediterranean). In


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Fig. 6. Schematic models of the emplacement of ophiolites in subduction processes. (a) Obducted ophiolites are often associated with me´langes at their base ((1) Ho¨ck et al. 2002; (2) Collins & Robertson 1997; (3) Ferrie`re et al. 1988; (4) Searle & Malpas 1980). This type of me´lange is a consequence of the thrusting of the ophiolitic sequence over a platform. The matrix is composed either of serpentinite or of sediments. (b) Accretionary complexes display ultramafic and/or mafic rocks embedded in a sedimentary matrix, but are not overlain by an ophiolitic sequence. Often such mafic rocks were deeply buried at blueschist metamorphic conditions ((5) Cloos 1986; (6) Parkinson 1996; (7) Bousquet 2007; (8) Bousquet et al. 2002; (9) Schwartz et al. 2000; (10) Brandon & Calderwood 1990).

addition, me´langes play an important role in ophiolite emplacement in many areas such as Albania, the Himalayas (e.g. Corfield et al. 1999; Robertson 2000; Ho¨ck et al. 2002). Facies analysis, and geochemical and structural studies show that most of the accreted me´langes range in setting from the distal continental rise to open oceanic. In summary, me´langes beneath ophiolites are produced by obduction of large ophiolites onto continental margins. They record the obduction processes of oceanic crust and deep-sea sediments (Robertson 2004). Such me´langes are formed during the thrusting of the ophiolites onto the continental margin.

Accretionary complexes The second type of me´lange is very widespread but is not found directly beneath large overriding ophiolites (Fig. 6b). Such me´langes are typically complex and often affected by multiple deformation events such that any genetic distinction between sedimentary and tectonic origins is difficult and unreliable. The main impetus for the recognition of me´langes as recording subduction of oceanic crust came instead from studies of the Franciscan Complex. This me´lange includes ophiolitic material (e.g. serpentinite), but is not overlain by any ophiolite, as is the case in

many accretionary complexes (e.g. Alps, Bousquet et al. 2002; Bousquet 2008; Sulawesi, Parkinson 1996; Olympic Mountains, Brandon & Vance 1992). Blocks in the Franciscan Complex include blueschist metamorphic rocks with glaucophane and lawsonite, which indicate HP– LT metamorphism. All the blocks, like ultramafic rocks, are embedded in a fine-grained matrix composed mainly of black shale. No consensus on the tectonic or sedimentary origin of the Franciscan Complex exists yet; some researchers have inferred a mainly tectonic origin (e.g. Cloos 1984), whereas others have envisaged a mainly sedimentary origin (e.g. Cowan 1978). Accretionary complexes similar to the Franciscan Complex occur widely along the Tethys belt from the Alps (Bousquet et al. 2002) to New Caledonia (Potel et al. 2006) including the Eastern Mediterranean region (Jolivet et al. 1998; Robertson 2004), Turkey (Okay et al. 1996), Iran (Gansser 1974) and Sulawesi (Parkinson 1996). All examples, displaying HP– LT metamorphism, are regionally associated with subduction, but are not directly overlain by ophiolites. The various me´langes, rather than the ophiolites themselves, are important indicators of the former existence of oceanic areas. The main reason is that most of the ophiolites preserve unusual tectonic settings (e.g. suprasubduction-zone genesis),

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Fig. 7. Geological map of the western part of the Bou-Azzer inlier (modified after Leblanc 1975). Cross-sections 1 and 2 are shown in figure 8.

whereas most ‘normal’ mid-ocean ridge basalt (MORB) was subducted. When such subduction takes place only fragments of oceanic crust and sediments are preserved, including serpentinite, volcanic seamounts and related pelagic sediments.

The Bou-Azzer ophiolite: me´lange beneath ophiolites or accretionary complex? Saquaque et al. (1989) and Hefferan et al. (2002) described the Bou-Azzer inlier as resulting from an accretionary complex formed during the

Fig. 8. Geological cross-sections across the Bou-Azzer inlier (locations are shown in fig. 7; modified after Leblanc 1975; Saquaque et al. 1989). Structural relations between the units clearly show that the me´lange sequence is located structurally above the basement and below the ophiolitic sequence.


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Fig. 9. Model for the emplacement of the Bou-Azzer ophiolitic suite. 750 –700 Ma, north-dipping subduction of the Pan-African ocean; 680 –660 Ma, formation of me´lange and HT greenschist facies metamorphic overprint during obduction and terrane assembly; 650–640 Ma, syncollisional magmatism crosscutting earlier tectonic structures. A northern continent is assumed based on the absence of Pan-African age oceanic crust north of Bou-Azzer.

Pan-African subduction and not representing a coherent sample of Precambrian oceanic lithosphere obducted onto the Eburnean margin (Leblanc 1976). In this model the whole Bou-Azzer inlier is considered to be an accretionary complex formed by a huge me´lange juxtaposed to the

ophiolite (Fig. 7). However, coherent petrographic and stratigraphic relationships can be traced along strike, in some cases for kilometres (Leblanc & Billaud 1978; Church 1991; Leblanc & MoussinePouchkine 1994). The major part of the inlier is composed of different ophiolite sequences

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including ultramafic rocks that are juxtaposed along faults and shear zones (Fig. 8), indicating early top-to-the-south movements (Leblanc 1975). The me´lange is limited to the southern part of the inlier at the contact with the basement (Fig. 8). The matrix of the me´lange varies from serpentinite to strongly deformed volcanosedimentary rocks. Blocks mainly consist of ophiolitic fragments, metagreywackes and quartzites. Occurrences of metabasaltic breccias within the me´lange testify to the tectonic origin of the blocks. The me´lange forms slices interspersed between the ophiolite sequence in the north and a basement in the south. This geometry, already described by Saquaque et al. (1989), suggests that the me´lange forms the base of the ophiolitic rocks. Thus we interpret the whole sequence as an ophiolite obducted southwards onto the WAC platform with its underlying me´lange (Fig. 9). After the obduction of the ophiolite, we note accretion of terranes, such as island arcs, to the ophiolitic complex. During terrane assembly, the ophiolites were dismembered and thickened, allowing HT greenschist metamorphic conditions (Fig. 9). During collision between the northern continent and the WAC, diorite to quartz diorite syntectonic plutons and diabase dykes intruded the different accreted terranes. Radiometric dates on various granodiorites of Bou-Azzer have yielded ages between 650 and 640 Ma (Inglis et al. 2004, 2005). All units were subsequently deformed and metamorphosed under lower greenschist-facies conditions, indicated by the growth of chlorite within shear zones in granodiorite. The entire inlier appears to be a complex ensemble of diverse igneous, metamorphic and sedimentary rock units that have been accreted, juxtaposed and deformed at different times during the closure of the Pan-African ocean (Fig. 9) and not assembled at the same time in an accretionary complex. Several studies (Saquaque et al. 1989; Hefferan et al. 2000, 2002; Ennih & Lie´geois 2001) clearly show that the Bou-Azzer ophiolitic suite is the remnants of a fore-arc assemblage that evolved above a north-dipping subduction zone. Despite this, the dip orientation of the subduction in the Moroccan Anti-Atlas during the Pan-African orogen is still controversial (Gasquet et al. 2005). However, the present geometry of the ophiolite suite (Soulaimani et al. 2006), the accretion sequence with a me´lange at its base at the contact with the WAC basement in the south, and the early top-to-the-south sense of shear combined with the calc-alkaline volcanism in the north (Saghro massif, Saquaque et al. 1992) clearly allow us to argue for a north-dipping subduction.

The present study shows that the ophiolite suite of Bou-Azzer is not part of an accretionary complex, but is an obducted ophiolite with a me´lange at its base. The ophiolite sequence was dismembered and delaminated during its accretion onto the WAC margin. No evidence for HP–LT (blueschist-facies) metamorphic conditions can be found in these rocks. Metamorphic conditions experienced by the ophiolites are limited to HT greenschist-facies conditions (5–6 kbar, 500–550 8C), which are typical conditions for collision or obduction settings. We therefore consider the ophiolites of Bou-Azzer not to be similar to the HP ophiolites of the Franciscan Complex or medium-pressure ophiolites of Sanbagawa, but rather similar to the Albanian or Cyprus ophiolites, which are dismembered ophiolite sequences overprinted by greenschist-facies conditions. R.B. thanks R. Caby for fruitful discussions and for his encouragement to write this paper. Funding by the French– Moroccan project MA No. 13 and the University of Basel are greatly appreciated. This paper benefited from constructive comments of K. Hefferan, J. C. Schumacher and D. Marquer. N. Ennih is thanked for his helpful editorial work.

References A DMOU , H. & J UTEAU , T. 1998. Dećouverte d’un syste`me hydrothermal oceánique fossile dans l’ophiolite antećambrienne de Khzama (massif du Siroua, Anti-Atlas marocain). Comptes Rendus de l’Academie des Sciences, Se´rie IIA, 327, 335– 340. B ALDWIN , J. A., B OWRING , S. A. & W ILLIAMS , M. L. 2003. Petrological and geochronological constraints on high pressure, high temperature metamorphism in the Snowbird tectonic zone, Canada. Journal of Metamorphic Geology, 21, 81– 98. B ALDWIN , J. A., B OWRING , S. A., W ILLIAMS , M. L. & W ILLIAMS , I. S. 2004. Eclogites of the Snowbird tectonic zone: petrological and U –Pb geochronological evidence for Paleoproterozoic high-pressure metamorphism in the western Canadian Shield. Contributions to Mineralogy and Petrology, 147, 528– 548. B ERMAN , R. G. 1988. Internally-consistent thermodynamic data for minerals in the system Na2O–K2O–Ca–MgO– FeO–Fe2O3 –Al2O3 –SiO2 –TiO2 –H2O–CO2. Journal of Petrology, 29, 445–522. B OUOUGRI , E. 2003. The Moroccan Anti-Atlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton—discussion. Precambrian Research, 120, 179–183. B OUSQUET , R. 2008. Metamorphic heterogeneities within a same HP unit: overprint effect or metamorphic mix? Lithos, DOI: 10.1016/j.lithos.2007.09.010. B OUSQUET , R., O BERHA¨ NSLI , R., G OFFE´ , B., J OLIVET , L. & V IDAL , O. 1998. High pressure– low temperature metamorphism and deformation in the Bu¨ndnerschiefer of the Engadine window: implications for the regional evolution of the eastern Central Alps. Journal of Metamorphic Geology, 16, 657– 674.

ME´LANGES AND OPHIOLITES, MOROCCO B OUSQUET , R., G OFFE´ , B., V IDAL , O., O BERHA¨ NSLI , R. & P ATRIAT , M. 2002. The tectono-metamorphic history of the Valaisan domain from the Western to the Central Alps: new constraints for the evolution of the Alps. Geological Society of America Bulletin, 114, 207– 225. B RANDON , M. T. & C ALDERWOOD , A. R. 1990. Highpressure metamorphism and uplift of the Olympic subduction complex. Geology, 18, 1252–1255. B RANDON , M. T. & V ANCE , J. A. 1992. Tectonic evolution of the Cenozoic Olympic subduction complex, Washington State, as deduced from fission track ages for detrical zircons. American Journal of Science, 292, 565– 636. B ROWN , E. H. 1977. The crossite content of Ca-amphibole as a guide to pressure of metamorphism. Journal of Petrology, 18, 53–72. C ABY , R. 1994. Precambrian coesite from nothern Mali: first record and implications for plate tectonics in the trans-Saharan segment of the Pan-African belt. European Journal of Mineralogy, 6, 235– 244. C HOPIN , C. 2003. Ultrahigh-pressure metamorphism: tracing continental crust into the mantle. Earth and Planetary Science Letters, 212, 1 –14. C HURCH , W. R. 1991. Comment on ‘Precambrian accretionary tectonics in the Bou Azzer –El Graara region, Anti-Atlas, Morocco’. Geology, 19, 285– 287. C LOOS , M. 1984. Flow me´langes and the structural evolution of accretionary wedges. In: RAYMOND , L. A. (ed.) Me´langes: Their Nature, Origin and Significance. Geological Society of America, Special Papers, 198, 71–79. C LOOS , M. 1986. Blueschists in the Franciscan Complex of California: Petrotectonic constraints on uplift mechanisms. In: E VANS , B. W. & B ROWN , E. H. (eds) Blueschists and Eclogites. Geological Society of America, Memoirs, 164, 77–94. C OLLINS , A. S. & R OBERTSON , A. H. F. 1997. Lycian me´lange, southwestern Turkey: An emplaced Late Cretaceous accretionary complex. Geology, 25, 255–258. C ONNOLLY , J. A. D. & P ETRINI , K. 2002. An automated strategy for calculation of phase diagram sections and retrieval of rock properties as a function of physical conditions. Journal of Metamorphic Geology, 20, 697–708. C ORFIELD , R. I., S EARLE , M. P. & G REEN , O. R. 1999. Photang thrust sheet: an accretionary complex structurally below the Spontang ophiolite constraining timing and tectonic environment of ophiolite obduction, Ladakh Himalaya, NW India. Journal of the Geological Society, London, 156, 1031–1044. C OWAN , D. S. 1978. Origin of blueschist-bearing chaotic rocks in Franciscan Complex, San Simeon, California. Geological Society of America Bulletin, 89, 1415–1423. C OWAN , D. S. & B RANDON , M. T. 1994. A symmetrybased method for kinematic analysis of large-slip brittle fault zones. American Journal of Science, 294, 257–306. DE C APITANI , C. & B ROWN , T. H. 1987. The computation of chemical equilibrium in complex systems containing non-ideal solutions. Geochimica et Cosmochimica Acta, 51, 2639– 2652.

