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 ge´ologique (Reguibat field trip, Mauritania 2004), Dramane Dembe´le´, now Director of the Direction Nationale de la Ge´ologie 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. Me´not (St Etienne, France) A. Michard (Paris, France) E. Moores (California, USA) P. Morzadec (Rennes, France) H. Ouanaimi (Marrakech, Morocco) Y. Osanai (Fukuoka, Japan)
viii
K. Ouzegane (Algiers, Algeria) C. Passchier (Mainz, Germany) J. J. Peucat (Rennes, France) A. Pouclet (Orle´ans, 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
vii 1
19
33
53
77 111
147
169 203
217 233
249 265
285
303
vi
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 Draˆa 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
329
345
385
409
433 453
467 483
499 509
525
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.
2
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 & Gonc¸uog˘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
3
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
N. ENNIH & J.-P. LIEGEOIS
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
THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT
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
N. ENNIH & J.-P. LIEGEOIS
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
THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT
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
N. ENNIH & J.-P. LIEGEOIS
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 de´collements 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,
THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT
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
N. ENNIH & J.-P. LIEGEOIS
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 (Patin˜o-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
THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT
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 (Patin˜o-Douce 1992); boxes for the three kind of peraluminous granitoids are based to the analyses compiled by Patin˜o-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
THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT
(a) 1000
1000
(b)
1000
Chondrite normalized
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
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
N. ENNIH & J.-P. LIEGEOIS
(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
THE WEST AFRICAN CRATON OF THE MOROCCAN METACRATONIC ANTI-ATLAS BELT
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 Ge´ologiques, the Direction Nationale de la Ge´ologie 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.
15
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
N. ENNIH & J.-P. LIEGEOIS
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 Pen˜o´n 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´troge´ne´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´burne´enne 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. Ge´ologie 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 ne´oprote´rozoı¨que terminal et pale´ozoı¨que de la re´gion de Tata, Anti-Atlas centre-occidental, Maroc: de´formation polyphase´e, 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 Pre´cambrien supe´rieur? Bulletin de la Socie´te´ Ge´ologique de France, 20, 851–857. N AVEZ , J. 1995. De´termination d’e´le´ments en traces dans les roches silicate´es par ICP-MS. Rapport Annuel du De´partement de Ge´ologie et Mine´ralogie 1993–1994 Muse´e 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.
17
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 Ge´ologie, Faculte´ des Sciences, Universite´ Abdou Moumouni, BP: 10662, Niamey, Niger (e-mail:
[email protected] or
[email protected]) 2
De´partement de Ge´osciences, UFR Sciences et Techniques, Universite´ de Franche-Comte´, 16, route de Gray, 25030 Besanc¸on 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
21
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
22
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
REE, Nd, U– Pb, AND BIRIMIAN CRUSTAL GENESIS
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
REE, Nd, U– Pb, AND BIRIMIAN CRUSTAL GENESIS
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
REE, Nd, U– Pb, AND BIRIMIAN CRUSTAL GENESIS
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
REE, Nd, U– Pb, AND BIRIMIAN CRUSTAL GENESIS
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 exe´cution 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 ge´ologiques du Liptako a` 1/ 100 000 et 1/200 000. Direction de la Recherche Ge´ologique et Minie`re, Ministe`re des Mines et de la Ge´ologie, 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 oce´anique birimien pre´coce au Liptako nige´rien (Sirba): ge´ologie, ge´ochronologie et ge´ochimie. 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 juve´nile Birimienne d’Afrique de l’Ouest. Donne´es 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 Ke´ne´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 ge´ologie du sillon de Te´ra, E´tude ge´ochronologique. Unpublished report, Ministry of Mines and Energy of Niger, Niamey, 49–59. K OUAMELAN , A. N. 1996. Ge´ochronologie et ge´ochimie des formations arche´ennes et Prote´rozoı¨ques de la Dorsale de Man en Coˆte d’lvoire. Implications pour la transition Arche´en 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 arche´ennes (3,15 Ga) au sein du magmatisme
31
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 Re´union Annuelle des Sciences de la Terre, Toulouse, France. Socie´te´ Ge´ologique 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 donne´es 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 ge´ochimique de la ceinture birimienne de Diagorou– Darbani, Liptako, Niger occidental (Afrique de l’Ouest). Thesis, University of Franche-Comte´, Besanc¸on. S OUMAILA , A. & K ONATE´ , M. 2005. Caracte´risation de la de´formation dans la ceinture birimienne (pale´oprote´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 oce´anique ou environnement d’arc/bassin arrie`re-arc
32
A. SOUMAILA ET AL.
oce´anique. Comptes Rendus Ge´osciences, 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 ge´odynamique de l’Afrique de l’Ouest entre 2.2 Ga et 2 Ga: le style arche´en des ceintures vertes et des ensembles se´dimentaires Birimiens du nord-est de la Coˆte d’Ivoire. Bulletin de la Socie´te´ Ge´ologique 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 (Taˆc¸araˆt –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 Taˆc¸araˆt –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 Taˆc¸araˆt– Inemmauˆdene Shear Zone (TISZ). (The shear zone is named after two
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 33–52. DOI: 10.1144/SP297.3 0305-8719/08/$15.00 # The Geological Society of London 2008.
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 Taˆc¸araˆt (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.
36
R. M. KEY ET AL.
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
MESOARCHAEAN SHIELDS, NW MAURITANIA
37
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.
38
R. M. KEY ET AL.
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),
MESOARCHAEAN SHIELDS, NW MAURITANIA
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
MESOARCHAEAN SHIELDS, NW MAURITANIA
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
MESOARCHAEAN SHIELDS, NW MAURITANIA
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.
Taˆc¸araˆt– 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 Taˆc¸araˆt 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.
MESOARCHAEAN SHIELDS, NW MAURITANIA
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 Taˆc¸araˆt –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
MESOARCHAEAN SHIELDS, NW MAURITANIA
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
data-point error ellipses are 2σ
data-point error ellipses are 2σ
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).
MESOARCHAEAN SHIELDS, NW MAURITANIA
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) Aoue´oua 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 Taˆc¸araˆt–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 Taˆc¸araˆt–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. Donne´es ge´ochronologiques nouvelles sur l’Arche´en 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. Ge´ologie 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 ante´cambrien de la Mauritanie occidentale. Bulletin de la Socie´te´ Ge´ologique 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 arche´ennes. Exemples naturels et mode`lisation thermome´canique. Me´moires de Ge´osciences, 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 ge´ologique 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 ge´ologique 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 ge´ologiques 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 arche´enne. Exemple de l’Amsaga (R.I. Mauritanie), Dorsale Re´guibat (Craton Ouest Africain). Me´moire de Ge´osciences 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;
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 53–75. DOI: 10.1144/SP297.4 0305-8719/08/$15.00 # The Geological Society of London 2008.
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; Le´corche´ 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 Taoude´ni 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 & Le´corche´ 1989; Le´corche´ 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).
GUELB MOGHREIN MINERALIZATION
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
58
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
GUELB MOGHREIN MINERALIZATION
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).
60
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)
GUELB MOGHREIN MINERALIZATION
(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
GUELB MOGHREIN MINERALIZATION
(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.
GUELB MOGHREIN MINERALIZATION
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
GUELB MOGHREIN MINERALIZATION
(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
GUELB MOGHREIN MINERALIZATION
(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
GUELB MOGHREIN MINERALIZATION
(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
GUELB MOGHREIN MINERALIZATION
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 ge´ologique et mine´ralogique du gisement d’Akjoujt, Mauritanie. PhD thesis, Universite´d’ Orle´ans. 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 donne´es lithostructurales, pe´trographiques, mine´ralogiques et geochimiques sur le gisement de
75
cuivre d’Akjoujt et son environnement ge´ologique (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 Ge´ographie et de l’Ame´nagement 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
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 77–109. DOI: 10.1144/SP297.5 0305-8719/08/$15.00 # The Geological Society of London 2008.
78
M. KAHOUI ET AL.
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 Ge´ologique 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 & Ze´nia 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
79
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).
80
M. KAHOUI ET AL.
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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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).
81
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
82
M. KAHOUI ET AL.
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).
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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.
84
M. KAHOUI ET AL.
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).
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
NNE–SSW limits of sub-basins in the Taoudeni basin, which are filled with Neoproterozoic and Palaeozoic deposits (Villeneuve & Corne´e 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 & Ze´nia 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
86
M. KAHOUI ET AL.
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 & Ze´nia (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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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
88
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.
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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
90
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.
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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).
92
M. KAHOUI ET AL.
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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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
94
M. KAHOUI ET AL.
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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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.