245

C APITANI , C. 1994. Gleichgewichts-Phasendiagramme: Theorie und Software. Berichte der Deutschen Mineralogischen Gesellschaft, 6, 48. D’L EMOS , R. S., I NGLIS , J. D. & S AMSON , S. D. 2006. A newly discovered orogenic event in Morocco: Neoproterozic ages for supposed Eburnean basement of the Bou Azzer inlier, Anti-Atlas Mountains. Precambrian Research, 147, 65– 78. E NNIH , N. & L IE´ GEOIS , J.-P. 2001. The Moroccan AntiAtlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton. Precambrian Research, 112, 289–302. E RNST , W. G. 1962. Synthesis, stability relations, and occurrence of riebeckite and riebeckite – arfvedsonite solid-solutions. Journal of Geology, 70, 689–736. E VANS , B. W. 1990. Phase relation of epidote blueschists. Lithos, 25, 3 –23. F ERRIE` RE , J., B ERTRAND , J., S IMANTOV , J. & D E W EVER , P. 1988. Comparaison entre les formations volcano-de´tritiques (‘Me´langes’) du Malm des Helle`nides internes (Othrys, Eubeé): implications geódynamiques. Bulletin of the Geological Society of Greece, 20, 223– 235. F REY , M., D E C APITANI , C. & L IOU , J. G. 1991. A new petrogenetic grid for low-grade metabasites. Journal of Metamorphic Geology, 9, 497–509. G ANSSER , A. 1974. The ophiolitic me´lange, a world-wide problem on Tethyan examples. Eclogae Geologicae Helvetiae, 67, 479–507. G ASQUET , D., L EVRESSE , G., C HEILLETZ , A., A ZIZI S AMIR , M. R. & M OUTTAQI , A. 2005. Contribution to a geodynamic reconstruction of the Anti-Atlas (Morocco) during Pan-African times with the emphasis on inversion tectonics and metallogenic activity at the Precambrian –Cambrian transition. Precambrian Research, 140, 157–182. G OFFE´ , B. & B OUSQUET , R. 1997. Ferrocarpholite, chloritoı¨de et lawsonite dans les me´tapelites des unite´s du Versoyen et du Petit St Bernard (zone valaisanne, Alpes occidentales). Schweizerische Mineralogische und Petrographische Mitteilungen, 77, 137– 147. G REENLY , E. 1891. The Geology of Anglesey. Memoir, Geological Survey of Great Britain. H EFFERAN , K. P., A DMOU , H., K ARSON , J. A. & S AQUAQUE , A. 2000. Anti-Atlas (Morocco) role in Neoproterozoic Western Gondwana reconstruction. Precambrian Research, 103, 89–96. H EFFERAN , K. P., A DMOU , H., H ILAL , R. ET AL . 2002. Proterozoic blueschist-bearing me´lange in the AntiAtlas Mountains, Morocco. Precambrian Research, 118, 179–194. H O¨ CK , V., K OLLER , F., M EISEL , T., O NUZI , K. & K NERINGER , E. 2002. The Jurassic South Albanian ophiolites: MOR- vs SSZ-type ophiolites. Lithos, 65, 143– 164. H OFFMANN , C. 1972. Natural and synthetic ferroglaucophane. Contributions to Mineralogy and Petrology, 34, 135– 149. H OLLAND , T. J. B. 1988. Preliminary phase-relations involving glaucophane and applications to highpressure petrology– New heat capacity and thermodynamic data. Contributions to Mineralogy and Petrology, 99, 134–142. DE

246

R. BOUSQUET ET AL.

H OLLAND , T. J. B. & P OWELL , R. 1998. An internally consistent thermodynamic data set for phases of petrological interest. Journal of Metamorphic Geology, 16, 309– 343. H UNZIKER , P. 2003. The stability of tri-octahedral Fe2þ – Mg–Al chlorite. A combined experimental and theoretical study. PhD Thesis, Universita¨t Basel. I NGLIS , J. D., M AC L EAN , J. S., S AMSON , S. D., D’L EMOS , R. S., A DMOU , H. & H EFFERAN , K. P., 2004. A precise U–Pb zircon age for the Bleida granodiorite, Anti-Atlas, Morocco: implications for the timing of deformation and terrane assembly in the eastern Anti-Atlas. Journal of African Earth Sciences, 39, 277 –283. I NGLIS , J. D., D’L EMOS , R. S., S AMSON , S. D. & A DMOU , H. 2005. Geochronological constraints on Late Precambrian intrusion, metamorphism, and tectonism in the Anti-Atlas Mountains. Journal of Geology, 113, 439– 450. J AHN , B.-M., C ABY , R. & M ONIE´ , P. 2001. The oldest UHP eclogites of the world: age of UHP metamorphism, nature of protoliths and tectonic implications. Chemical Geology, 178, 143– 158. J OLIVET , L., G OFFE´ , B., B OUSQUET , R., O BERHA¨ NSLI , R. & M ICHARD , A. 1998. Detachments in high-pressure mountain belts, Tethyan examples. Earth and Planetary Science Letters, 160, 31–47. K ARPOV , I. K., C HUDNENKO , K. V., K ULIK , D. A. & B YCHINSKII , V. A. 2002. The convex programming minimization of five thermodynamic potentials other than Gibbs’ energy in geochemical modeling. American Journal of Science, 312, 281– 311. K ATZIR , Y., A VIGAD , D., M ATTHEWS , A., G ARFUNKEL , Z. & E VANS , B. 2000. Origin, HP/LT metamorphism and cooling of ophiolitic me´langes in southern Evia (NW Cyclades), Greece. Journal of Metamorphic Geology, 18, 699–718. K USKY , T. M. & P OLAT , A. 1999. Growth of granite– greenstone terranes at convergent margins, and stabilization of Archean cratons. Tectonophysics, 305, 43–73. K USKY , T. M., B RADLEY , D. C., H AEUSSLER , P. J. & K ARL , S. 1997. Controls on accretion of flysch and me´lange belts at convergent margins: Evidence from the Chugach Bay thrust and Iceworm me´lange, Chugach accretionary wedge, Alaska. Tectonics, 16, 855–878. L EBLANC , M. 1975. Ophiolites prećambriennes et gıˆtes arseńie´s de Cobalt (Bou Azzer—Maroc). Doctorat d’E´tat, Universite´ Paris VI. L EBLANC , M. 1976. Proterozoic oceanic crust at Bou Azzer. Nature, 261, 34–35. L EBLANC , M. 1981. The Late Proterozoic ophiolites of Bou Azzer (Morocco): Evidence for Pan-African plate tectonics. In: K RO¨ NER , A. (ed.) Precambrian Plate Tectonics. Elsevier, Amsterdam, 435–451. L EBLANC , M. & B ILLAUD , P. 1978. Volcanosedimentary copper deposit on a continental margin of Upper Proterozoic Age—Bleida (Anti-Atlas, Morocco). Economic Geology, 73, 1101– 1111. L EBLANC , M. & L ANCELOT , J. R. 1980. Geodynamic interpretation of Pan-African region (Late Precambrian) in Anti-Atlas (Morocco) from geological and geochronological data. Canadian Journal of Earth Sciences, 17, 142 –155.

L EBLANC , M. & M OUSSINE -P OUCHKINE , A. 1994. Sedimentary and volcanic evolution of a Neoproterozoic continental margin (Bleida, Anti-Atlas, Morocco). Precambrian Research, 70, 25–44. L E P ICHON , X., B ERGERAT , F. & R OULET , M.-J. 1988. Plate kinematics and tectonics leading to the Alpine belt formation: a new analysis. In: C LARK , S. P., B URCHFIEL , B. C. & S UPPE , J. (eds) Processes in Continental Lithospheric Deformation. Geological Society of America, Special Papers, 218, 111–131. M ARUYAMA , S. & L IOU , J. G. 1998. Initiation of ultrahigh-pressure metamorphism and its significance on the Proterozoic– Phanerozoic boundary. Island Arc, 7, 6– 35. M ASSONNE , H.-J. & S ZPURKA , Z. 1997. Thermodynamic properties of white micas on the basis of high-pressure experiments in the systems K2O–MgO– Al2O3 – and K2O– FeO– Al2O3 – SiO2 – H2O. SiO2 – H2O Lithos, 41, 229– 250. M O¨ LLER , A., A PPEL , P., M EZGER , K. & S CHENK , V. 1995. Evidence for 2 Ga subduction zone—Eclogites in the Usagaran belt of Tanzania. Geology, 23, 1067– 1070. M O¨ LLER , A., M EZGER , K. & S CHENK , V. 1998. Crustal age domains and the evolution of the continental crust in the Mozambique Belt of Tanzania: Combined Sm– Nd, Rb–Sr, and Pb– Pb isotopic evidence. Journal of Petrology, 39, 749–783. N AIDOO , D. D., B LOOMER , S. H., S AQUAQUE , A. & H EFFERAN , K. P. 1991. Geochemistry and significance of metavolcanic rocks from the Bou Azzer– El Graara Ophiolite (Morocco). Precambrian Research, 53, 79–97. O BERHA¨ NSLI , R. 1986. Blue amphibole in metamorphosed Mesozoic mafic rocks from Central Alps. In: E VANS , B. W. & B ROWN , E. H. (eds) Blueschists and Eclogites. Geological Society of America, Memoirs, 164, 239– 247. O BERHA¨ NSLI , R., B OUSQUET , R., E NGI , M. ET AL . 2004. Metamorphic structure of the Alps. In: Commission for the Geological Map of the World (ed.) Explanatory Note to the Map ‘Metamorphic Structure of the Alps’. UNESCO, Paris. O’B RIEN , P. J. & R O¨ TZLER , J. 2003. High-pressure granulites: formation, recovery of peak conditions and implications for tectonics. Journal of Metamorphic Geology, 21, 3 –20. O KAY , A. I., S ATIR , M., M ALUSKI , H., S IYAKO , M., M ONIE´ , P., M ETZGER , R. & A KYU¨ Z , S. 1996. Paleo- and Neo-Tethyan events in northwest Turkey: geological and geochronological constraints. In: Y IN , A. & H ARRISON , M. (eds) Tectonics of Asia. Cambridge University Press, Cambridge, 420–441. P ARKINSON , C. D. 1996. The origin and significance of metamorphosed tectonic blocks in melanges: Evidence from Sulawesi, Indonesia. Terra Nova, 8, 312–323. P ARRA , T., V IDAL , O. & J OLIVET , L. 2002. Relation between the intensity of deformation and retrogression in blueschist metapelites of Tinos Island (Greece) evidenced by chlorite– mica local equilibria. Lithos, 63, 41–66. P OTEL , S., F ERREIRO -M A¨ HLMANN , R., S TERN , W. B., M ULLIS , J. & F REY , M. 2006. Very low-grade