96
M. KAHOUI ET AL.
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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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
98
M. KAHOUI ET AL.
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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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 & Ze´nia 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 & Ze´nia (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,
100
M. KAHOUI ET AL.
(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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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
102
M. KAHOUI ET AL.
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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
Sample: Analysis:
103
104
M. KAHOUI ET AL.
(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
NORTH AFRICAN DIAMONDIFEROUS PROVINCE
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. Re´activation du craton ouest-africain au Panafricain: pale´ocontraintes 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 donne´es 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 oce´anique birimien pre´coce au Liptako nige´rien (Sirba): ge´ologie, ge´ochronologie et ge´ochimie. 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
106
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 sye´nites me´socrates associe´es, stock plutonique sud Tinguicht, zone de jointure Yetti– Eglab (Dorsale Re´guibat, Alge´rie). Bulletin du Service Ge´ologique de l’Alge´rie, 14, 79–95. B ARDET , M. G. 1974. Les gisements kimberlitiques de l’Ouest africain. In: Ge´ologie du diamant. Deuxieme partie: Gisements de diamant d’Afrique. Me´moires du Bureau de Recherche Ge´ologique 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 Ne´oprote´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 Ge´ologique 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 pre´cambrien 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 ge´ologique 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 ge´ologique 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 arche´ennes, exemples naturels et mode´lisation thermome´canique. Me´moires de Ge´osciences 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 ne´oprote´rozoı¨ques sur la bordure nord du craton ouest africain. Bulletin de la Socie´te´ Ge´ologique 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. Ge´ologie 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 ge´ologiques au 1/500 0008 Tindouf–Eglab. Service de la Carte Ge´ologique 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 ge´ologiques pour le diamant exe´cute´es 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´talloge´nique (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´burne´en 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 Ge´ologique 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
107
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 apparente´es de Coˆte-d’Ivoire. Direction des Mines et de la Ge´ologie 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, Be´char. L AMEYRE , J. & L ASSERRE , M. 1967. Etude ge´ochronologique des sye´nites 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. Donne´es ge´ochronologiques sur l’axe pre´cambrien Yetti–Eglab en Alge´rie et en Mauritanie du Nord. Bulletin du Bureau de Recherche Ge´ologique 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 pale´omagne´tiques et pale´ontologiques en faveur d’une position antipodale du craton ouest africain et de la Chine du nord dans le super-continent Rodinia: conse´quences pale´oge´ographiques. Comptes Rendus Ge´oscience, 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´burne´en dans
108
M. KAHOUI ET AL.
les Massifs des E´glab et du Yetti (Dorsale Re´guibat, Alge´rie). Bulletin du Service Ge´ologique 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. De´couverte d’un diatre`me de kimberlite diamantife`re a` Se´gue´la en Coˆte-d’Ivoire. Comptes Rendus de Ge´oscience, 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 Ge´ologique 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 pre´cambienne 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. Donne´es stratigraphiques et pale´omagne´tiques de la re´gion Yetti– Eglab (Sahara occidental alge´rien). Bulletin du Bureau de la Recherche Ge´ologique 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 Ge´ologie. AOF, Dakar.
NORTH AFRICAN DIAMONDIFEROUS PROVINCE S OUGY , J. 1960. Les se´ries pre´cambriennes 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 Guine´e (Afrique de l’Ouest): caracte`res pe´trographiques, mine´ralogiques et ge´ochimiques. 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 pre´cise´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 ge´ologique. 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 Ge´ologique de l’Alge´rie, 8, 259– 272. V ILLEMUR , J. R. 1967. Reconnaissance ge´ologique et structurale du Nord du bassin de Taoudeni. Me´moires du Bureau de Recherche Ge´ologique 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 Ge´ographie et de l’Ame´nagement du Territoire, USTHB, BP 32, Dar el Beida 16111, Alger, Alge´rie (e-mail:
[email protected]) 2
Equipe Ge´obiosphe`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 Ge´osciences 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
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 111–146. DOI: 10.1144/SP297.6 0305-8719/08/$15.00 # The Geological Society of London 2008.
112
A. BENDAOUD ET AL.
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
114
A. BENDAOUD ET AL.
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
TIDJENOUINE METAPELITES EVOLUTION
(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
116
A. BENDAOUD ET AL.
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
TIDJENOUINE METAPELITES EVOLUTION
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
118
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
TIDJENOUINE METAPELITES EVOLUTION
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).
120
A. BENDAOUD ET AL.
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).
TIDJENOUINE METAPELITES EVOLUTION
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
TIDJENOUINE METAPELITES EVOLUTION
Sample: Analysis: Position:
(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)
TIDJENOUINE METAPELITES EVOLUTION
Sample: Analysis: Position:
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.
Sample: Analysis: Position:
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
TIDJENOUINE METAPELITES EVOLUTION
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
TIDJENOUINE METAPELITES EVOLUTION
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.
TIDJENOUINE METAPELITES EVOLUTION
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.
TIDJENOUINE METAPELITES EVOLUTION
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)
TIDJENOUINE METAPELITES EVOLUTION
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
Peak of Metamorphism
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
Peak of Metamorphism
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)
TIDJENOUINE METAPELITES EVOLUTION
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
TIDJENOUINE METAPELITES EVOLUTION
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.
TIDJENOUINE METAPELITES EVOLUTION
(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
A. BENDAOUD ET AL.
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´burne´en 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 synoroge´nique 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 Ge´odynamique, Ge´ologie de l’Inge´nieur et de Plane´tologie, FSTGAT –USTHB, B.P. 32 El Alia, Dar el Beida, 16111 Alger, Alge´rie
3
Equipe Ge´obiosphe`re actuelle et primitive, CNRS IPGP, Universite´ Denis-Diderot (Paris 7), case 89, 4 place Jussieu, 75252 Paris, France 4
Laboratoire de Ge´osciences 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;
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 147–167. DOI: 10.1144/SP297.7 0305-8719/08/$15.00 # The Geological Society of London 2008.
148
Z. ADJERID ET AL.
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).
149
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
150
Z. ADJERID ET AL.
(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.
ULTRAHIGH-I MINERALS, IN OUZZAL
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
152
Z. ADJERID ET AL.
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
ULTRAHIGH-I MINERALS, IN OUZZAL
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.
154
Z. ADJERID ET AL.
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)).
ULTRAHIGH-I MINERALS, IN OUZZAL
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
ULTRAHIGH-I MINERALS, IN OUZZAL
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
ULTRAHIGH-I MINERALS, IN OUZZAL
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
Tek58 (2) Reaction (10)
Tek58 (2) Reaction (10)
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).
160
Z. ADJERID ET AL.
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
ULTRAHIGH-I MINERALS, IN OUZZAL
161
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
162
Z. ADJERID ET AL.
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-
ULTRAHIGH-I MINERALS, IN OUZZAL
163
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).
164
Z. ADJERID ET AL.
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).
ULTRAHIGH-I MINERALS, IN OUZZAL
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.
166
Z. ADJERID ET AL.
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’Arche´en du Hoggar: ge´ochimie et ge´odynamique. Bulletin du Service Ge´ologique 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’Arche´en. 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 Ge´ologique 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
167
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 Sa˜o 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
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 169–201. DOI: 10.1144/SP297.8 0305-8719/08/$15.00 # The Geological Society of London 2008.
170
M. VILLENEUVE
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, Sa˜o 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.
WEST AFRICAN OROGENIC BELTS
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´ Ge´ologique 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,
171
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,
172
M. VILLENEUVE
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).
WEST AFRICAN OROGENIC BELTS
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
173
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.
174
M. VILLENEUVE
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;
WEST AFRICAN OROGENIC BELTS
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
175
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.
176
M. VILLENEUVE
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.
WEST AFRICAN OROGENIC BELTS
177
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;
WEST AFRICAN OROGENIC BELTS
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.
179
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.
180
M. VILLENEUVE
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.
WEST AFRICAN OROGENIC BELTS
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
181
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.
182
M. VILLENEUVE
(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.
WEST AFRICAN OROGENIC BELTS
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
183
(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).
WEST AFRICAN OROGENIC BELTS
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
185
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.
186
M. VILLENEUVE
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,
WEST AFRICAN OROGENIC BELTS
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
187
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
188
M. VILLENEUVE
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.
WEST AFRICAN OROGENIC BELTS
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
189
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
190
M. VILLENEUVE
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
WEST AFRICAN OROGENIC BELTS
191
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).
WEST AFRICAN OROGENIC BELTS
193
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.
194
M. VILLENEUVE
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
WEST AFRICAN OROGENIC BELTS
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
195
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).
196
M. VILLENEUVE
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
WEST AFRICAN OROGENIC BELTS
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 ge´ophysiques en Guine´e (Conakry) occidentale. In: Notice de la carte ge´ologique de la Re´publique de Guine´e. 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.
197
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 ge´ologique du Se´ne´gal oriental et de ses confins guine´o-maliens. Me´moires du BRGM, 40. B ASSOT , J. P. 1969. Aperc¸u sur les formations precambriennes et paleozoiques du Senegal oriental. Bulletin de la Socie´te´ Ge´ologique de France, 7, 160– 169. B ASSOT , J. P. & V ACHETTE , M. 1983. Donne´es nouvelles sur l’aˆge du massif de granitoı¨des du Niokolo–Koba (Se´ne´gal oriental); implications sur l’aˆge du stade pre´coce 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 pre´cambriennes et pale´ozoı¨ques du Se´ne´gal oriental. Bulletin de la Socie´te´ Ge´ologique de France, 5, 401–405. B ECHENNEC , F. 1980. Etude ge´ologique du Nord-Est de la Guine´e 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 arche´en du socle de la feneˆtre de Bou-Naga (Mauritanie); age panafricain des massifs intrusifs alcalins. 11e`me Re´union 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 ge´ologique de la Re´publique de Guine´e au 1/200 000. Service Ge´ologique de Guine´e, Conakry. B RINCKMAN , J., M EINHOLD , K. D., B AH , N. ET AL . 2003. Carte ge´ologique au 1/200 000 des chaıˆnes des Bassarides et les re´gions avoisinantes au Nord-Ouest de la Guine´e. BGR et Service Ge´ologique de Guine´e. Ministe`re des Mines, de la Ge´ologie 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 Ge´oscience, 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 ge´ologique de la chaıˆne des Mauritanides entre le paralle`le de Moudjeria et le fleuve Se´ne´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 Se´ne´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 photoge´ologique de la chaıˆne du Dhlou (Zemmour– Mauritanie septentrionale). Laboratoire de Ge´ologie, 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 ge´ologique du Mali occidental au 1/200 000. Feuilles Kankossa, Kayes, Kossanto. Re´publique du Mali, Direction Nationale de la Ge´ologie 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´ Ge´ologique 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 Se´ne´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 de´couverte de Silurien supe´rieur au sommet de la se´rie de Youkounkoun au Se´ne´gal oriental: implications pale´oge´ographiques et structurales. Bulletin de la Socie´te´ de Ge´ologie 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 oce´aniques au sein de la chaıˆne panafricaine des Mauritanides dans la re´gion des Bassaris (Guine´e –Se´ne´gal). Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 299, 65–70. G EVIN , P. 1960. Etudes et reconnaissances ge´ologiques 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´ Ge´olgique 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 ge´ologie du Se´ne´gal et des re´gions voisines. Bulletin de la Socie´te´ Ge´ologique 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. Re´activation de la zone frontale d’une chaıˆne polyphase´e: 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. De´couverte d’e´clogites 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 Se´ne´gal et de la Mauritanie. The`se d’Etat, Universite´ Aix– Marseille III, Marseille. L ILLE , R. 1967. Etude ge´ologique 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 ge´ologique et metalloge´nique 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 Se´ne´gal a` la Sierra Leone: interpre´tation ge´ophysique 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 donne´es lithostructurales, pe´trographiques, mine´ralogiques et ge´ochimiques du gisement de cuivre d’Akjoujt et son environnement ge´ologique (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 oce´anique limite´e 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 ge´ologique 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 Se´ne´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 ge´ologique du Maroc au 1/50 000, Feuille Award. Notes et Me´moires du Service Ge´ologique 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 franc¸aise. Bulletin de la Socie´te´ Ge´ologique de France, 18, 589–628. S ELIVERSTOV , Y. P. 1970. Notice explicative de la carte ge´ologique de la Re´publique de Guine´e 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 ge´ologique des guelbs Bou-Leriah (re´gion d’Aoucert, Sahara espagnol). Bulletin de la Socie´te´ Ge´ologique de France, 4, 436–445. S OUGY , J. 1964. Les formations pale´ozoı¨ques du Zemmour noir (Mauritanie septentrionale). Etude stratigraphique, pe´trographique et pale´ontologique. 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 pre´cambrien et sa couverture infracambrienne et pale´ozoı¨que), Afrique de l’Ouest. Bulletin de la Socie´te´ Ge´ologique 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 ge´ologique de la Re´publique de Guine´e au 1/200 000, feuille C-28– 12 (Sierumba). Technoexport, Conakry. T ORCHINE , N. S. 1976a. Notice explicative de la carte ge´ologique de la Re´publique de Guine´e 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 ge´ologique de la Re´publique de Guine´e au 1/ 200 000, feuilles D 28– 35 et D 28–29 (Youkounkoun). Technoexport, Conakry. T ROMPETTE , R. 1973. Le Pre´cambrien supe´rieur et le Pale´ozoı¨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 ge´ologique du sud du Fouta Djalon (Guine´e). Comptes Rendus Sommaire de la Socie´te´ Ge´ologique de France, 2, 55–59. V ILLENEUVE , M. 1982. Sche´ma lithostratigraphique des Mauritanides au Sud du Se´ne´gal et au Nord de la Guine´e d’apre`s les donne´es actuelles. Bulletin de la Socie´te´ Ge´ologique de France, 7, 249– 254. V ILLENEUVE , M. 1984. Etude ge´ologique 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 (Guine´e–Se´ne´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 (Guine´e–Se´ne´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 pale´ozoı¨que de Guine´e (Afrique de l’Ouest). Bulletin de la Socie´te´ Ge´ologique 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, Se´ne´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 Ge´ologie 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.
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 203–216. DOI: 10.1144/SP297.9 0305-8719/08/$15.00 # The Geological Society of London 2008.
204
R. CABY ET AL.
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).
205
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
206
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
SUBDUCTION METAMORPHISM, EASTERN MALI
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 (Kie´nast 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.
208
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.
SUBDUCTION METAMORPHISM, EASTERN MALI
(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
R. CABY ET AL.
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.
SUBDUCTION METAMORPHISM, EASTERN MALI
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
211
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
212
R. CABY ET AL.
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, Me´not & 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
SUBDUCTION METAMORPHISM, EASTERN MALI
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 de´collement 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
213
(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 (Me´not & 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).
SUBDUCTION METAMORPHISM, EASTERN MALI
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; Kie´nast 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 ge´ologique de l’Adrar des Iforas et du Gourma oriental’ supported by the Banque Europe´enne d’Investissement (BEI). We are grateful to the Direction Nationale de la Ge´ologie et des Mines (Bamako) and to GEOTER for logistical support and facilities. We thank C. Passchier and R. P. Me´not 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
215
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: Ge´ome´trie d’une suture pan-africaine. Bulletin de la Socie´te´ Ge´ologique 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 ge´ologique du Gourma oriental au 1/200 000 Ansongo– Amalaoulaou. Direction Nationale des Mines et de la Ge´ologie, Bamako, Mali. C ABY , R. 1979. Les nappes pre´cambriennes du Gourma dans la chaıˆne pan-africaine du Mali. Comparaison avec les Alpes occidentales. Revue de Ge´ologie Dynamique et de Ge´ographie 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 synoroge´nique 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 ge´ologique de l’Adrar des Iforas au 1/200 000. Direction Nationale de la Ge´ologie 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’E´nergie 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 ge´ochronologie de roches cristallophyliennes du bassin du Gourma (Mali): conse´quences ge´odynamiques. The`se 3e cycle, Universite´des Sciences, Montpellier. D E L A B OISSE , H. 1981. Sur le me´tamorphisme du micaschiste e´clogitique de Takamba (Mali) et ses conse´quences pale´oge´odynamiques au Pre´cambrien supe´rieur. Comptes Rendus Sommaires de la Socie´te´ Ge´ologique 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´ Ge´ologique 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 Pale´oprote´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 Pre´cambrien spe´rieur? Bulletin de la Socie´te´ Ge´ologique 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 ge´ologique du Gourma. Me´moires du BRGM, 53. R EICHELT , R. 1972. Ge´ologie du Gourma (Afrique occidentale). Un ‘seuil’ et un bassin du Pre´cambrien 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
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 217–231. DOI: 10.1144/SP297.10 0305-8719/08/$15.00 # The Geological Society of London 2008.
218
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
219
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).
220
K. ATTOH & P. M. NUDE
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
CARBONATITES IN DAHOMEYIDE SUTURE ZONE
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
221
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
222
K. ATTOH & P. M. NUDE
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
CARBONATITES IN DAHOMEYIDE SUTURE ZONE
223
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
224
K. ATTOH & P. M. NUDE
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. Me´not & 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
CARBONATITES IN DAHOMEYIDE SUTURE ZONE
225
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
226
K. ATTOH & P. M. NUDE
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.
CARBONATITES IN DAHOMEYIDE SUTURE ZONE
227
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).
228
K. ATTOH & P. M. NUDE
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–Sa˜o 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
CARBONATITES IN DAHOMEYIDE SUTURE ZONE
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
230
K. ATTOH & P. M. NUDE
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
231
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 ge´ologie, Universite´ Hassan II—Aı¨n Chock, Route d’El Jadida, B.P. 5366 Maaˆrif, Casablanca, Morocco
3
Laboratoire de Ge´ologie, 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
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 233–247. DOI: 10.1144/SP297.11 0305-8719/08/$15.00 # The Geological Society of London 2008.
234
R. BOUSQUET ET AL.
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.
235
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
236
R. BOUSQUET ET AL.
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.
ME´LANGES AND OPHIOLITES, MOROCCO
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.
238
R. BOUSQUET ET AL.
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
No Garnet
Pressure [kbar]
6
4
No Na-Amphibole
No Albite
0.1
0.3 0.4 0.5 0.6 0.7
No Na-Amphibole No Na-Amphibole
4
Riebeckite end-member 2
2
a k Do
300
0.5
6
8 0.3 0.2 0.1
No Na-Amphibole No Na-Amphibole
Pressure [kbar]
0.4
No Albite
600
a k Do
Alb-Gt-Chl Ep-Gln 0.5
0.4 0.3
6
0.6 0.2
No Na-Amphibole No Na-Amphibole
4
4
Fe-Glaucophane end-member 2 300
400 500 Temperature [°C]
(d) 10 No Garnet
Alb-Gt-Chl Ep-Gln
Th e r i
No Albite
600
i m no
8
400 500 Temperature [°C]
i m no
No Garnet
300
Pressure [kbar]
0.2
6
Stability field
(c) 10
a k Do
Alb-Gt-Chl Ep-Gln
Th e r i
Pressure [kbar]
8
Alb-Gt-Chl Ep-Gln
No Na-Amphibole
(b) 10
i m no
8
a k Do
i m no
No Albite
239
Th e r i
(a) 10
No Garnet
ME´LANGES AND OPHIOLITES, MOROCCO
400 500 Temperature [°C]
Glaucophane end-member 600
2 300
400 500 Temperature [°C]
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).