ME´LANGES AND OPHIOLITES, MOROCCO metamorphic evolution of pelitic rocks under highpressure/low-temperature conditions, NW New Caledonia (SW Pacific). Journal of Petrology, 47, 991–1015. P OWELL , R., H OLLAND , T. & W ORLEY , B. 1998. Calculating phase diagrams involving solid solutions via non-linear equations, with examples using THERMOCALC. Journal of Metamorphic Geology, 16, 577–588. R AMSAY , J. G. & H UBERT , M. 1987. The Techniques of Modern Structural Geology. Academic Press, London. R APP , R. P., S HIMIZU , N. & N ORMAN , M. D. 2003. Growth of early continental crust by partial melting of eclogite. Nature, 425, 605– 609. R EDDY , S. M., C OLLINS , A. S. & M RUMA , A. 2003. Complex high-strain deformation in the Usagaran Orogen, Tanzania: structural setting of Palaeoproterozoic eclogites. Tectonophysics, 375, 101–123. R OBERTSON , A. H. F. 2000. Formation of me´langes in the Indus suture Zone, Ladakh Himalaya by successive subduction–accretion, collisional and post-collisional processes during Late Mesozoic–Late Tertiary time. In: K HAN , M. A., T RELOAR , P. J., S EARLE , M. P. & J AN , M. Q. (eds) Tectonics of the Naga Parbat Syntaxis and the Western Himalaya. Geological Society, London, Special Publications, 170, 333– 374. R OBERTSON , A. 2004. Development of concepts concerning the genesis and emplacement of Tethyan ophiolites in the Eastern Mediterranean and Oman regions. Earth-Science Reviews, 66, 331–387. S AQUAQUE , A., A DMOU , H., K ARSON , J. A. & H EFFERAN , K. P. 1989. Precambrian accretionary tectonics in the Bou Azzer– El Graara region, Anti-Atlas, Morocco. Geology, 17, 1107– 1110. S AQUAQUE , A., B ENHARREE , M., A BIA , H., M RINI , Z., R EUBER , I. & K ARSON , J. A. 1992. Evidence for a Panafrican volcanic arc and wrench fault tectonics in the Jbel Saghro, Anti-Atlas, Morocco. Geologische Rundschau, 81, 1– 13. S CHWARTZ , S., L ARDEAUX , J.-M., G UILLOT , S. & T RICART , P. 2000. Diversite´ du me´tamorphisme

247

ećlogitique dans le massif ophiolitique du Monviso (Alpes occidentales, Italie). Geodinamica Acta, 13, 169– 188. S EARLE , M. P. & M ALPAS , J. 1980. Structure and metamorphism of rocks beneath the Semail ophiolite of Oman and their significance in ophiolite obduction. Transactions of the Royal Society of Edinburgh: Earth Sciences, 71, 247 –262. S ENGO¨ R , A. M. C. 1999. Continental interiors and cratons: any relation? Tectonophysics, 305, 1– 42. S OULAIMANI , A., J AFFAL , M., M AACHA , L., K CHIKACH , A., N AJINE , A. & S AIDI , A. 2006. Magnetic modelling of the Bon Azzer –El Graara ophiolite (central Anti-Atlas, Morocco). Geodynamic implications of the Panafrican reconstruction. Comptes Rendus Geosciences, 338, 153 –160. S PEAR , F. S. (ed.) 1993. Metamorphic Phase Equilibria and Pressure–Temperature– Time Paths. Mineralogical Society of America, Monograph. S TERN , R. J. 2004. Subduction initiation: spontaneous and induced. Earth and Planetary Science Letters, 226, 275– 292. S TERN , R. J. 2005. Evidence from ophiolites, blueschists, and ultrahigh-pressure metamorphic terranes that the modern episode of subduction tectonics began in Neoproterozoic time Geology, 33, 557–560. VAN DER V ELDEN , A. J. & C OOK , F. A. 1999. Proterozoic and Cenozoic subduction complexes: A comparison of geometric features. Tectonics, 18, 575 –581. VAN K EKEN , P. E., K IEFER , B. & P EACOCK , S. M. 2002. High-resolution models of subduction zones: Implications for mineral dehydration reactions and the transport of water into the deep mantle. Geochemistry, Geophysics, Geosystems, 3 (10), 1056, DOI: 10.1029/ 2001GC000256. V IDAL , O., T HEYE , T. & C HOPIN , C., 1994. Experimental study of chloritoid stability at high pressure and various fO2 conditions. Contributions to Mineralogy and Petrology, 118, 256– 270. Y ARDLEY , B. W. D. 1989. An Introduction to Metamorphic Petrology. Longman, Harlow.

Gold mineralization in the Proterozoic Bleida ophiolite, Anti-Atlas, Morocco ABDELHAY BELKABIR1, MICHEL JE´BRAK2, LHOU MAACHA3, M. RACHID AZIZI SAMIR3 & ATMANE MADI4 1

Universite´ Cadi Ayyad, Faculte´ des Sciences et Techniques de Marrakech, De´partement de Geólogie, B.P. 549, Marrakech 40 000, Morocco (e-mail: [email protected]) 2

Universite´ du Que´bec a` Montreál, De´partement des Sciences de la Terre et de l’Atmosphe`re, CP 8888 Centre Ville, Montreál, Que´bec H3C 3P8, Canada 3

Reminex Exploration, 235, Lot. Hamra, Marrakech, Morocco 4

Akka Gold Mining, Tafraout, Morocco

Abstract: The newly discovered (1998) West Bleida gold mineralization (3 tonnes metal Au) lies west of the main Moroccan Bleida copper deposit (1981– 1991) in the central Anti-Atlas (southern Morocco). It is hosted by metamorphosed and deformed mafic to intermediate volcanic rocks that are part of the Neoproterozoic tholeiitic volcanosedimentary series forming the stratigraphically upper part of the Bou Azzer ophiolite sequence. Strong sericitization and local silicification are associated with mineralization. These altered rocks represent a proximal hydrothermal alteration halo around the West Bleida ore zones. Normative chlorite characterizes the metamorphic assemblage away from the ore zones. Gold mineralization primarily occurs as deformed gold-bearing quartz veins and disseminations in Cu-rich chert zones (chalcopyrite–malachite), Fe-rich lithofacies and breccia zones. Gold is accompanied by small amounts of copper sulphides (,1% modal chalcopyrite). Scanning electron microscope– energy dispersive spectrometry analyses of gold grains from veins and disseminations reveal the presence of palladium as inclusions of Pd–As –Sb, Pd–Bi –Se and Pd–Te mineral phases. An electron microprobe study confirms the presence of two types of gold. The first is an alloy of Au–Ag–Pd, typically bordered by small grains of Pd and Bi (Te,Sb) phases and associated with a metamorphic assemblage. Isomertieite, Pd11(Sb2,As2), was identified as one of the phases. The second type of gold is electrum (10% Ag, 90% Au), which is always associated with fractures and occurs with hematite and white mica. Based on its form and habits, West Bleida gold reflects two distinct generations of fluid activity. The primary event precipitated Au –Ag –Pd alloys from Au–Pd-bearing hydrothermal fluids and produced auriferous quartz veins and disseminations within mafic rocks of the Bleida ophiolitic accretionary complex. It was structurally and lithologically controlled. This early event is preserved in the deeper (and thus fresher) zones more than 80 m below the surface. Intense tectonic overprinting obscures the genetic relationship between vein and disseminated styles of mineralization, both of which contain Pd-rich gold, but some of the auriferous quartz veins are observed to crosscut disseminated mineralization. Two possible hypotheses are considered: the pre-tectonic root of a volcanogenic massive sulphide system, or a late tectonic orogenic (mesothermal) deposit. The presence of Pd minerals and anomalous cobalt concentrations suggest a source in ultramafic rocks. The second event, characterized by inclusion-free electrum, occurred much later and represents the alteration and weathering of the primary Pd-rich gold assemblage by oxidizing surface fluids. It affected all mineralized units and structures to a depth of 80 m. This post-tectonic surficial alteration also caused leaching of Cu-sulphides, which may explain their low abundances in the upper parts of the ore zones.

The Anti-Atlas of southern Morocco is a longstanding gold–copper province (Leblanc & Lancelot 1980). One of its more notable systems is the Bleida copper deposit, which was mined from 1981 to 1997 (Leblanc & Billaud 1978). In 1996, the MANAGEM Company discovered gold mineralization in the Neoproterozoic ophiolite complex of the West Bleida area using the results of stream and soil sediment geochemistry followed by

trenching and rock sampling. With a measured resource of 2.5 million tonnes at 2 g t21 Au, the West Bleida deposit constitutes the first major gold discovery in the central part of the Anti-Atlas. A study by Barakat et al. (2002) suggested that gold at West Bleida is confined to gold –quartz veins and may represent an epithermal deposit. The results of Reminex’s recent exploration work and initial definition drilling on the West


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Bleida project have provided significant new data that have implications for the origin of gold in this area. Gold at West Bleida occurs as auriferous quartz veins and disseminations in mafic rocks of the ophiolitic complex. This ultramafic geological environment is rather unusual for an epithermal deposit (Hedenquist et al. 2000) and classic epithermal models cannot adequately explain the platinum group element (PGE) content of some of the goldbearing facies. This paper uses the results of detailed geological and mineralogical studies to develop an alternative model in which the West Bleida deposit most likely represents the root of a volcanogenic massive sulphide system or a metamorphosed orogenic (mesothermal?) gold system in an ophiolitic accretionary complex that can be considered as a probable mesothermal gold deposit, similar to those in the Archaean belts of Canada (Card et al. 1989) and Australia (Groves & Foster 1991).

Geological setting Regional setting The Bleida district is located in the central AntiAtlas, along the southern edge of the Bou Azzer-El Graraa (BAG) inlier (Fig. 1). Leblanc (1981) concluded that the ophiolite of the inlier represents a mid-ocean ridge basalt (MORB)-type environment generated within an oceanic rift system, and subdivided the BAG inlier into five stratigraphic sequences from Precambrian to Cambrian (Fig. 1). The Precambrian basement is Palaeoproterozoic in age and comprises gneiss, amphibolite and leucogranite. Basement rocks were affected by the 2 Ga Eburnean orogeny (Rb –Sr: Charlot 1982; U –Pb: Thomas et al. 2004; U –Pb: Gasquet et al. 2005), and constitute the northern edge of the West African craton. The Neoproterozoic sequence

Fig. 1. Geological map of the Anti-Atlas showing the location of Bou Azzer– El Graara inlier (modified from Leblanc 1981).

GOLD MINERALIZATION IN BLEIDA OPHIOLITE

unconformably overlies the Palaeoproterozoic rocks. Two main units characterize the Neoproterozoic sequence. The lower unit (Cryogenian) includes the ophiolite complex and associated diorites in the axial zone of the BAG inlier (Rb–Sr: 788+8 Ma, Clauer 1976). It consists of a basal unit of ultramafic rocks overlain by (in order) cumulate gabbros, isotropic gabbros, rare sheeted dykes and basaltic rocks. These rocks host significant Co, Ni and As mineralization, as well as numerous chromite deposits. The uppermost unit is a mafic volcaniclastic series (Leblanc & Billaud 1978, 1990; Saquaque et al. 1989). The late Neoproterozoic sequence of the Ouarzazate Supergroup (Ediacaran) unconformably overlies the earlier Neoproterozoic sequences (Anti-Atlas Supergroup; Bleida Group) (see Thomas et al. 2004; Gasquet et al. 2005). The dominant lithological units are rhyolite (U –Pb: 580 + 15 Ma, Juery et al. 1974) and volcaniclastic rocks. The Proterozoic – Cambrian series (Adoudounian) overlies the Palaeoproterozoic and Neoproterozoic sequences. It consists of dolomite and grey sandstone units with minor intercalations of felsic volcanic rocks (U –Pb: 534 + 10 Ma, Ducrot & Lancelot 1977). The youngest manifestations of igneous activity in the BAG inlier are dolerite dykes (Ar –Ar: 196 + 1.8 Ma; Sebei et al. 1991). The supracrustal rocks of the BAG inlier experienced polyphase Pan-African deformation. The first recognizable event, D1, overprints the midNeoproterozoic ophiolite complex (Odin 1994) and the Palaeoproterozoic basement. It represents a major tectonometamorphic orogenic event that accompanied southwestward obduction of the ophiolite complex (685 + 15 Ma, Clauer 1976; Leblanc & Lancelot 1980). In the West Bleida area (Fig. 1), the result was the thrusting of the ophiolite complex over a continental margin sequence of argillitic sediments that were metamorphosed to sericite schists. Ophiolitic rocks were dismembered by NW –SE- and NE–SW-trending regional faults. In the Neoproterozoic formations, D1 is defined by strong penetrative foliation and metre-scale intrafolial folds, and was accompanied by low-grade greenschist metamorphism dated at 623 + 10 Ma (U– Pb data, Leblanc & Billaud 1978). Syntectonic quartz diorites and the posttectonic granodioritic Bleida pluton (U –Pb, 615 + 12 Ma, Ducrot 1979) are also associated with this event. F1 folds are NE –SW trending and reflect a major N120-oriented compressional event. Ophiolite obduction was oblique and associated with sinistral-reverse strike slip movement along NE –SW-oriented thrust faults (Emran & Chorowicz 1992). The late Pan-African D2 event is characterized by N030-shortening resulting in N110- to

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N130-trending F2 folds verging to the NNE (Emran & Chorowicz 1992), as well as conjugate fault sets trending N060 and N170. These faults have been dated at 608 + 12 Ma (U–Pb data, Ducrot 1979).