240
R. BOUSQUET ET AL.
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
ME´LANGES AND OPHIOLITES, MOROCCO
241
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),
242
R. BOUSQUET ET AL.
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.
ME´LANGES AND OPHIOLITES, MOROCCO
243
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
244
R. BOUSQUET ET AL.
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. De´couverte d’un syste`me hydrothermal oce´anique fossile dans l’ophiolite ante´cambrienne 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, Eube´e): implications ge´odynamiques. 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 pre´cambriennes et gıˆtes arse´nie´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
e´clogitique 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 Ge´ologie, B.P. 549, Marrakech 40 000, Morocco (e-mail:
[email protected]) 2
Universite´ du Que´bec a` Montre´al, De´partement des Sciences de la Terre et de l’Atmosphe`re, CP 8888 Centre Ville, Montre´al, 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
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 249–264. DOI: 10.1144/SP297.12 0305-8719/08/$15.00 # The Geological Society of London 2008.
250
A. BELKABIR ET AL.
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
251
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).
252
A. BELKABIR ET AL.
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
GOLD MINERALIZATION IN BLEIDA OPHIOLITE
253
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.
A. BELKABIR ET AL.
Average 4
254
Facies:
GOLD MINERALIZATION IN BLEIDA OPHIOLITE
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).
255
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.
256
A. BELKABIR ET AL.
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
GOLD MINERALIZATION IN BLEIDA OPHIOLITE
257
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
258
A. BELKABIR ET AL.
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.
GOLD MINERALIZATION IN BLEIDA OPHIOLITE
259
(a)
Banded volcanic rocks Mafic dyke Breccia lithofacies Fe-rich lithofacies Argillic lithofacies Cu-rich chert
Au (g/t)
10
1.0
0.1
0.01 0
20
40
60
80
Sample location in metres
(b)
Au-quartz veins Banded volcanic rocks
100
Cu-rich chert Fe-rich lithofacies
Au (g/t)
10 1 0.1 0.01 0.001 0
50
100
150
200
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
260
A. BELKABIR ET AL.
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
GOLD MINERALIZATION IN BLEIDA OPHIOLITE
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
261
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.
262
A. BELKABIR ET AL.
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 Conceic¸a˜o mine; Olivo et al. 1995) located in the Itabira district, southern Sa˜o 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 Conceic¸a˜o 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
GOLD MINERALIZATION IN BLEIDA OPHIOLITE
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
263
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 pale´ocirculations fluides dans l’indice d’or de Bleida (Anti-Atlas, Maroc). Comptes Rendus Ge´osciences, 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´ve´nements e´burne´ens et panafricains dans l’Anti-Atlas marocain: Apport de la me´thode ge´ochimique Rb/Sr. Notes et Me´moires du Service Ge´ologique du Maroc, 313. C ISSE´ , A. 1989. Ge´ome´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. Ge´ochimie isotopique du strontium des milieux se´dimentaires. Application a` la ge´ochronologie 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´ Ge´ologique de France, (7), XXI, 495–499. D UCROT , J. & L ANCELOT , J. R. 1977. Proble`me de la limite Pre´cambrien– 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 polyphase´e dans la boutonnie`re pre´cambrienne de Bou Azzer (Anti-Atlas central, Maroc): apport de l’imagerie spatiale Landsat-MSS et de l’analyse structurale de terrain. Sciences Ge´ologiques, 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
264
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 Pre´cambrien III du Haut Atlas et le proble`me de la limite Pre´cambrien–Cambrien (abstract). Sciences de la Terre, 2nd Annual Meeting, Nancy, 230. L EBLANC , M. 1981. Ophiolites pre´cambriennes et gıˆtes arse´nie´s de cobalt (Bou-Azzer, Maroc). Notes et Me´moires du Service Ge´ologique 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 ge´odynamique du domaine pan-africain (Pre´cambrien terminal) de l’Anti-Atlas (Maroc) a` partir de donne´es ge´ologiques et ge´ochronologiques. 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 & Corne´e 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.
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 265–283. DOI: 10.1144/SP297.13 0305-8719/08/$15.00 # The Geological Society of London 2008.
266
A. TOUIL ET AL.
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 & Corne´e 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.
268
A. TOUIL ET AL.
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
270
A. TOUIL ET AL. (Na+K)A<0.5 ; Ti<0.5 Tremolite
Act Hbl
Mg/(Mg+Fe)
0.8
(Na+K)A>0.5 ; Ti<0.5 1
Tr Hbl Magnesio-Hbl Tsch Hbl
0.6 Actinolite
0.4 FerroActinolite
0.2
FeAct Hbl
Ferro-Hbl
Silicic Edenite
0.6
0
Mga Has Hbl
Magnesian Hastingsite
Has Hbl
Hastingsite
Fe Ed Hbl
0.4 FerroEdenite
0.2
Early gabbro Ifouachguel granodiorite Ait Nebdas gabbro-diorite Ait Nebdas tonalite Askaoun diorite Askaoun granodiorite Affela N’ouassif diorite Ida ou Illoun diorite Imdghar granodiorite Affela N' ouassif granodiorite and granite Ida ou Illoun granodiorite and granite
MagnesioHastingsite
Ed Hbl
Silicic Ferro-Edenite
FeFerroTsch Tschermakite Hbl
Mg Has Hbl
Edenite
0.8 Tschermakite
Mg/(Mg+Fe)
1
0 8
7.5
7
6.5
6
5.5
8
7.5
7
6.5
6
5.5
Fig. 4. Leake et al. (1997) classification diagram for amphiboles from Siroua.
the other samples, XFe values are around 0.28. Clinopyroxene from the early gabbro is characterized by relatively high contents of Al and Ti, increasing with differentiation. Titanium increase is consistent with the late crystallization of ilmenite suggested by the textural relationships. The relatively high Ti content of early gabbros is illustrated on the diagrams of Leterrier et al. (1982) for the discrimination of magmatic series (Fig. 3). On a Ti v. Ca þ Na diagram, all clinopyroxene compositions fall within the tholeiitic–calc-alkaline field, whereas on a Ti þ Cr v. Ca diagram, clinopyroxene from the early gabbros falls within the nonorogenic basalt field, whereas that from other rocks falls in the orogenic field.
Amphibole Amphibole is present in nearly all rock types, as single crystals in the matrix or as inclusions, or in association with clinopyroxene. The crystal cores have often a more brownish colour, especially in the Ida ou Illoun and Ifouchguel massifs. All analyses correspond to calcic amphibole in the Leake et al. (1997) classification scheme. Based on variations in alkali contents and XMg values, three compositional groups are distinguished (Fig. 4). Group 1 comprises amphiboles from the basic rocks of all plutons and from the Nebdas, Askaoun and Ifouachguel granodiorites. This group is characterized by relatively low alkali contents (0.03 , (Na þ K)A , 0.5 a.p.f.u.), low titanium (0.06 , Ti , 0.39 a.p.f.u.), and XMg between 0.5 and 0.9. The compositions vary from tschermakitic hornblende, forming the brown cores, through magnesio-hornblende, to actinolite-rich hornblende on rims. Group 2 comprises amphiboles from the Ida ou Illoun and Affela N’ouassif granodiorites. Alkalis (0.03 , (Na þ K)A , 0.5 a.p.f.u.) and titanium contents (0.06 , Ti , 0.39 a.p.f.u.) are lower than
in Group 1, and XMg is relatively low (0.2 , XMg , 0.38). Group 2 amphiboles are classified as ferro-hornblendes. Group 3 comprises a few analyses of amphiboles from granodiorite and granite from Ida ou Illoun and Affela N’ouassifat. They are distinguished from Group 2 amphiboles by a high alkali content (0.51 , (Na þ K)A , 0.68 a.p.f.u.) and are classified as edenitic ferro-hornblendes.
Biotite Distribution and abundance of biotite is highly variable between and within the various rock types. Biotite is not found in the most basic rocks (gabbros and diorites sensu stricto) and becomes relatively abundant in more differentiated basic rocks. It becomes a significant component in granites and granodiorites, but it is still less abundant than amphibole in the latter. Based on XFe values, two composition groups are distinguished. Group 1 comprises biotites with lower XFe (0.4 , XFe , 0.5), from the Ida ou Illoun, Askaoun, Imdghar diorites and Askaoun granodiorites. Group 2 comprises biotites with higher XFe (0.7 , XFe , 0.9), from the Ida ou Illoun and Afella N’ouassif granodiorites and granites. In the diagram used by Abdel-Rahman (1994) for discrimination of magmatic series based on biotite composition, Group 1 biotites plot in the calcalkaline field, those from Group 2 plot in the alkaline field (Fig. 5). The same affinities are found using the diagrams (not reproduced here) of Nachit et al. (1985), where Groups 1 and 2 plot respectively in the calc-alkaline and subalkaline domains. To summarize the results of the mineral chemistry, three main groups are distinguished among the various rock types and plutons. One group corresponds to early gabbros and diorites, found in Askaoun, where the clinopyroxene compositions,
NEOPROTEROZOIC SIROUA GRANITOIDS 18
Peraluminous 16
Al2O3
Calc-alkaline
14
Alkaline
12
10 0
3
6
9
12
15
MgO
Fig. 5. Abdel-Rahman (1994) Al2O3 v. MgO plot of Siroua biotites. Symbols as for Figure 4.
namely the relatively high Ti contents, suggest crystallization from tholeiitic magmas. The second group comprises gabbros, diorites and tonalites from Nebdas, granodiorites and associated gabbro–dioritic enclaves from Askaoun and Ifouachguel, and basic rocks (gabbro–diorites) from Ida ou Illoun, Affela N’ouassif and Imdghar. The composition of clinopyroxene, amphibole and biotite characterizes the calc-alkaline affinity of this group. The third group comprises granodiorites and granites from Ida ou Illoun, Affela N’ouassif and Imdghar, where the relatively high Fe contents of amphiboles and biotites point to a sub-alkaline affinity.