Local setting Lithologies and stratigraphy. The West Bleida study area is located 180 km SW of the city of Ouarzazate and 8 km west of the old Bleida copper mine (Fig. 1). The stratigraphic sequence is dominated by intermediate to mafic volcanic rocks (andesite to andesite – basalt composition) and contains volumetrically minor amounts of Fe-rich and cherty horizons. The volcanic facies account for 65% of exposed rocks. Other rock types include small quartz diorite and tonalite stocks, microdiorite and dolerite dykes, and tectonic slices of gneiss, gabbro and pyroxenite. The distribution of the various lithological units and geological features of the study area is presented in Figure 2, and photomicrographs of the principal lithologies are displayed in Figure 3. Volcanic rocks are represented by two main facies: (1) a regional and volumetrically dominant banded aphanitic intermediate to mafic facies defined by alternating dark and light layers of actinolite –chlorite –epidote and carbonate– sericite –quartz, respectively (Fig. 3a); (2) a spotted facies of more restricted extent defined by large (3– 4 mm; Fig. 3b) poikiloblastic crystals of cordierite that locally replace the quartz–sericite – chlorite –carbonate assemblage, as well as devitrification textures and broken quartz–plagioclase crystals suggesting a crystal tuff protolith of intermediate composition. The spotted facies is confined to the southwestern part of the area (Fig. 2). Results of X-ray diffractometry show that hornblende in both facies is magnesio-hornblende, chlorite is chlinochlore, epidote is clinozoisite and plagioclase is almost pure albite. In weathered samples, hematite, goethite and clay minerals are also present. The banded facies is characterized by low SiO2 contents (46– 56%), moderate Al2O3 (average of 12%; Table 1), moderate MgO (average of 6%) and low TiO2 (average of 0.08%). This composition confirms an andesitic to basaltic composition for the protolith. The spotted facies reveals intermediate SiO2 contents (average of 57%), high Al2O3 (average of 17%), low MgO (average of 4.2%) and low TiO2 (average of 0.09%). The texture of these rocks suggests an origin similar to that of cordierite-porphyroblastic Archaean rocks known as ‘dalmatianite’ in the Noranda district of Canada that have been interpreted as evidence of metamorphosed Mg-metasomatism (Riverin & Hodgson 1980).

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Fig. 2. Local geological features of the study area with the main structural and lithological units, and stereographic projections of some structural elements (lower hemisphere).

At West Bleida, however, subsequent leaching may have affected the rocks and further investigation is required. Iron-rich horizons (reddish colour, crumbly to weakly consolidated) in both spotted and banded volcanic units measure 5 – 30 m thick with horizontal and downdip extensions of up to 100 m (confirmed by drilling). The rocks are characterized by specular hematite – chlorite – quartz – epidote assemblages and 20 – 30% SiO2, suggesting either a banded iron formation protolith or Fe-rich volcanic rocks. Most primary textures were destroyed by deformation and metamorphic recrystallization. Copper-rich chert horizons (grey colour, massive and microcrystalline) are also present, typically comprising a quartz – sericite – calcite – albite

mineral assemblage with lesser amounts of chlorite, malachite, epidote, oxides and leucoxene. The horizons measure 0.4 – 5 m thick with horizontal and downdip extensions of up to 50 m (confirmed by drilling). Lithogeochemical analyses reveal a high SiO2 content (up to 70 wt%) and copper enrichment (3 – 4% Cu). They probably represent chert beds or hydrothermally silicified horizons. Stratigraphic correlations within the BAG inlier show that the West Bleida volcanic rocks correspond to the upper part of the ophiolite sequence (Saquaque et al. 1989; Hefferan et al. 2000). The units form a homoclinal east –west-trending structure with a moderate dip to the north. To the south, a major fault system marks the contact with


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Fig. 3. Representative photomicrographs showing the principal lithofacies of West Bleı¨da: (a) banded volcanic rocks; (b) spotted lithofacies showing deformed and retrograte altered cordierite; (c) an aphanitic and altered microdiorite dyke; (d) gneissic rocks showing evidence of high strain; (e) typical breccia lithofacies; (f) Fe-rich lithofacies with hematite layers (black) and quartz-rich layers (white).

metamorphosed sedimentary rocks (sericite schists) of the continental margin (Fig. 2). All rocks are weakly to moderately schistose and are crosscut by major long-lived fault systems known as Rouimiat 1 and Rouimiat 2 (see the ‘Structural geology’ section). Dykes and small stocks intrude the volcanic rocks. Most dykes are nonfoliated microdiorites and dolerites strongly altered to hematite and chlorite (see Fig. 3c). They trend mainly NW –SE or NE–SW (Fig. 2). Some weakly deformed north– south-trending aplite dykes occur in the central

part of the area. Small irregular stocks of strained quartz diorite and tonalite that occur within fault zones are medium- to coarse-grained and lack metamorphic aureoles. In the southwestern part of the study area, a tectonic slice of gabbro-pyroxenite lies adjacent to the Rouimiat 1 Fault (Fig. 2). A metamorphic assemblage of chlorite, amphibole, carbonate, epidote and biotite characterizes this lithological entity. The gneissic rocks are quartz– feldspar-rich and biotite-rich, display high-strain textures (Fig. 3d), and probably represent the remnants of a sedimentary protolith.

Table 1. Chemical analyses of the various lithologies in the West Bleida area Qtz dio

Gneiss SD

(%) SiO2 Al2O3 MnO MgO CaO Na2O K2O P2O5 TiO2 Fe2O3 LOI SUM

56.22 13.30 0.10 4.46 6.61 3.34 1.63 0.21 0.70 7.62 4.85 98.85

(ppm) Au Ag As B Ba Be Bi Cd Co Cr Cu Ge Li Mo Nb Ni Pb Sb Sn Sr V W Y Zn Zr

0.07 0.26 1.78 16.67 354 0.73 5.73 1.26 25 77 48 1.15 25.80 5.53 8 32 9 1.50 5.75 259 61 4.13 20.99 84 75

Number:

Amphibolite

Bd vol

Sp vol

Fe-rich vol

Cu-rich vol

1

1

Average 13

SD

Average 11

SD

Average 12

SD

Average 13

SD

8.13 4.45 0.05 3.59 4.26 2.41 0.51 0.18 0.78 4.32 3.14 0.55

61.89 10.92 0.20 2.21 9.57 0.01 0.14 0.26 0.73 11.16 2.77 99.60

47.48 16.52 0.22 2.61 15.83 0.15 0.14 0.15 0.54 10.29 3.40 97.19

54.45 15.56 0.16 5.58 7.64 1.92 0.73 0.19 0.73 9.12 3.56 99.47

7.40 3.46 0.05 2.55 3.96 1.22 0.51 0.06 0.30 1.89 0.75 0.25

55.76 14.91 0.15 4.71 6.54 3.06 1.09 0.18 0.71 8.74 3.63 99.30

6.86 1.16 0.05 1.77 4.28 1.70 0.58 0.06 0.13 3.13 1.75 0.47

50.61 14.48 0.16 4.43 9.04 1.20 1.85 0.15 0.67 9.69 7.14 99.26

10.15 4.26 0.08 2.96 8.22 1.10 1.79 0.07 0.37 3.64 9.53 0.82

59.27 13.33 0.14 4.58 5.60 2.17 1.16 0.20 0.68 7.66 4.76 99.36

12.49 5.23 0.05 3.11 5.22 1.94 1.12 0.10 0.34 2.98 3.57 0.51

0.08 0.42 1.87 31.95 145 0.48 8.48 1.29 13 98 26 1.16 17.36 3.36 2 28 15 0.02 6.49 189 33 5.25 23.04 44 68

0.06 0.1 0.5 0.5 107 1.0 1.0 0.6 50 127 206 0.2 9.3 6.8 28 104 166 24.0 1.5 160 260 1.5 25.7 91 19

0.06 0.1 16.8 0.5 53 1.3 1.0 1.6 27 48 25 7.5 9.6 11.8 27 25 1 1.5 5.2 1850 159 1.5 12.4 78 38

0.60 0.55 3.56 0.93 175 0.87 3.72 2.14 40 235 43 2.99 26.71 7.61 17 121 16 1.50 4.44 195 110 2.59 16.46 102 37

1.64 1.13 3.52 1.36 124 0.21 3.20 1.40 13 221 29 4.12 21.70 3.16 7 82 17 0.03 5.08 93 34 1.94 6.03 26 45

2.24 0.05 7.97 18.97 212 1.13 6.67 2.82 34 134 84 0.21 37.15 7.08 15 82 40 1.50 3.21 307 98 2.96 16.04 107 29

7.07 0.21 10.06 27.46 64 0.25 5.68 1.14 10 186 111 0.21 12.56 1.63 3 66 19 0.30 4.17 154 25 4.50 2.31 23 22

0.23 3.19 18.42 1.01 422 0.92 23.92 1.97 44 188 657 2.69 20.17 8.62 15 76 29 4.67 13.94 387 133 2.68 15.55 114 26

0.34 8.91 38.30 1.76 356 0.55 55.28 1.61 35 183 1647 3.62 12.36 3.77 8 51 42 7.00 29.80 373 62 3.54 8.57 55 40

1.27 0.70 2.94 2.38 244 0.99 7.65 1.92 36 180 128 0.80 27.60 7.71 12 88 27 1.50 4.23 183 94 7.93 12.43 104 33

2.04 1.70 2.97 4.58 195 0.53 11.95 1.64 17 222 174 1.29 18.72 2.11 7 74 18 0.03 4.96 138 46 11.20 4.49 35 41

SD, standard deviation; qtz dio, quartz diorite; vol, volcanic rocks; Sp, spotted; Bd, banded; LOI, loss on ignition.

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Average 4

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Facies:


Structural geology Polyphase deformation in the West Bleida area is expressed by the subvertical attitude of the volcanic strata, the development of ductile planar and linear elements, the presence of local mesoscopic folds and veins, and brittle events that include brecciation and the development of two major fault systems and related subsidiary faults (Fig. 2). The dominant D1 fabric in the area is S1 schistosity. It is defined everywhere by the planar alignment of sericite, chlorite and quartz –feldspar aggregates, and ranges from weakly developed to strongly penetrative. It has an average regional strike of N280 and a dip of 608 to the north (Fig. 2), indicating north–south regional shortening. Both banded and spotted volcanic rocks display transposition of compositional layering into S1. The schistosity is better developed within the spotted volcanic unit owing to its rheological properties (.70% quartz þ sericite). The S1 schistosity is weakly developed within competent mafic dykes, where deformation is more evident as fracture networks; both styles of D1 deformation are more evident in dykes that are subparallel to the general S0 stratigraphic trend and boudinaged. A weakly developed stretching lineation (Lm) defined by quartz –feldspar –mica mineral elongation is associated with the S1 fabric and plunges moderately to the SW (Fig. 2). Brittle –ductile shears in the area are related to D1 deformation. These shear zones are marked by intensification of the S1 fabric and locally developed protomylonitic textures and intercalated chlorite schist. The irregular and discontinuous shear zones strike N080 –N100 and locally host syntectonic goldbearing quartz veins. The angular relationships between the S1-parallel shear zone walls and S –C fabrics indicate a dominantly reverse movement toward the SSW. Small-scale shears (,5 cm thick) form a metre-scale conjugate network within some of the heterogeneous and altered lithologies. Veins of calcite, chlorite and quartz locally fill these small shear zones. Both dextral and sinistral movements are indicated by small offsets of several lithological markers. Outcrops of volcanic rocks display asymmetric metre-scale folds that range in shape from open to isoclinal. The F1 structures represent intrafolial syn-S1 folds that display Z-asymmetry and affect all types of volcanic rocks. They trend NE–SW (Fig. 2) and reflect a regional N120-oriented D1 shortening event (Leblanc & Billaud 1978). The D1 event was accompanied by low-grade metamorphism as demonstrated by the predominance of chlorite –epidote–albite – actinolite assemblages. Metamorphic intensity increases in proximity to the quartz diorite and tonalite stocks and locally reached amphibolite facies (Cisse´ 1989).