Whole-rock chemistry Table 1 shows representative bulk compositions of samples from the various plutons. Major, trace and rare earth element abundances have been determined by plasma emission spectrometry (complete dissolution of about 250 g of rock powder in three acids (HNO3, HCl, and HF at 1008C) and X-ray fluorescence, using a X Philips PW 1404 system operating at 40 kV and 70 mA. Further analytical details have been given by Touil (1999) and El Khanchaoui (2001). The early gabbros and diorites from Askaoun have relatively basic compositions (SiO2 between 47.80 and 52.92%) and a limited range of variation; samples from other plutons are generally more acidic and define a general trend of increasing SiO2 and alkali contents and decreasing contents of TIO2, Fe2Ot3, Al2O3, MgO and CaO. Most samples show a metaluminous character (Clarke 1981), with A/CNK ratio [molar ratio Al2O3/ (CaO þ Na2O þ K2O)] ,1; this is consistent with
271
the relatively low Al contents found in biotites (Al2O3 , 14.7%). Based on variations in major and trace element contents, five main groups of rocks are distinguished (Fig. 6a–c). Group 1 comprises the early gabbros and diorites from Askaoun, which are characterized by low potassium (K2O 0.19–0.68%) and high TiO2, Fe2Ot3, Al2O3 and P2O5 (Fe2O3 varies from 14 to 15.24%; P2O5 from 6.25 to 5.12% and TiO2 from 2.70 to 3.06%), as well as high Y, Zr and Nb. Judging from the low ignition loss values, the chemical character of these rocks is considered as primary, and the low alkali contents, together with high concentrations of iron and titanium, point to a tholeiitic affinity, which is also indicated by the clinopyroxene composition. Group 2, the Ait Nebdas trondhjemites, is characterized by high Na2O, Al2O3 and SiO2 contents and low K2O. The contents of transition elements (Ni, Cr, V and Sc), incompatible elements (Zr, Th, Nb, Ce, Y, Rb) and Ba is low, but Sr is relatively abundant. These features are comparable with those of continental trondhjemites (Arth 1979; Barker 1979) and are in contrast to those of oceanic plagiogranites (Coleman & Peterman 1975). Group 3 comprises gabbros and diorites from Ida ou Illoun, Affela N’ouassif and Imdghar, and granodiorites and associated gabbro–dioritic enclaves from Askaoun and Ifouachguel. Relatively low contents of iron (Fe2O3 , 8.74%) and titanium (TiO2 , 1.6%), and high potassium levels (K2O varying from 1.42 to 2.86%, for SiO2 between 51.98 and 61%) are indicative of a high-K calc-alkaline series. This chemical affinity is consistent with the results of typological studies of zircons from the Askaoun (Regragui 1997) and Ifouachguel (El Khanchaoui et al. 2001) plutons, and from the Ida ou Illoun diorites (Jouider 1997), which have shown preferential developments of the [100] prism over the [110], and of the [101] pyramid over the [211]. In the IA –IT diagram (where A corresponds to the average development of pyramids (Agpaicity, Alkalinity and Acidity) and T corresponds to the average development of prisms (Temperature)) (Pupin 1980), these zircon populations plot in the field of potassic calc-alkaline rocks formed from a dry and hot magma with a significant mantle component. Crystallization temperatures are estimated at 850 + 50 8C and 825 + 50 8C for the Askaoun and Ifouachguel plutons, respectively. Similar temperatures (812 + 60 8C) were obtained on Askaoun gabbro–diorites by Touil (1999) on orthopyroxene–clinopyroxene pairs, using the Wood & Banno (1973) thermometer. The Askaoun and Ifouachguel granodiorites show limited variation ranges that can be accounted for by processes of fractionation, involving clinopyroxene, of hybrid
272
Table 1. Representative major and trace element analyses of Siroua granitoids Sample
Nb11*
Nb30*
Nb5b*
IF1*
7B3
8
J127
47.89 2.93 12.96 15.08 0.33 6.25 7.12 3.52 0.20 0.33 2.60 99.21 64 2 170 42 203 14 4 4 20 141 59 371 64 38 56 17.5 43.7 na 36.7 8.8 2.5 na na 8.4 na na na 4.4 na 2.69
52.48 0.87 12.34 11.58 0.26 7.95 7.68 2.87 0.88 0.10 2.40 99.41 320 23 341 16 58 3 1 88 na 267 100 264 321 32 38 6.1 15.1 2.8 13.4 3.8 1.7 3.9 0.9 5.2 1.1 3.2 0.4 2.9 0.4 1.43
59.40 0.85 19.00 4.93 0.12 2.53 3.86 5.31 1.80 0.21 2.00 100.01 300 61 555 13 54 2 1 6 na 92 10 103 13 8 28 5.6 10.8 2.0 9.5 2.7 1.1 2.8 0.6 3.2 0.6 1.6 0.2 1.4 0.2 2.67
73.71 0.13 14.98 1.32 0.04 0.42 1.64 5.26 1.16 0.01 1.2 99.87 234 31 465 6 80 2 1 14 na 56 2 8 2 2 78 11.8 25.1 3.6 14.0 2.6 0.8 2.0 0.3 1.4 0.3 0.8 0.1 0.7 0.1 10.66
64.42 0.63 15.62 4.30 0.18 1.71 3.57 4.13 2.99 0.15 1.6 99.30 786 100 353 27 135 8 8 33 na 134 5 81 6 12 45 25.2 49.2 7.2 25.5 5.0 1.3 1.0 1.0 5.4 1.1 3.1 0.5 3.2 0.4 5.30
54.47 0.88 16.32 7.75 0.21 4.19 5.82 4.18 2.85 0.16 1.51 98.34 600 91 374 23 97 6 6 18 17 113 24 174 50 25 27 18.7 43.4 na 23.3 4.7 1.2 na na 4.0 na na na 2.6 na 4.88
61.17 0.67 15.52 5.50 0.20 2.75 3.72 3.29 3.43 0.14 2.02 98.41 890 106 342 23 133 9 13 32 17 142 16 99 20 14 18 20.9 49.3 na 24.4 4.3 1.1 na na 3.7 na na na 2.4 na 5.73
52.70 1.01 16.26 8.53 0.21 6.72 8.04 3.00 1.27 0.13 1.46 99.33 247 43 240 19 80 5 5 5 16 60 91 170 193 24 39 23.3 51.4 na 29.4 5.0 1.4 na na 4.2 na na na 3.0 na 5.28
A. TOUIL ET AL.
SiO2 TiO2 Al2O3 Fe2OT3 MnO MgO CaO Na2O K2 O P2O5 LOI Total Ba Rb Sr Y Zr Nb Th Pb Ga Zn Ni V Cr Sc Co La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu (La/Yb)N
114
Table 1. Continued J145
120
TT5*
TT18*
TT16*
NK64*
NK107*
NK7*
SiO2 TiO2 Al2O3 Fe2OT3 MnO MgO CaO Na2O K2O P2O5 LOI Total Ba Rb Sr Y Zr Nb Th Pb Ga Zn Ni V Cr Sc Co La Ce Pr Nd Sm Eu Gd Tb Dy Ho
65.09 0.74 14.36 7.02 0.14 0.93 3.15 3.61 3.63 0.16 0.60 99.43 1064 127 147 67 350 12 24 14 22 93 6 37 8 24 19 98.6 211.0 na 102.0 21.8 2.2 na na 15.6 na
71.75 0.36 13.46 3.16 0.07 0.44 1.88 3.27 4.46 0.08 0.63 99.55 855 169 104 32 149 6 14 21 18 45 5 18 10 8 6 41.7 80.8 na 36.9 7.1 1.0 na na 6.5 na
52.76 1.37 16.47 10.31 0.33 3.85 6.25 4.23 1.63 0.22 2.30 99.72 320 65 230 29 200 10 3 22 na 140 8 160 84 26 34 40.8 71.3 10.7 41.6 7.9 2.4 7.0 1.3 7.1 1.4
66.81 0.73 14.27 6.91 0.06 0.89 0.58 3.87 3.60 0.12 2.00 99.84 1107 120 92 34 242 11 10 17 na 38 2 40 7 22 98 69.5 113.7 19.1 55.3 9.4 2.0 8.4 1.5 8.2 1.7
72.70 0.25 12.57 3.10 0.02 0.45 0.34 3.10 5.65 0.08 1.20 99.46 434 120 50 30 111 7 8 23 na 18 3 3 2 4 125 42.0 82.2 10.2 36.1 6.5 0.9 5.4 1.0 5.9 1.2
48.62 0.94 14.66 13.60 0.40 7.56 7.86 2.65 1.04 0.22 1.77 99.32 405 32 193 20 98 7 3 na na 143 78 319 400 49 62 10.0 21.2 0.0 9.9 3.4 1.0 3.5 0.0 3.6 0.0
66.54 0.45 14.68 5.20 0.05 1.92 2.25 4.89 3.48 0.13 1.79 101.38 1182 94 136 36 202 7 15 na na 103 10 142 84 13 84 na na na na na na na na na na
70.57 0.48 13.74 3.20 0.04 1.49 0.59 4.67 3.85 0.10 1.50 100.23 797 99 95 31 280 8 12 8 na 41 6 24 6 9 72 55.6 109.0 12.8 43.7 7.6 1.2 6.8 1.2 7.2 1.4
273
(Continued)
NEOPROTEROZOIC SIROUA GRANITOIDS
Sample
*na, not analysed; 114, Early Gabbro; Nb11*, Ait Nebdas diorite; Nb30*, Ait Nebdas tonalite; Nb5b*, Ait Nebdas trondhje´mite; IF1*, Ifouachguel granodiorite; 7B3, Askaoun diorite; 8, Askaoun granodiorite; J127, Ida ou Illoun diorite; J145, Ida ou Illoun granodiorite; 120, Ida ou Illoun granite; TT5*, Affela N’ouassif diorite; TT18*, Affela N’ouassif granodiorite; TT16*, Affela N’ouassif granite; NK64*, Imdghar gabbro and diorite; NK107*, Imdghar granodiorite; NK7*, Imdghar granite. Samples from El Khanchaoui (2001).
na na na na na 2.0 0.0 1.8 0.3 3.66 3.0 0.3 2.9 0.4 9.74 3.9 0.4 3.4 0.4 8.00 na na 80 na 8.27 Er Tm Yb Lu (La/Yb)N
na na 3.7 na 7.48
4.6 0.5 4.2 0.5 11.04
NK107* NK64* TT16* TT18* TT5* 120 J145 Sample
Table 1. Continued
4.1 0.5 3.6 0.5 10.27
A. TOUIL ET AL.
NK7*
274
magmas formed by mixing between crustal melts and basic components represented by the associated gabbro–diorite enclaves (Touil 1999). Group 4 comprises gabbros, diorites and tonalites from Nebdas. These rocks are, from a mineralogical point of view, comparable with rocks from Group 2, but, considering their lower contents of incompatible elements (Zr, Y, Nb, Th, Ce), transition elements (Ni, Cr) and alkalis (K, Rb), they are classified as low-K calc-alkaline rocks. Major and trace element variations are consistent with a genetic link, through fractional crystallization processes, between the gabbro–diorites and the tonalites. This link is also suggested by the REE patterns, which show a progressive increase of the heavy/light REE ratio (HREE/LREE) from the gabbro–diorites ((La/Yb)N ¼ 2.42– 3.29) to the tonalites ((La/Yb)N ¼ 2.67). Group 5 corresponds to granodiorites and monzogranites from the Ida ou Illoun, Affela N’ouassif and Imdghar plutons. The evolution of bulk composition from the granodiorites to the monzogranites is consistent with a process of fractional crystallization involving amphibole, biotite and feldspars. The evolution of REE patterns is characterized by a decrease of both LREE and HREE levels and a strongly deepening negative Eu anomaly. The rocks of Group 5 are characterized by high concentrations in K2O (from 3.5 to 7%, for SiO2 from 67 to 78%) and enrichments in Fe, Th, Zr, Nb, Ce and Y. This indicates that they cannot be derived from a typical calc-alkaline magma. They have a definite alkaline affinity, as shown also by the low-Al, high-Fe biotite compositions. Typological studies on zircons from the Ida ou Illoun granodiorites and monzogranites (Jouider 1997) have shown the development of [100] prismatic types and [101] pyramidal types and a high frequency of P3 subtypes, corresponding to a domain overlapping the typological fields of alkaline and subalkaline series. Temperatures estimated from these studies are around 850 8C.
Magma affinities and geodynamic setting Basic and intermediate rocks The presence of two main magmatic trends in the basic rocks of the Siroua district, with, on one hand, a tholeiitic group represented by the early gabbros and diorites of Askaoun, and, on the other hand, a calc-alkaline trend represented by the basic and intermediate rocks of the other plutons (Nebdas, Ifouchguel, Askaoun, Ida ou Illoun, Affela N’ouassif Imdghar), is corroborated by several classical bulk chemistry plots, including AFM projection (Irvine & Baragar 1971),
NEOPROTEROZOIC SIROUA GRANITOIDS 8
275
8 K 2O
Na2O
6
6
shoshonitic high-K
4
4
2
2
medium-K low-K 0
0 45
55
75 SiO2
65
10
45
55
65
75 SiO2
16 CaO
Fe2O3
8 12
6 8 4
4 2
0
0 45
55
65
45
75 SiO2
20
55
65
75 SiO2
4 TiO2
Al2O3 18
16 2 14
12
0
10 45
55
65
75 SiO2
45
55
65
75 SiO2
Ifouachguel granodiorite Ait Nebdas trondhjémite Ait Nebdas tonalite Early gabbros Ait Nebdas gabbro and diorite Affela N'ouassif diorite Askaoun diorite Ida ou Illoun gabbro and diorite Ida ou Illoun granodiorite and granite Askaoun granodiorite Imdghar granodiorite and granite Imdghar gabbro and diorite Affela N'ouassif granodiorite and granite
Fig. 6. Major (wt%) and trace (ppm) elements v. SiO2 (wt%) in the Siroua granitoids. K2O v. SiO2 diagram; separating lines are from Rickwood (1989).
276
A. TOUIL ET AL.
4
FeO*
TiO2
3
Tholeiitic Tholeiitic
Early gabbros Ait Nebdas gabbro and diorite Askaoun diorite Ida ou Illoun gabbro and diorite Affela N’ouassif diorite Imdghar gabbro and diorite
2
Calc-alkaline 1 Calc-alkaline FeO*/MgO
Na2O+K2O
0 0
2
1
3
MgO
4
Ti/100
1 Nb/Y
0.6
A lk al in e
0.8
Tholeiitic
A
D 0.4
B C
0.2
Zr
Zr/P2O5
0 0
0.05
0.1
3*Y
0.2
0.15 T
Fig. 7. Miyashiro (1974) TiO2 v. FeO /MgO, Irive & Baragar (1971) Na2O þ K2O– FeOT –MgO, Winchester & Floyd (1977) Nb/Y v. Zr/P2O5 and Pearce & Cann (1973) Zr–Ti/100–3*Y (A, low K-tholeiites; B, ocean-floor basalts; C, calc-alkaline basalts; D, within-plate basalts) discrimination diagrams for Siroua basic rocks.
TiO2 v. FeO*/MgO (Miyashiro 1974), and Nb/Y v. Zr/P2O5 (Winchester & Floyd 1977) (Fig. 7). In the geodynamic projections (Zr –Ti/100 –3Y and Zr–Ti/100– Sr/2; Pearce & Cann 1973), the gabbro–diorite from Nebdas and the early gabbros plot in the field of anorogenic and oceanic basalts, whereas the other compositions plot in the calc-alkaline domain (Fig. 7). On chondrite-normalized REE diagrams and extended normalization diagrams (Thompson et al. 1982), the early gabbros have spidergrams similar to those of back-arc tholeiites or enriched mid-ocean ridge basalt (MORB) (Figs 8 and 9), but they differ from the latter in showing negative anomalies in Nb and Sr, strong enrichment in large ion lithophile elements (LILE), Th, La, Ce and Zr, and a negative slope. The gabbro–diorites from Ait Nebdas, whose weak potassic character has been already mentioned, give spidergrams that differ from those of the other groups of calc-alkaline rocks: they show relatively low levels in most incompatible elements and a less pronounced niobium negative anomaly. On the
other hand, the high-K calc-alkaline rocks from Ifouchguel, Askaoun, Ida ou Illoun, Affela N’ouassif and Imdghar are represented by a series of similar spidergrams, parallel to each other, showing higher enrichment in LILE and LREE than those of Nebdas. Such patterns are similar to those of the post-collisional potassic arc magmas defined by Muller et al. (1992). A crustal contamination of the parental magmas of all basic rocks is suggested by the systematic presence of negative anomalies in niobium on spidergrams, and by their high Th/Nb (.3) and Ba/Zr ratios (.1.2). The evolution of these ratios would indicate an increasing degree of contamination from the early gabbros, through the Nebdas gabbro–diorite, to gabbro–diorites and diorites of other plutons.