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The late Pan-African D2 event manifests as S2a, S2b and F2 folds within volcanic rocks, and as the sinistral lateral offsets of dykes and small quartz diorite stocks along N045-trending faults (Fig. 2). Conjugate crenulation cleavages, S2a and S2b, occur within high-strain zones and may define the axial plane of F2 folds. Visible F2 structures fold S1 schistosity into open to isoclinal, asymmetric metre-scale folds that trend N110– N130. They are the product of a D2 NE–SW to NW–SE shortening episode. Kink bands define a rare S3 fabric that is developed along narrow, discrete small-scale shear bands. Breccia zones observed throughout the area represent an important host for gold mineralization. The zones strike NE –SW to NW– SE; they measure 1–5 m in width and have a maximum horizontal and downdip extension of 200 m (see Fig. 4). The breccias are polymictic and contain angular to rounded, rotated fragments of foliated volcanic rocks in a matrix of quartz, calcite, iron oxides, malachite and chalcopyrite (Fig. 3e). Intense silicification and open-space fillings of gangue minerals, with or without sulphides, characterize some of the breccias. Using textural criteria defined by Je´brak (1997), these breccias are interpreted as tectonic in origin, with later collapse and hydrothermal infilling. The east – west-trending oblique-reverse faults, Rouimiat 1 and Rouimiat 2 (R1 and R2 in Fig. 2), are the largest structures in the area. The faults are well exposed at the surface and form linear crests of north-dipping carbonateand serpentine-bearing rocks that extend for more than 6 km in strike length (Fig. 2). The Rouimiat faults crosscut all other veins and faults. They truncate volcano-sedimentary sequences and displace a NNE – SSW-trending Jurassic dyke and the breccia zones. Striae and other brittle structural elements indicate dextral and sinistral movements. A dominant oblique reverse movement is also suggested by the angular relationship between the Rouimiat faults and the subsidiary fault sets (concave towards the WSW), and by the 200 m dextral displacement of a 1 m thick garnetiferous tuff horizon acting as a stratigraphic marker unit. The secondary sets of brittle faults (Fig. 2) trend N045, N070 and N280 to N300, and dip moderately to the north. They contain striae that are consistent with a composite reverse – horizontal displacement. The faults affect quartz veins, intrusive rocks and gneisses, and are typically barren or uneconomic with respect to gold. The Rouimiat 1 and Rouimiat 2 fault systems are in turn overprinted by a system of subhorizontal undulatory faults marked by weakly auriferous quartz and malachite veins of 4 – 6 cm width.

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Fig. 4. (a) Simplified geological map through the breccia zone, showing ore zone location relative to breccia. (b) Simplified geological cross-section E5 within the breccia zone. (c) Schematic cross-section illustrating the deformed quartz vein (qv) at the contact of mafic dyke (md). (d) Schematic cross-section showing the geometry and age relationship between an Au– quartz vein and late fault-filled malachite (Fma). ss, slip-surface; chlz, chlorite zone; bv, banded volcanic rocks; tr, trench; loc, location.

Mineralization Gold is present in anomalous quantities in weathered volcanic and intrusive rocks throughout the West Bleida study area (background values of 80–90 ppb Au; Reminex 2004). This auriferous zone was discovered in 1996 when the MANAGEM Company conducted a regional stream and soil sediment geochemistry exploration programme followed by trenching and sampling

of local targets. In 1998, MANAGEM announced a measured resource of 2.5 million tonnes at 2 g t21 Au for the West Bleida deposit based on trenching results and follow-up definition drilling. The deposit defines an east –west-trending zone of 500–600 m width and 1 km lateral extent. Gold mineralization occurs in veins and as disseminations. It is inhomogeneous at the scale of the study area in terms of grade distribution, nature of mineralization and related alteration. Figures 4a,b


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Fig. 5. Photomicrographs of gold-bearing vein: (a) a weakly weathered vein sample, with presence of chalcopyrite –quartz–gold assemblage and malachite; (b) quartz–chalcopyrite vein partially altered to hematite; (c) weathered gold– quartz vein hematite-filled microfractures as well as gypsum; (d) surficial alteration of gold–quartz vein with hematite and rare miersite and BaSO4 alteration assemblage; SEM-EDS analysis; (e) typical texture of specular hematite with the gold-quartz vein; SEM-EDS analysis. Cp, chalcopyrite; Qz, quartz; Ma, malachite; Gy, gypsum; Hem, hematite; Cal, calcite; Fe(ox), iron oxide.

and 7 present plan and cross-sectional views of drilled ore zones.

Disseminated gold Rocks containing disseminated gold constitute the bulk of the ore (Figs 4a,b and 7). Assays range from 0.5 ppm to 22 g t21 Au with an average grade of 3.2 g t21 Au. Disseminated-type ore zones, defined by the presence of free gold and minor sulphides, are lens-shaped, subparallel to S1, and generally dip to the north. They measure several metres to tens of metres thick, and have a horizontal length of 1–50 m and a downdip extent of up to 200 m. Zones of disseminated gold surround, or are adjacent to, auriferous quartz veins (see below; Fig. 7). The results of trench sampling confirm a correlation between gold and particular lithologies or brecciation, with the highest-grade zones preferentially concentrated in Cu-rich, stratabound cherty-textured quartz horizons (average 5.6 g t21 Au, n ¼ 25), Fe-rich volcanic lithofacies

(average 3.8 g t21 Au, n ¼ 34), and breccia zones (average 2.8 g t21 Au, n ¼ 320). Figure 8 presents the results from two such trenches. The nature of the auriferous disseminated zones varies according to the host lithology, but they are typically characterized by the presence of sulphides (,1%; mainly chalcopyrite and lesser pyrite), an increase in the friability of the host rocks, and the development of alteration-related mineralogical changes (see the ‘Hydrothermal alteration’ section).

Gold-bearing quartz veins At the regional scale, auriferous quartz veins are grouped into Z-shaped lenses. The veins are hosted by fault zones that create wide deformation corridors offset by post-D2 faults (Fig. 2). Veins strike east –west and have variable dips to the north or south, but mostly to the north (Fig. 2). They range in thickness from 2 cm to 2 m and have lengths up to 100 m. Veins are laminated

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Fig. 6. Various types of occurrences of gold and palladium and their variable shapes. (a) Typical weathered gold grain showing an irregular porous particle and infilling of Fe-oxides in the pits. (b) Free gold grains disseminated within a mafic protolith. Hbl, Hornblende; Chl, chlorite; SEM-EDS analysis. (c) Gold grain spatially (and paragentically) associated with palladium (Sb, As) in Fe-rich mafic volcanic rocks. (Note the smooth and mutual boundary between gold and palladium.) SEM-EDS analysis. (d) Gold hosting palladium inclusions in a breccia zone; the palladium displays secondary (surficial) redistribution of Cu and Te. (Note the intergrowth of specular hematite laths with the Au– Pd association.) SEM-EDS analysis. (e) Palladium (As, Sb, Cu, S) inclusions within free gold grain. SEM-EDS analysis. (f) Detail of last photograph showing a spatial distinction between Cu– S and Pd (As, Sb) phases in the Au– Pd association. SEM-EDS analysis. (g) Gold enclosed in clinozoisite within a breccia zone. SEM-EDS analysis. (h) Quartz vein showing weathered in situ gold in microfracture intersections.


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Sample location in metres

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Cu-rich chert Fe-rich lithofacies

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Sample location in metres

Fig. 7. Idealized cross-sections of ore zones of West Bledia from drill core, within (a) secondary dispersion halo around Au– quartz veins and (b) ore zone confined to Fe-rich lithofacies. vol, volcanic rocks; DDH, diamand drill hole.

and occur in the cores of brittle shear zones, typically oriented subparallel to S1. The veins contain thin slivers of foliated wall rocks. They show evidence of various degrees of strain, indicating that the vein material was introduced during several stages of deformation. The shear veins are boudinaged and folded by F1 and F2 (Fig. 4c). They are crosscut by N150-trending microdiorite dykes as well as by breccia zones. Veins have an average gold content of 0.65 ppm Au. The least weathered auriferous veins observed in drill intersections more than 80 m below the surface contain quartz, minor carbonates, pyrite (1–2%) and small amounts (,1%) of Cu-sulphide minerals including chalcopyrite, covellite and chalcocite (Fig. 5a and b). Intensely oxidized quartz veins (ferruginous quartz) are characterized by specular hematite, malachite, gypsum and traces of barite and miersite (Fig. 5c– e). Hematite, the dominant iron oxide in the veins, is observed in drill core to 150 m depth.

Weakly auriferous veins Two types of post- and late-tectonic (D1) economically barren quartz veins occur throughout the area:

Fig. 8. Diagrams illustrating the correspondence between Au contents and lithological facies within (a) trench 51 and (b) trench 33.

malachite – quartz veins and quartz-only veins. Malachite –quartz veins occur within narrow, subhorizontal late tectonic faults and local small-scale shears. Malachite is massive and locally exceeds 90% of the vein material. The veins are 3 –10 cm in thickness and 10– 50 cm in length. Malachite is less abundant below depths of 50 m. Malachite – quartz veins reveal low gold grades with an average value of 130 ppb Au. Analyses of deeper (fresher) drill intersections indicate that these veins are not surrounded by a secondary gold halo. Quartz-only veins occur in en echelon arrays (up to 2 m wide and 15 m in lateral extent) that crosscut gold-bearing veins. They trend north–south and have a subvertical dip. Locally the veins display centimetre-scale dextral or sinistral offsets along S3 foliation, suggesting late reactivation of S3 foliation. The weighted average value of 16 samples is very low (80 ppb Au).

Gold grain chemistry and textures Reflected light microscopy and scanning electron microscope –energy dispersive spectrometry (SEM-EDS) analysis of 150 samples and more than 200 gold point analyses revealed that gold

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grain sizes range from 10 to 40 mm within some gold-bearing quartz veins, and from 20 to 700 mm in disseminated zones within Fe-rich, Cu-rich cherty, or brecciated horizons. SEM-back-scattered electron (BSE) analyses of gold grains from both veins and disseminated mineralization have revealed the presence of palladium-bearing intergrowths. Pd-bearing grains appear restricted to zones more than 80 m below the surface and occur in association with metamorphic minerals (clinozoisite, chlorite and amphibole). Preliminary investigations of the Pd –Au association indicate that Pd inclusions (up to 2 g t21 Pd) are present as Pd– As –Sb, Pd–Bi –Se and Pd– Te phases (Fig. 6c–f). Copper minerals in the Pd-rich gold mineral assemblages include covellite, chalcopyrite and malachite (Fig. 6e and f). A second type of gold grain, generally restricted to the upper 80 m of the deposit, lacks inclusions and typically occurs in association with hematite and white mica, filling S1 foliation planes and other microfractures (Fig. 6h). This type of gold presents a wider range of habits and sizes, ranging from 10 to 700 mm wide, and grains may show stressed surfaces with goethite and clay minerals filling micro-pits (Fig. 6a and h). Analyses reveal the gold to be electrum; Au:Ag ratios and Ag contents are similar (,1:10; 4 –20 wt% Ag) for gold from veins and disseminations. Four samples with visible gold grains from breccia zones and Cu-rich chert have been studied using a SX 50 electron microprobe at McGill University in Quebec, Canada. The 22 point analyses confirm the presence of two types of gold: (1) gold as electrum (10% Ag and 90% Au) observed in four samples; (2) gold as an Au–Ag– Pd alloy, typically bordered by small grains of Pd and Bi (Te, Sb) phases (Fig. 9). Isomertieite, Pd11(Sb2,As2), was identified but appears restricted to samples within the Cu-rich cherty horizons.

Fig. 9. SEM– BSE image of gold grain from the Cu-rich chert.

Paragenesis Textures and structural relationships, including the spatial association with shear zones subparallel to S1, the preferred orientation of sericite and chlorite in quartz vein inclusions, and deformation by D1 folding and boudinage, collectively suggest pre- to early syn-D1 growth for disseminated and vein-type gold. Disseminated gold from this primary event contains inclusions enriched in Pd, As, Sb, Te, Bi, Cu and Se. Gold grains that lack inclusions and occur in association with hematite and white mica are interpreted as primary Pd-rich gold grains that were intensely altered by oxidizing surface fluids. Despite the wide range of secondary gold forms and habits, the original Au:Ag ratio was not markedly affected by supergene processes. The primary relationship between vein and disseminated gold is severely overprinted by D1 and D2 deformation events that structurally transposed both goldbearing ore types. However, vein gold is locally observed crosscutting Cu-rich cherty and Fe-rich volcanic rocks that contain disseminated gold, suggesting that at least some of the auriferous quartz veins post-date disseminated mineralization. Copper is certainly primary in this paragenesis, existing with palladium as inclusions in gold and as a stable mineral with gold and pyrite (Figs 5a,b and 6f).