Felsic rocks Based on whole-rock and mineral chemistry data, four main groups are distinguished in the granitoid rocks of the area: (1) Nebdas trondhjemite; (2) Nebdas tonalite; (3) Askaoun and
NEOPROTEROZOIC SIROUA GRANITOIDS 500
500
100
100
10
10
Rock/chondrite
277
1
1
0.5
0.5
Ba Rb Th K. Nb
La Ce Sr Nd P. Sm Zr
Ti Tb Y Tm Yb
Early gabbro Ait Nebdas gabbro and diorite Askaoun diorite Ida ou Illoun gabbro and diorite Affela N’ouassif diorite Imdghar gabbro and diorite E-MORB BAT N-MORB
Ba Rb Th K. Nb
La Ce Sr Nd P. Sm Zr
Ti Tb Y. Tm Yb
Fig. 8. Extended normalization diagrams (Thompson et al. 1982) for the Siroua basic rocks from some geodynamic sites. Enriched (E)-MORB and normal, (N)-MORB from Sun & McDonough (1989) and back-arc tholeiites (BAT) from Wood (1980). 100
Precollisional tonalite
100
post-collisional High K-Calc-alkaline granite
ophiolitic plagiogranite
Rock/chondrite
0.1 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
0.1 La Ce Pr Nd
1000
1000
100
100
10
10
1
1
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Hybrid alkaline-high K-calc-alkaline granite
0.1 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
0.1 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Fig. 9. Chondrite-normalized (Sun & McDonough 1989) REE patterns for Siroua granitoids. Comparison with ophiolitic plagiogranite of Khzama (Samson et al. 2004), pre-collision tonalite, post-collision high K-calc-alkaline and hybird alkaline-high K-calc-alkaline granite from Tuareg shield (Lie´geois et al. 1998).
278
A. TOUIL ET AL.
Ifouachguel granodiorite; (4) granodiorite and monzogranite from Ida ou Illoun, Affela N’ouassif and Imdghar. (1) The Nebdas trondhjemite is characterized by high Sr (210 –465 ppm) and low Rb (28– 40 ppm) and K2O (0.4–1.3%). The Rb/Sr (0.06–0.15) and K/Rb (71 –208) ratios are comparable with those of continental trondhjemites (Coleman & Peterman 1975; Drummont & Defant 1990; Defant et al. 1991; Samson et al. 2004). This similarity is also found in the REE patterns (Fig. 9), which show a strong LREE/HREE fractionation, with (La/Yb)N between 5.7 and 11.8, and a conspicuous positive anomaly in europium. However, the evolution of the REE patterns; namely, the concomitant decrease of total content and increase of Eu anomaly, appears incompatible with the derivation of the trondhjemites from the associated gabbro– diorites and tonalites through fractional crystallization. The ‘sliding normalization’ parameters SNX and SNY defined by Lie´geois et al. (1998) have relatively low values, respectively 0.19– 0.38 and 0.08 – 0.14, comparable with those of pre-collisional tonalite – trondhjemite–granodiorite
(TTG) (Fig. 10). In the geodynamic diagrams of Pearce et al. (1984), the Nebdas trondhjemite falls, because of its low contents of Rb, Nb and Y, in the field of volcanic arc granitoids (Fig. 11). Spidergrams normalized to the ‘oceanic ridge granite’ reference values of Pearce et al. (1984) confirm this chemical affinity, with a slight enrichment in LILE and a strong depletion in high field strength elements (HFSE), especially niobium (Fig. 11). Such features are considered characteristic of continental trondhjemites formed by partial melting of a subducting oceanic crust (Arth 1979; Size 1984). (2) The Nebdas tonalite is characterized by low concentrations in LILE, HFSE and transition elements, considered typical of calc-alkaline arc magmas (Fig. 11). The ‘sliding normalization’ parameters SNX and SNY have low values, respectively 0.23 –0.27 and 0.25– 0.36, comparable with those of pre-collision TTG (Fig. 10). (3) Granodiorite from Askaoun and Ifouachguel has chemical features typical of high-K2O calcalkaline granites: weak enrichment in Fe2O3, Rb, Y and Nb, high K2O content, and K2O/Na2O of 0.5–1. This affinity is confirmed by the position
4.5 Ait Nebdas tonalite Ait Nebdas trondhjemite Ifouachguel granodiorite Askaoun granodiorite Ida ou Illoun granodiorite and granite Affela N’ouassif granodiorite and granite Imdghar granodiorite and granite
SNY = mean [Rb-Th-U-Ta]NYTS
4.0
3.5
3.0
2.5
2.0
HKCA + Shoshonitic
1.5
1.0
0.5
0.0 0.0
Pre-collisional
0.5
Alkaline
1.0
1.5
2.0
2.5
3.0
SNX = mean [Zr-Ce-Sm-Y-Yb]NYTS
Fig. 10. SNY v. SNX diagram for Siroua granitoids (Lie´geois et al. 1998).
NEOPROTEROZOIC SIROUA GRANITOIDS
279
1000
1000 Nb
Rb
Y/44 Syn-COLG WPG
WPG 100
100
Alcalin VAG Syn-COLG 10
10
ORG
VAG ORG
post-collision Syncollision 1
1 1
10
100
Y
1
1000
1000
1000
100
100
10
100
Nb+Y
1000
Rb/100
Nb/16
1000
Hybride Alkaline-high K-calc-alkaline Post collisional High K-Calc-alkaline granite
100
Precollisional tonalite 10
10
10
1
1
1
0.1
0.1
0.1
Cyprus Trondhjemite
0.01 Sr K
0.01 Rb Ba Th
Nb Ce P
Zr
Sm Ti Y
Yb
0.01 Sr K
Rb Ba Th
Nb Ce P
Zr
Sm Ti Y
Yb
Sr K
Rb Ba Th
Nb Ce P
Zr
Sm Ti Y
Yb
Fig. 11. Pearce et al. (1984), Nb (ppm) v. Y (ppm) and Rb (ppm) v. (Nb þ Y) (ppm) diagrams; Thie´blement & Cabanis (1990), (Rb/100)–(Y/44) –(Nb/16) and Pearce et al. (1984) exteneded normalization diagrams of Siroua granites. Symbols as for Figure 10. Comparison with pre-collisional and post-collisional granites from Tureg shield (Lie´geois et al. 1998) and Cyprus trondhjemite (Coleman & Peterman 1975).
in the AFM plot of Irvine & Baragar (1971) and in the geodynamic diagrams of Pearce et al. (1984) (Fig. 11). On Pearce et al. (1984) multi-element diagrams, the rocks of this group show strong enrichments in Rb, Th and Ce and depletions in Nb and Zr, which are considered typical of calc-alkaline magmas from island arcs (Perfit et al. 1980). The spidergrams and the values of SNX (0.67–0.71 and 0.49–0.65, respectively for the Ifouachguel and Askaoun granodiorites) and SNY (respectively 0.96 –1.02 and 1.05– 1.48) are comparable with those of many similar occurrences where a post-collisional character is established (Muller et al. 1992; Kuster & Harms 1998; Lie´geois et al. 1998). A post-collisional emplacement of these granodiorites is also consistent with their position in the Rb/100 –Y/44 –Nb/16 diagram of Thie´blemont & Cabanis (1990) (Fig. 11). (4) The significant iron enrichment observed in the granodiorites and monzogranites from Ida ou Illoun, Affela N’ouassif and Imdghar indicates an alkaline affinity, which is also shown by the compositions of biotites and amphiboles. Additional evidence for the sub-alkaline character of these rocks is given by the value of the MgO/TiO2 ratio: around 1.4 + 0.2, which is intermediate, according to a compilation by Bilal & Giret (1997), between
values typical of calc-alkaline (2 + 0.4) and anorogenic granites (1 + 0.5). In the Pearce et al. (1984) diagrams for geodynamic discrimination based on Rb, Y and Nb contents, these granites plot in an intermediate domain between the fields of island arc and within-plate granites, but they plot clearly within the field of anorogenic granites on the Thie´blemont & Cabanis (1990) discrimination diagram (Rb/100–Y/44 –Nb/16, Fig. 11). The spidergrams drawn with Pearce et al. (1984) normalizing values (Fig. 11), showing strong enrichments in Rb, Th and Ce and depletions in Nb and Zr, are similar to those of granites from Hoggar, interpreted by Azzouni-Sekkal et al. (2003) as transitional between calc-alkaline and alkaline. SNX and SNY are highly variable, especially in Ida ou Illoun granites, but they are generally intermediate between those of alkaline anorogenic and post-collisional high-K2O calcalkaline granites (Fig. 10).
Discussion It is now possible to discuss, in the light of the petrographical and chemical character of the rocks and the possible geodynamic settings, the types of
280
A. TOUIL ET AL.
associations encountered in the Siroua massif and their relations with other magmatic features in other districts of the belt. The early gabbros and diorites from Askaoun have compositions similar to those of oceanic and intraplate basalts. Such rocks, commonly encountered in continental rift settings (Harper & Link 1986; Hellingwerf 1987), have been reported from a number of sites in the Anti-Atlas belt. In the Bou-Azzer inlier (Central Anti-Atlas), affinities with MORB and within-plate basalts have been found (Leblanc & Moussine-Pouchkine 1994) for basaltic rocks from the volcanoclastic Bleida– Tachdamt, which is an equivalent of the Siroua series. According to Leblanc & MoussinePouchkine (1994), the emplacement of these basalts is related to the formation of deep crustal faults during the opening of a pull-apart basin in the Ediacaran. In the Jbel Saghro, the Tagmout tholeiites, Anou N’Izme pillow lavas and Imiter series form a consistent group of active margin formations that reflect an episodic stage of oceanization of Ediacaran age (Marini & Ouguir 1990; Mokhtari et al. 1995). In a similar way, the Askaoun gabbro–diorites may have been emplaced in a back-arc setting during an extensional episode. This would be consistent with the extensional setting (Regragui 1997; De Beer et al. 2000) of the Siroua detrital series in the Ediacaran. The gabbro–diorites, tonalites and trondhjemites from Nebdas represent calc-alkaline rocks with no conspicuous potassium enrichment. Such series are commonly encountered in subduction zone settings and are considered as products of the partial fusion of lithospheric mantle, with a contribution from subducted oceanic crust. The gabbro–diorites and associated felsic rocks from Ifouchguel and Askaoun, and the gabbro– diorites from Ida ou Illoun, Affela N’ouassif and Imdghar, represent high-K2O calc-alkaline rocks. Such rocks are found emplaced in a thickened crust following a collision event. The basic magmas parental to these rocks are considered here, as in similar magmatic associations described by Lie´geois & Black (1987), Lie´geois et al. (1987, 1998) and Kuster & Harms (1998), as products of the partial melting of a lithospheric mantle modified by fluids percolating from the underlying oceanic crust. The hybrid rocks, of monzodiorite to granodiorite composition, observed in Askaoun and Ifouachguel, would result from interactions between these basaltic magmas and anatectic melts produced by fusion of the lower crust at their contacts. The composition differences observed between Askaoun and Ifouachguel can be related to differences in conditions of crystallization, or to different degrees of interaction between mantle- and crust-derived magmas.