Hydrothermal alteration Primary gold mineralization was accompanied by extensive hydrothermal alteration of the host rocks resulting in net changes in whole-rock chemistry. The primary hydrothermal alteration halo is obscured by supergene alteration and weathering in the first 50 –80 m below the surface.

Geochemical mapping of altered rocks Mineral assemblages in the least-altered mafic rocks are typical of greenschist-facies conditions: albite, chlorite, sericite, epidote, actinolitic hornblende, quartz and calcite, with minor amounts of biotite and hematite. Original textures are generally preserved away from structural corridors and gold zones (Fig. 3a and b), but fluid activity in hydrothermally altered zones destroyed primary textures and promoted the development of penetrative schistosity as a result of mineralogical changes. Hydrothermal activity is reflected by an overall increase in quartz, chlorite, sericite, hematite, titanite and calcite (Fig. 5d and e). The normative mineral index method of Piche´ & Je´brak (2004), successfully used to quantify alteration at the regional scale in the Abitibi Subprovince of Quebec (Canada), was used to quantify the type


and intensity of alteration affecting the volcanic sequence in the West Bleida study area. The available data include: (1) a set of 850 samples that were analysed by inductively coupled plasma atomic emission spectrometry (ICP-AES) for major and trace elements by MANAGEM during the earliest phases of exploration; these data cover the entire area and are representative of different rock types; (2) a new set of 55 samples from various units that were analysed by X-ray fluorescence (XRF) for major and trace elements at X-Ray Assay Ltd. in Ontario, Canada (Table 1). The widespread coverage of the data allowed normative isograd index maps to be constructed across the entire study area. Chlorite (CI) and sericite (SI) indexes are defined as CI ¼ ½ðMgO þ Fe2 O3 Þ=ðMgO þ Fe2 O3 þ CaO þ SiO2 Þ 100

ð1Þ

SI ¼ ½K2 O=ðK2 O þ Na2 OÞ 100:

ð2Þ

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Figure 10 displays normative index maps for sericite and chlorite alteration. The maps demonstrate that pervasive silicification and hematization are localized, occurring in narrow zones that are either structurally or lithologically controlled. Although the geometric distribution of chlorite and sericite alteration reflects the combined effects of metamorphic and hydrothermal processes, the normative index method allowed the two alteration types to be distinguished. The normative chlorite index depicts the southern sericite schist domain as a homogeneous low separated from the volcanic rocks by the Rouimiat 1 Fault (Fig. 10b). North of Rouimiat 1, chloritization displays an east–west structural grain, corresponding to either massive hydrothermal chloritization or retrograde replacement of chlorite by sericite – calcite –quartz assemblages; the latter was observed by petrographic analysis and is considered more likely. Sericitization is most prominent in the southern sericite schist domain and in spotted volcanic rocks (Fig. 10a). In both facies (spotted volcanic

Fig. 10. Isograd maps of normative (a) sericite and (b) chlorite within the West Bleida area. Maps reveal the north continental margin sericitic schist and confirm the viability of the normative calculations.

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and sericite-schist), sericite is dominantly metamorphic, as reflected by its ubiquitous and widespread distribution. However, the sericite index map also defines narrow east –west-trending domains that coincide with quartz veins and Cu-rich cherty and Fe-rich volcanic units. These narrow domains are interpreted to be hydrothermal in origin. Figure 10a also reveals a new NW–SE-trending corridor that was not observed during field mapping. This corridor of medium- to high-grade sericitization is transposed into S1 and may represent a metamorphically segregated layer or a fine-grained volcanic rock. At the local scale, intensely mineralized and altered rocks associated with both styles of gold mineralization (vein-type and disseminated) were investigated. Changes in colour and mineralogy were compared with lithogeochemical data, and a petrographic study characterized the main alteration facies prior to normative analyses. The north –south sericite index profile intersects several pervasive alteration zones (index .20). Strong sericitization is closely associated with auriferous quartz veins, whereas sericitization and silicification characterize the copper-rich cherty zones. Silicification as microcrystalline quartz and local quartz infillings characterizes parts of the breccia zones. In thin sections, this relationship is shown by the intense replacement of metamorphic chlorite by sericite, clay minerals and quartz. The chlorite index profile confirms the dominantly metamorphic character of this mineral, which is supported by the relatively high abundance of chlorite (CI . 25) throughout the various facies in the area. Despite the widespread distribution of normative chlorite, the index profile reveals some pervasively altered zones. Most of these zones are spatially related to the banded and sheared volcanic rocks.

Discussion The latest study at the West Bleida gold deposit suggested that it represents a structurally controlled epithermal system related to an unidentified magmatic source (Barakat et al. 2002). This model is based on a fluid inclusion study that indicated that quartz veins (undifferentiated) formed in a geothermal environment in which P–T conditions decreased from 0.5 kbar and 300 8C to 40 bars and 150 8C at an approximate depth of 2 km (Barakat et al. 2002). However, very few true epithermal deposits are known in ophiolite environments (Sillitoe & Hedenquist 2003), and the West Bleida area generally lacks the geological context and felsic volcanic rock association that are key characteristics of adularia–sericite or acid –

sulphate epithermal gold systems (Heald et al. 1987; Cooke & Simmons 2000; O’Brien et al. 1999). The West Bleida gold deposit does not display evidence of any pre-tectonic open-space fillings and presents a rather unusual Pd –Au signature. Other types of depositional environment may be considered: iron oxide –gold– copper deposits (IOCG, or Olympic Dam-style; Haynes 2000), orogenic gold deposits (Olivo et al. 1995; Groves et al. 1998) and volcanogenic massive sulphide (VMS)-related stockworks (Franklin et al. 2005). In the well-documented Brazillian Au–Pd occurrences (e.g. the Conceiçaõ mine; Olivo et al. 1995) located in the Itabira district, southern Saõ Francisco craton, gold and palladium mineralization is hosted by altered iron formation and consists of a single quartz vein parallel to the S1 foliation. The quartz vein is deformed and altered to Fe hydroxides and kaolinite; the deposit is interpreted as an epigenetic model for gold and palladium concentration (Olivo et al. 1995). The very specific assemblage of the Bleida gold mineralization recalls the Pd –As– Sb and other complex phases, such as Se- and Te-bearing minerals, identified in the gold grains of the Conceiçaõ deposit. Such a mineral association is consistent with highly oxidizing saline fluids that are able to carry Cu, Au and PGE complexes at various temperatures. IOCG deposits have been described in mafic environments, such as the Salobo (Brazil) or Ernest Henry (Australia) deposits. In Western Nevada, the Humbodlt Mountains district displays some Cu–Au– Fe occurrences associated with a Jurassic mafic intrusive body (Barton & Johnson 1996) and regional sodic alteration. This style of deposit usually displays both structural and magmatic associations. West Bleida does not display evidence of large iron concentrations, nor does it appear to contain other key minerals (fluorite, barite and scapolite) that are often associated with this style of deposit. However, some similarities could be related to the same oxidizing environment. Orogenic gold deposits are known in the AntiAtlas Province (e.g. Akka Gold Mine) as syn- to post-tectonic mineralization related to major shear zones. These deposits may contain PGE in association with quartz, pyrite and pyrrhotite. Similar associations have been described in Muruntau, Sukhoi Log, Waterberg (Transvaal) and Chudnoe (Polar Ural) within fuchsite –albite-rich rocks that suggest a mafic to ultramafic environment (Distler et al. 1996). Fluid inclusions in these cases indicate an abundance of saline fluids. Although the tectonic style of the West Bleida mineralization does not appear to fit perfectly with an orogenic gold environment, it remains possible that the mineralization was partly or totally syntectonic. VMS-related


stockworks are enriched in gold in several ophiolitic environments, such as Cyprus and Oman (Herzig 1988; Lescuyer et al. 1988). Although not abundant, the PGE can be recuperated at the end of the electrolytic ore extraction process. The Lasaill mine in Oman, for example, makes that country a small-scale PGE producer. The West Bleida deposit displays several features that are reminiscent of VMS-style mineralization: spatial coincidence between gold dissemination zones and specific stratigraphic horizons, stockwork and disseminated ore types, and dalmatianite-style alteration. In this scenario, the Cu-rich chert would be genetically related to the Bleida ophiolitic environment (Leblanc & Billaud 1978). Nevertheless, the lack of a zinc anomaly is puzzling. Further work is needed to demonstrate fully the pre-tectonic setting of all the region’s deposits and the intensity of remobilization processes that may have occurred during late Pan-African deformation. Undoubtedly, the future underground mining developments on the Bleida Au-project will lead us to better understanding of the geometry and crosscutting age-relationships of Au–Pd mineralization and eventually allow us to make comparisons with other well-known Au–Pd deposits.

Conclusions At West Bleida, gold mineralization is associated with Cu and Pd and is structurally and lithologically controlled. Gold occurs principally in quartz veins and disseminations within mafic rocks of the Bleida ophiolitic accretionary complex. Gold grains display several habits reflecting a primary hydrothermal event and later overprinting by surficial alteration and weathering. The structural style suggests that primary gold mineralization represents the root of a pre-tectonic VMS system, and/or a late tectonic mesothermal system. The West Bleida gold mineralization within the Neoproterozoic volcanic rocks of the central AntiAtlas represents an unusual style of Au mineralization in Morocco, as the orogenic character of gold is not dominant as in the Akka gold deposit (western Anti-Atlas) or in the Tiouit gold deposit (eastern Anti-Atlas). The presence of gold in both lithological and structural settings implies a complex geological history for gold implacement and must be taken in account in further gold exploration programmes. The authors wish to express their thanks to H. Gibson, D. Lentz and C. Beaumont-Smith for their constructive comments on earlier drafts of this paper. We also thank the Reminex staff for their financial support and for providing access to properties and documents. Constructive

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and very through reviews by R. Chapman and C. Marignac are gratefully acknowledged.

References B ARAKAT , A., M ARIGNAC , C., B OIRON , M. C. & B OUABDELLI , M. 2002. Caracte´risation des paragene`ses et des paleócirculations fluides dans l’indice d’or de Bleida (Anti-Atlas, Maroc). Comptes Rendus Geósciences, 334, 35–41. B ARTON , M. D. & J OHNSON , D. A. 1996. Evaporiticsource model for igneous related Fe-oxide (REE– Cu– Au) mineralization. Geology, 24, 259–262. C ARD , K. D., P OULSEN , K. H. & R OBERT , F. 1989. The Archean Superior Province of the Canadian Shield and its lode gold deposits. In: K EAYS , R. R., R AMSAY , W. R. H. & G ROVES , D. I. (eds) The Geology of Gold Deposits: the Perspective in 1988. Economic Geology Monography, 6, 19– 36. C HARLOT , R. 1982. Caracte´risation des e´veńements e´burneéns et panafricains dans l’Anti-Atlas marocain: Apport de la me´thode geóchimique Rb/Sr. Notes et Me´moires du Service Geólogique du Maroc, 313. C ISSE´ , A. 1989. Geóme´trie et cine´matique de la de´formation Panafricaine majeure dans le district minier de Ble´¨ıda (Bou Azzer –El Graara, Anti-Atlas central, Maroc). PhD Thesis, Unive´rsite´ de Marrakech. C LAUER , N. 1976. Geóchimie isotopique du strontium des milieux se´dimentaires. Application a` la geóchronologie de la couverture du craton ouest-africain. PhD Thesis, University of Strasbourg. C OOKE , D. R. & S IMMONS , S. E. 2000. Characteristics and genesis of epithermal gold deposits. Reviews in Economic Geology, Gold in 2000, 13, 221–244. D ISTLER , V. V., M ITROFANOV , G. L., N EMEROV , V. K., K OVALENKER , V. A., M OKHOV , A. V., S EMELKINA , L. K. & Y UDOVSKAYA , M. A. 1996. Modes of occurrence of the platinum group elements and their origin in the Sukhoi Log gold deposit (Russia). Geology of Ore Deposits, 38, 413–428. D UCROT , J. 1979. Datation a` 615 Ma de la granodiorite de Bleida et conse´quences sur la chronologie des phases tectoniques, me´tamorphiques et magmatiques pan-africaines dans l’Anti-Atlas marocain. Bulletin de la Socie´te´ Geólogique de France, (7), XXI, 495–499. D UCROT , J. & L ANCELOT , J. R. 1977. Proble`me de la limite Prećambrien– Cambrien. E´tude radiochronologique par la me´thode U –Pb sur zircons du volcan du Jbel Boho (Anti-Atlas marocain). Canadian Journal of Earth Sciences, 14, 2771– 2777. E MRAN , A. & C HOROWICZ , J. 1992. La tectonique polyphaseé dans la boutonnie`re prećambrienne de Bou Azzer (Anti-Atlas central, Maroc): apport de l’imagerie spatiale Landsat-MSS et de l’analyse structurale de terrain. Sciences Geólogiques, Bulletins, 45, 121– 134. F RANKLIN , J. M., G IBSON , H. L., J ONASSON , I. R. & G ALLEY , A. G. 2005. Volcanogenic massive sulfide deposits. Economic Geology, 100th Anniversary Volume, 523– 560. G ASQUET , D., L EVRESSE , G., C HEILLETZ , A., A ZIZI S AMIR , M. R. & M OUTTAQUI , A. 2005. Contribution to a geodynamic reconstruction of the Anti-Atlas

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A. BELKABIR ET AL.

(Morocco) during Pan-African times with the emphasis on inversion tectonics and metallogenic activity at the Precambrian– Cambrian transition. Precambrian Research, 140, 157–182. G ROVES , D. I. & F OSTER , R. P. 1991. Archean lode gold deposits. In: G ROVES , D. I. (ed.) Gold Metallogeny and Exploration. Blackie, Glasgow, 63–103. G ROVES , D. J., G OLDFARB , R. J., G EBRE -M ARIAM , M., H AGEMANN , S. G. & R OBERT , F. 1998. Orogenic gold deposits: A proposed classification in the context of their crustal distribution and relationships to other gold deposit types. Ore Geology Reviews, 13, 7 –27. H AYNES , S. J. 2000. Geology and Wine 2. A geological foundation for terroirs and potential sub-appellations of Niagara Peninsula wines, Ontario, Canada. Geoscience Canada, 27, 67– 87. H EALD , P., F OLEY , N. K. & H AYBA , D. O. 1987. Comparative anatomy of volcanic-hosted epithermal deposits: Acid– sulphate and adularia–sericite type. Economic Geology, 82, 1 –26. H EDENQUIST , J. W., A RRIBAS , A. A. R. & G ONZALEZ -U RIEN , E. 2000. Exploration for epithermal deposits. Reviews in Economic Geology, Gold in 2000, 13, 245– 278. H EFFERAN , K. P., A DMOU , H., K ARSON , J. A. & S AQUAQUE , A. 2000. Anti-Atlas (Morocco) role in Neoproterozoic Western Gondwana reconstruction. Precambrian Research, 103, 89–96. H ERZIG , P. M. 1988. A mineralogical, geochemical, and thermal profile through the Agrokipia B hydrothermal sulphide deposit, Troodos Ophiolite Complex, Cyprus. In: F REIDERICH , G. H. & H ERZIG , P. M. (eds) Base Metal Sulphide Deposits. Springer, Berlin, 182– 215. J E´ BRAK , M. 1997. Hydrothermal breccias in vein-type ore deposits: A review of mechanisms, morphology and size distribution. Ore Geology Reviews, 12, 111– 134. J UERY , A., L ANCELOT , J. R., H AMET , J., P ROUST , F. & A LLE` GRE , C. J. 1974. L’aˆge des rhyolites du Prećambrien III du Haut Atlas et le proble`me de la limite Prećambrien–Cambrien (abstract). Sciences de la Terre, 2nd Annual Meeting, Nancy, 230. L EBLANC , M. 1981. Ophiolites prećambriennes et gıˆtes arseńie´s de cobalt (Bou-Azzer, Maroc). Notes et Me´moires du Service Geólogique du Maroc, 280. L EBLANC , M. & B ILLAUD , P. 1978. A volcanosedimentary copper deposit on a continental margin of upper proterozoic age: Bleida, Anti-Atlas, Morocco. Economic Geology, 73, 1101– 1111. L EBLANC , M. & B ILLAUD , P. 1990. Zoned and recurrent deposition of Na– Mg–Fe–Si exhalites and Cu– Fe sulphides along synsedimentary faults (Bleida, Morocco). Economic Geology, 85, 1759– 1769.

L EBLANC , M. & L ANCELOT , J. R. 1980. Interpre´tation geódynamique du domaine pan-africain (Prećambrien terminal) de l’Anti-Atlas (Maroc) a` partir de donneés geólogiques et geóchronologiques. Canadian Journal of Earth Sciences, 17, 142– 155. L ESCUYER , J. L., O UDIN , E. & B EURIER , M. 1988. Review of the different types of mineralization related to the Oman ophiolitic volcanism. In: E L -G ABY , S. & G REILING , R. O. (eds) Proceeding of Seventh Quadrenniel IAGOD Symposium. Schweizerbart, Stuttgart, 489–500. O’B RIEN , S. J., D UBE´ , B. & O’D RISCOLL , C. F. 1999. High-sulphidation, epithermal-style hydrothermal systems in late Neoproterozoic Avalonian rocks on the Burin peninsula, Newfoundland: implications for gold exploration. Current Research, Newfoundland Department of Mines and Energy, Geological Survey, Reports, 99, 275– 296. O DIN , G. S. 1994. Geological time scale. Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 318, 59–71. O LIVO , G. R., G AUTHIER , M., B ARDOUX , M., L AEA˘ O D E S A´ , E., F ONSECA , J. T. F. & S ANTANA , F. C. 1995. Palladium-bearing gold deposit hosted by Proterozoic Lake Superior-type iron formation at the Caue˛ Iron Mine, Itabira district, Southern Sa˘o Francisco Craton, Brazil: geologic and structural controls. Economic Geology, 90, 118–134. P ICHE´ , M. & J E´ BRAK , M. 2004. Normative minerals and alteration indices developed for mineral exploration. Journal of Geochemical Exploration, 82, 59–77. REMINEX 2004. Rapport annuel, Far West, projet Bleida. Reminex, Marrakech. R IVERIN , G. & H ODGSON , J. 1980. Wall-rock alteration at the Millenbach Cu– Zn mine, Noranda, Quebec. Economic Geology, 75, 424–444. S AQUAQUE , A., A DMOU , H., K ARSON , J., H EFFERAN , K. & R EUBER , I. 1989. Precambrian accretionary tectonics in the Bou Azzer– El Graara region, Anti-Atlas, Morocco. Geology, 17, 1107–1110. S EBAI , A., F E´ RAUD , G., B ERTRAND , H. & H ANES , J. 1991. 40Ar/39Ar dating and geochemistry of tholeiitic magmatism related to the early opening of the central Atlantic rift. Earth and Planetary Science Letters, 104, 455– 472. S ILLITOE , R. H. & H EDENQUIST , J. W. 2003. Linkages between volcanic settings, ore-fluid compositions, and epithermal precious metal deposits. In: S IMMONS , S. F. & G RAHAM , I. (eds) Volcanic, Geothermal, and Ore-forming Fluids: Rulers and Witnesses of Processes within the Earth. Society of Economic Geologists, Special Publications, 10, 315– 343. T HOMAS , R. J., F EKKAK , A., E NNIH , N. ET AL . 2004. A new lithostratigraphic framework for the Anti-Atlas Orogen, Morocco. Journal of African Earth Sciences, 39, 217–226.

Petrology and geochemistry of the Neoproterozoic Siroua granitoids (central Anti-Atlas, Morocco): evolution from subduction-related to within-plate magmatism AHMED TOUIL1, AHMID HAFID1, JACQUES MOUTTE2 & ABDELMAJID EL BOUKHARI3 1

De´partement des Sciences de la Terre, Faculte´ des Sciences et Techniques, Gue´liz BP 549, Marrakech, Morocco (e-mail: [email protected]) 2

ENSMSE, centre SPIN, 158 cours Fauriel 42023 Saint-Etienne, France

3

De´partement des Sciences de la Terre, Faculte´ des Sciences, Semlalia, Marrakech, Morocco Abstract: The Siroua massif includes many plutons of Neoproterozoic age. The mineralogical and geochemical character of the plutons allows us to describe an evolution of the magmatism, in space and time, from a subduction-related type in the northern part, to a within-plate subalkaline type in the southern part. The first magmatic activity coeval with the closing of the Khzama oceanic basin in the north is little evolved and of oceanic type (dominantly gabbros and basalts). It is followed by a low potassic calc-alkaline magmatism (gabbro–diorites, tonalites and trondhjemites of Nebdas pluton) and by a voluminous highly potassic calcalkaline magmatism (Askaoun and Ifouachguel plutons) that marks the collisional period. The end of crustal uplift and the beginning of the extension is marked in the south by a sub-alkaline magmatism corresponding to the Ida ou Illoun, Imdghar and Affela N’ouassif granites. Magmatic activity, in the Siroua massif, is marked at the end of the Neoproterozoic (PIII) by a continental tholeiite with an alkaline affinity, which occurs as dykes crosscutting the Neoproterozoic granites, and later by dominantly alkaline granites.

The Siroua massif is a Precambrian segment of the Anti-Atlas Belt, located at about 100 km NE of Agadir, on the northern edge of the West African craton. It is limited to the north by the South Atlas Fault and to the south by the Anti-Atlas Major Fault, which represents the southern limit of the mobile belt that developed, during the Pan-African orogeny, on the northeastern edge of the West African craton (WAC; Choubert 1963; Leblanc 1976; Leblanc & Lancelot 1980) (Fig. 1). The Anti-Atlas Major Fault corresponds to the southwestern boundary of an aulacogen that formed along the northern margin of the WAC during the early Neoproterozoic and the actual northern limit of the WAC is located at the South Atlas Fault (Ennih & Lie´geois 2001, 2003). The Siroua massif comprises a basement of Palaeoproterozoic and Cryogenian age (PI and PII, after Choubert & Faure-Muret 1970) and a cover of Ediacaran and late Proterozoic (PIII) to Meso-Cenozoic ages. The Eburnean orogeny (c. 2000 Ma; Thomas et al. 2004) is only locally present, as lenses of gneisses and strongly foliated amphibolites found to the SE of Jbel Siroua (N’Kob district). However, the Pan-African formations (c. 750 –520 Ma, corresponding to Cryogenian and Ediacaran stages) crop out widely

throughout the Siroua district. Three main formations are distinguished. (1) The Bleida Group (of early Neoproterozoic age) is represented mainly in the eastern part of the study area, where plutonic formations (e.g. the Tourtit orthogneiss of Samson et al. 2004) are less developed. It comprises the Khzama and N’Kob ophiolitic complex and the volcanoclastic formations of the Imghlay series (El Boukhari 1991). A tonalite protolith from an orthogneiss of the ophiolitic sequence has been dated at 743 Ma (Sm –Nd, De Beer et al. 2000) and a plagiogranite from Tasriwine ophiolite has been dated at 762 þ1/22 Ma (Samson et al. 2004). The Imghlay series is affected by the major Pan-African event (660 Ma, Thomas et al. 2004), which is considered to have been responsible for the obduction of ophiolitic complexes on the West African craton (Villeneuve & Corneé 1994). The emplacement of the Khzama ophiolite has been dated at 663 + 13 Ma (De Beer et al. 2000). In the western part of the massif, where plutonic rocks dominate, the Bleida Group is represented by hectometre-scale quartzites, serpentinites and mafic to ultramafic amphibolites (Regragui 1997; Jouider 1997), which are enclosed within the plutonic formations.


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Fig. 1. Location of Siroua massif in the Anti-Atlas belt.

(2) The Sarhro Group (of Late Neoproterozoic age) comprises a volcanoclastic sequence (the Siroua Series) overlying unconformably the Bleida Group and intruded by granitoid rocks whose emplacement ages range from 680 to 600 Ma (Charlot 1982; De Beer et al. 2000; Thomas et al. 2000, 2004). The volcanoclastic series, whose thickness may reach locally 2000 m, is composed of basalts, andesites, sandstones, silts and conglomerates. They were accumulated in a back-arc basin that shows a horst and graben structure controlled by east –west normal faults. This structure is related to an extensional stage of late Cryogenian age (El Boukhari 1991; Regragui 1997; Gress et al. 2000). The deformation of the series during the late Pan-African stage is related to the closure of the Khzama oceanic basin (El Boukhari 1991). This stage has been dated at 615 + 12 Ma (Clauer et al. 1982) and 579 + 1.2 Ma (Inglis et al. 2004) in the Bou-Azzer inlier (central Anti-Atlas), where the Tidiline series has palaeogeographical characteristics (Saquaque et al. 1989; Leblanc & Moussine-Pouchkine 1994; Villeneuve & Corneé 1994) similar to those of the Siroua series. The late Pan-African stage does not induce significant metamorphic re-equilibration of the volcanoclastic series. The volcanoclastic series is intruded by numerous granitic plutons (Assarag suite). In the western part of Siroua, the Askaoun and Ida ou Illoun granites intrude respectively the northern and southern areas. The eastern part

comprises, in north, near Khzama, the Ait Nebdas, Tifratine, Ifouachguel and Tamassirt massifs, and, in the south, near N’kob, the Arg, Affela N’ouassif, Taloust and Imdghar massifs (Fig. 2). (3) The Ouarzazate Group (Late Neoproterozoic) consists of volcanic and volcanoclastic rocks unconformably overlying the Sarhro Group. These are intruded by pinkish granites and microgranites (Amassine, Tifnout and Immorghane granites) emplaced between 595 and 520 Ma (De Beer et al. 2000; Thomas et al. 2000, 2002, 2004). Several studies have already been devoted to the geochemistry and mineralogy of single plutons of the Assarag suite in the western part (Touil 1999; Touil et al. 1999b) and the eastern part of the area (El Khanchaoui 2001; El Khanchaoui et al. 2001), but comprehensive studies are currently lacking concerning the plutonic activity of the area, its evolution in space and time, and its possible relationships with the geodynamic history. The aim of the present study is to propose, based on a regional survey of the geochemical and mineralogical characters of the granitic rocks, a general scheme of the magmatic evolution of the Siroua massif in the Neoproterozoic. The plutons selected for the present study are located on both sides of the Khzama and N’Kob ophiolitic formations. In the eastern district, they consist of the Ait Nebdas and Ifouachguel plutons, located respectively on the northern and southern side of the Khzama ophiolites, and the Affela

NEOPROTEROZOIC SIROUA GRANITOIDS

267

Fig. 2. Geological map of Siroua belt with the location of studied plutons.

N’ouassif and Imdghar plutons, located to the south of the former. In the western district, the Assarag suite is represented mainly by the Askaoun and Ida ou Illoun plutons (Fig. 2).

Geological and petrographic characteristics of the rock types The crystalline basement of the Siroua massif is essentially composed of Assarag granitic suite. The largest plutons occur in the western part (Askaoun and Ida ou Illoun). Their intrusion induces in the Sarhro volcano-sedimentary group a contact metamorphism marked by the recrystallization of phyllites to large poikilitic crystals of muscovite and to biotite. The plutons are unconformably overlain by the Ouarzazate Group basal conglomerates, containing pebbles derived from them, and by volcanic rocks.

The Askaoun pluton The Askaoun complex constitutes a large batholith covering an area of about 35 km 22 km. Charlot (1982) obtained a whole-rock Rb/Sr isochron age of 699 + 10 Ma. De Beer et al. (2000) obtained a sensitive high-resolution ion microprobe (SHRIMP) U/Pb zircon age of 575 + 8 Ma. The dominant rock type is a dark-coloured monzodiorite–granodiorite, with a medium-grained

(,7 mm) equigranular texture. The batholith contains locally swarms of centimetre- to metre-scale enclaves, with rounded or lobate contours. The composition of the enclaves ranges from gabbrodiorite to quartz-diorite. The mineral assemblage in both rock types includes plagioclase (An50 – 20, up to An70 in the enclaves), clinopyroxene, amphibole, biotite, quartz and orthoclase. Orthopyroxene is found in some enclaves. Accessory minerals include apatite, zircon, magnetite and ilmenite. In the Tizkht and Tarniwine areas (respectively in the south and the NW of Askaoun), the granodiorites enclose earlier intrusions of foliated gabbro to diorite compositions (early gabbros and diorites), whose mineral assemblage comprises plagioclase, clinopyroxene, apatite, zircon, ilmenite and locally, orthopyroxene, amphibole, biotite or K-feldspar. The Amlouggi tonalites described by Thomas et al. (2002) in the Askaoun massif belong to the Ouarzazate intrusive rocks (Ediacaran intrusions), which intrude sharply the Askaoun granodiorite in the form of centimetre-scale veins or as intrabatholitic masses forming outcrops of about 50 m across (Touil 1999). These observations are in agreement with the age of 586 Ma obtained on the Amlouggi tonalites, and indicate that the age of 575 + 8 Ma obtained for the Askaoun granodiorite cannot be retained. Additional dating on the Askaoun granodirite, as indicated by De Beer et al. (2000), is imperative.

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The Ida ou Illoun pluton The Ida ou Illoun pluton, located in the south of Askaoun, covers an area of about 18 km 7 km. Ages of 610 + 13 Ma and 614 + 10 Ma have been obtained on this pluton, by a whole-rock Rb/Sr isochron (Charlot 1982) and by U/Pb on zircon (Thomas et al. 2000), respectively. The pluton comprises masses of basic rocks surrounded by acidic rocks of granodiorite and monzogranite compositions. The basic rocks (the Tamtattarn diorite of Thomas et al. 2002), which vary from gabbros to quartz-diorites, are found as hectometre-scale massifs and as metre-scale enclaves disseminated in the acidic rocks. Their texture changes from medium grained, in cores of larger masses, to fine grained toward borders and in the enclaves. Contacts with the host granites are either clear-cut or gradational. In the latter case, the presence of a zone of intermediate composition, resulting from the interaction between basic and acidic magmas, suggests that the two magmas were emplaced at approximately the same time. The mineral assemblage of the basic rocks comprises plagioclase (An60 – 35), amphibole and clinopyroxene as major phases, with orthopyroxene, biotite, ilmenite, apatite, zircon and titanite as minor phases. Quartz and orthoclase are sometimes present. The granodiorite is a coarse-grained equigranular dark-coloured rock composed of plagioclase (An45 – 20), amphibole, biotite, quartz and orthoclase, with zircon, apatite, allanite, ilmenite, magnetite, pyrite and pyrrhotite as accessories. The dominant rock type is a monzogranite (the Mzil granite of Thomas et al. 2002), with the same mineral assemblage. It is characterized by large megacrysts of K-feldspar, whose size (around 15 mm, up to 35 mm) tends to decrease toward the contacts with the granodiorite, resulting in a gradational relationship between the two rock types. Enclaves of basic rocks are less abundant in monzogranites than in granodiorites.

The Nebdas pluton The Nebdas pluton, about 3 km 1.5 km in size, is intrusive, in its southern part, in the Khzama ophiolitic complex, and is overlain, in the north, by volcanoclastic formations of the Ouarzazate Group. Its central part comprises mainly gabbros, diorites and tonalites grading into each other. Their mineral assemblage comprises clinopyroxene, plagioclase (albitized An5 – 11) and amphibole, with the addition of quartz and biotite in diorites and tonalites, and of K-feldspar in tonalites, and with apatite, zircon and ilmenite as accessory minerals. These rocks are intruded by more leucocratic rocks, of trondhjemitic composition (plagioclase

54 –58%, quartz 38–42%, biotite 2–4%, with accessory titanite and apatite), which become dominant toward the margins of the pluton.

The Ifouachguel pluton The Ifouachguel pluton is a stock of about 5 km2 exposed to the south of the Nebdas plutons. In the south and the north, the massif is intrusive in the volcanoclastic Sarhro Group, and it is intruded in the NW by the pinkish Immorghane granite of Ediacaran age (PIII). The massif is composed dominantly of a melanocratic granodiorite, and locally contains swarms of centimetre-scale dioritic enclaves of variable shape, either angular or rounded. The granodiorite mineral assemblage comprises clinopyroxene, zoned plagioclase, amphibole, biotite, perthitic orthoclase and quartz; the latter two are sometimes intergrown as myrmekites. The accessories are apatite, zircon and ilmenite. The Ifouachguel granodiorite appears to be similar to the Askaoun granodiorite.

The Affela N’ouassif pluton The Affela N’ouassif pluton, located to the east of the N’Kob ophiolites, like a prolongation of the Ida ou Illoun pluton, covers an area of about 12 km 2 km. In north and east the massif is intrusive in the Imghlay series (Bleida Group), and in the south and west it is in faulted contact with or unconformably covered by ignimbrites and rhyolites of Ouarzazate Group or by Neogene phonolites. The massif comprises a large mass of basic rocks, ranging from gabbros to quartz-diorite, surrounded by granodiorite and monzogranite. The basic rocks form a relatively continuous mass, 2 km 0.8 km, in the northeastern part of the pluton, and appear also as metre-scale enclaves disseminated in the granodiorite and monzogranite. Their texture changes from medium grained in the core of the larger mass to fine grained toward its borders and in the enclaves. The contacts with the host granites are either clear-cut or gradational. The mineral assemblage of the basic rocks comprises plagioclase (An48 – 33), amphibole, clinopyroxene, biotite, ilmenite, apatite, zircon and titanite. Quartz and orthoclase are present in the quartz-diorites. The granodiorite is dark-coloured and coarse grained, and shows a gradual enrichment in pinkish K-feldspar toward its contacts with the monzogranite. The mineral assemblage comprises plagioclase (An36 – 26), amphibole, biotite, quartz, orthoclase, ilmenite, apatite and zircon. The largest part of the pluton is occupied by monzogranite, characterized by the presence of K-feldspar megacrysts (around 5 mm in diameter, locally up

NEOPROTEROZOIC SIROUA GRANITOIDS

to 40 mm). Their only mineralogical difference from the granodiorite is the absence of amphibole.

The Imdghar pluton The Imdghar pluton (the Ouafalla granite of Thomas et al. 2002), located to the east of the N’Kob ophiolites, in an eastern prolongation of the Affela N’ouassif pluton, is intrusive in the Sarhro Group in the south, and is unconformably covered by the volcanoclastic Ouarzazate Group in the north. This massif is also characterized by the association of basic rocks, mainly dioritic, and acidic rocks (granodiorite and monzogranite). The dioritic rocks, which range from gabbros to monzodiorites, are distributed mainly in the northeastern part of the pluton, as hectometre-scale masses surrounded by granodiorites, and as small intrusions, of decametre-scale, that cut across the Saghro Group. They are dark-coloured, of variable grain size, and show variable degrees of chloritization. The mineral assemblage comprises plagioclase (An49 – 35), amphibole, biotite, ilmenite, apatite and zircon; in addition, clinopyroxene is found in gabbros, and quartz and orthoclase in monzodiorites. The granodiorites are light grey rocks, with large orthoclase grains, plagioclase, amphibole, biotite and quartz as major minerals, and ilmenite,

269

apatite and zircon as accessories. The monzogranite, which is dominant in the southern part of the pluton, is a coarse-grained pinkish rock, essentially composed of K-feldspar, quartz, plagioclase and biotite, and containing ilmenite, apatite and zircon as accessories.

Mineral composition Pyroxene Clinopyroxene is found in gabbro and gabbrodiorite, where it is occasionally associated with orthopyroxene. In the acidic rocks, clinopyroxene is found only in the Askaoun and Ifouachguel granodiorites, as tiny inclusions (,1 mm diameter) in plagioclase or as relicts in amphibole cores. In the Morimoto et al. (1988) classification diagram (Fig. 3), the clinopyroxene from the less differentiated early gabbros of Askaoun massif is an augite and shows a trend toward increasing Fe and decreasing Ca contents, whereas that from the gabbrodiorite, Askaoun and Ifouachguel granodiorite, and Ida ou Illou gabbrodiorite is an augite –diopside. Significant variations of XFe are observed only in early gabbros (0.15 , XFe , 0.39) and Askaoun gabbrodioritic enclaves (0.16 , XFe , 0.37); in

Wo Diopside

Hedenbergite

Early Gabbro Askaoun diorite Askaoun granodiorite Ifouachguel granodiorite Ida ou Illoun diorite

Augite Pigeonite Enstatite

Ferrosilite

Fs

En

0.08

0.08 Ti

Ti+Cr 0.06

0.06 Tholeitic and calc-alkaline basalts

Alkaline basalts No orogenic basalts

0.04

0.04

0.02

0.02

0 0.4

0.6

0.8

1 Ca+Na

0 0.4

Orogenic basalts

0.6

0.8

Fig. 3. Wollastonite– enstaite– ferrosilite diagram (Morimoto et al. 1988); Ti v. Ca þ Na and Ti þ Cr v. Ca (Leterrier et al. 1982) for pyroxenes from Siroua.

1

Ca

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A. TOUIL ET AL. (Na+K)A

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