The transitional character of the granitoids from Ida Ou Illoun, Affela N’ouassif and Imdghar has been pointed out. These rocks have a close kinship with the ‘trans-alkaline’ monzonitic granites (Lameyre 1987), in which the ferro-magnesian minerals are consistently iron-rich, in contrast to the magnesian compositions generally encountered in monzonitic series. Such granites are commonly interpreted as markers of the transition from an orogenic calc-alkaline to an anorogenic alkaline magmatism, and represent the last magmatic activity of the post-collision stage (Lie´geois et al. 1998). This magmatism can be linked, as in Adrar des Iforas (Lie´geois et al. 1998), NE Africa (Kuster & Harms 1998) or Tibet (Turner et al. 1996), with a lithospheric delamination and subsequent asthenospheric upwelling that triggered the fusion of an enriched lithospheric mantle and induced melting of lower continental crust. The pluton emplacement seems to be related to the same deep lithospheric discontinuities that control this lithospheric delamination (Bonin 1990; Lie´geois & Black 1987; Black & Lie´geois 1993). In the Siroua massif, the shape and distribution of the Ida ou Illoun, Affela N’ouassif and Imdghar plutons suggest their emplacement along large-scale east –west to NE –SW strike-slip faults. The emplacement of the subalkaline granites shortly after the calc-alkaline gabbro–diorites supposes that the transition from one magmatic calc-alkaline type to the other subalkaline type takes place within a short time span.
Conclusion The Assarag Suite (Late Neoproterozoic plutonism) of the Siroua massif displays an evolution of the magmatism, in space and time, from subductionrelated in the northern part to within-plate subalkaline in the southern part. Similar evolutions have been described in other areas such as Adrar des Iforas, Mali (Lie´geois & Black 1987; Lie´geois et al. 1987, 1998), or Hoggar, Algeria (AzzouniSekkal et al. 2003). The oceanic subduction, concomitant with the closing of the Khzama oceanic basin, was accompanied by magmatism of oceanic type, dominantly basaltic (early gabbros of Askaoun massif and basalts associated with the volcanoclastic sequence), and sedimentation of thick volcanoclastic series in a back-arc basin. Related to the closing of this basin, magmatism of low-K calc-alkaline character developed, producing the Nebdas gabbro– diorite, tonalite and trondhjemite. The crustal thickening subsequent to the continental collision was marked by a strong magmatic activity, represented by the Askaoun and Ifouachguel granitoids, of high-K2O calc-alkaline character. The production of calc-alkaline magmas in the
NEOPROTEROZOIC SIROUA GRANITOIDS
subduction system continued during crustal upwelling (basic rocks of the Ida ou Illoun, Imdghar and Affela N’ouassif plutons). After uplift ceased, extension became significant, and the sub-alkaline granites of the Ida ou Illoun, Imdghar and Affela N’ouassif plutons were emplaced, marking the end of the Cryogenian (Late Neoproterozoic) magmatic cycle. During the Ediacaran (late Neoproterozoic), and still under an extensional regime, continental tholeiites with an alkaline signature (Touil et al. 1999a) were emplaced in the Siroua massif, as dykes cutting across the Assarag Suite and Sarhro Group. We warmly thank A. Giret for his constructive critiques.
References A BDEL -R AHMAN , A. M. 1994. Nature of biotites from alkaline, calc-alkaline, and peraluminous magmas. Journal of Petrology, 35, 525–541. A RTH , J. G. 1979. Some trace elements in trondhjemites; their implications to magma genesis and paleotectonic setting. In: B ARKER , F. (ed.) Trondhjemites, Dacites and Related Rocks. Elsevier, Amsterdam, 123–132. 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 ARKER , F. 1979. Trondhjemite: definition, environment and hypothesis origin. In: B ARKER , F. (ed.) Trondhjemites, Dacites and Related Rocks. Elsevier, Amsterdam, 1– 12. B ILAL , E. & G IRET , A. 1997. Aluminium saturation index and MgO/TiO2 ratio: two parameters influenced by PH2O and discriminant in magmatic series. International Symposium on Granites and Associated Mineralizations II, Salvador, Bahia, Brazil, 98– 100. 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 ONIN , B. 1990. From orogenic to anorogenic settings: evolution of granitoids suite after a major orogenesis. Geological Journal, W. S. Pitcher Special Issue, 25, 261–270. C HARLOT , R. 1982. Caracte´risation des e´ve´nements e´burne´ens et panafricains dans l’Anti-Atlas marocain. Apport de la me´thode ge´ochronologique Rb– Sr. Notes et Me´moires du Service Ge´ologique du Maroc, 313. C HOUBERT , G. 1963. Histoire ge´ologique de l’Anti-Atlas. Notes et Me´moires du Service Ge´ologique du Maroc, 162. C HOUBERT , G. & F AURE -M URET , A. 1970. Livret guide de l’excursion Anti-Atlas occidental et central. Colloque International sur les Correlations du Pre´cambrien. Notes et Me´moires du Service Ge´ologique du Maroc, 229.
281
C LARKE , D. B. 1981. The mineralogy of peraluminous granites: a review. Canadian Mineralogist, 19, 3 –17. C LAUER , N., C ABY , R., J EANETTE , D. & T ROMPETTE , R. 1982. Geochronology of sedimentary and metasedimentary Precambrian rocks of west African craton. Precambrian Research, 18, 53– 71. C OLEMAN , R. G. & P ETERMAN , Z. 1975. Oceanic plagiogranite. Journal of Geophysical Research, 80, 1099– 1108. D E B EER , C. H., C HEVALLIER , L. P., D E K OCK , G. S., G RESSE , P. G. & T HOMAS , R. J. 2000. Me´moire explicative de la carte ge´ologique du Maroc au 1/50 000, feuille Siroua. Notes et Me´moires du Service Ge´ologique du Maroc, 395b. D EFANT , M. J., M AURY , R. C., R IPLEY , E. M., F EINGENSON , M. D. & J ACQUES , D. 1991. An example of island-arc petrogenesis: geochemistry and petrology of southern Luzon arc, Philippines. Journal of Petrology, 32, 455– 500. D RUMMONT , M. S. & D EFANT , M. J. 1990. A model for trondhjemite –tonalite– dacite genesis and crustalgrowth via slab melting, Archean to modern comparisons. Journal of Geophysical Research, 95, 21503–21521. E L B OUKHARI , A. 1991. Magmatisme et me´tase´diments associe´s du prote´rozoı¨que supe´rieur de la zone de N’Kob (Siroua SE, Anti-Atlas Central, Maroc). Une ophiolite forme´e et mise en place sur la marge du craton ouest-Africain. PhD Thesis, University Cadi Ayad, Marrarkech. E L K HANCHAOUI , T. 2001. Le magmatisme Ne´oprote´rozoı¨que du massif de Siroua (Anti-Atlas Central, Maroc): marqueur de l’e´volution d’un contexte type d’arc a` un contexte intraplaque. PhD Thesis, University Cadi Ayad, Marrarkech. E L K HANCHAOUI , T., L AHMAM , M., E L B OUKHARI , A. & E L B ERAAOUZ , H. 2001. Les granitoı¨des ne´oprote´rozoı¨ques de Khzama, Anti-Atlas central, Maroc: marqueurs de l’e´volution d’un magmatisme d’arc a` un magmatisme alcalin. Journal of African Earth Sciences, 32, 655– 676. 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 NNIH , N. & L IE´ GEOIS , J. P. 2003. The Moroccan AntiAtlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton: reply to comments by E. H. Bouougri. Precambrian Research, 120, 185– 189. G RESS , P. G., D E B EER , C. H., C HEVALLIER , L. P., D E K OCK , G. S. & T HOMAS , R. J. 2000. Me´moire explicative de la carte ge´ologique du Maroc au 1/50 000, feuille Tachoukacht. Notes et Me´moires du Service Ge´ologique du Maroc, 393b. H ARPER , G. D. & L INK , P. K. 1986. Geochemistry of upper Proterozoic rift-related volcanics, northern Utah and southeastern Idaho. Geology, 14, 864– 867. H ELLINGWERF , R. H. 1987. Formation of sulfide deposits and its relation to sodic and potassic alteration of Proterozoic metabasites in the Sasca rift basin, Bergslagen, Sweden. Mineralium Deposita, 22, 53–63.
282
A. TOUIL ET AL.
I NGLIS , J. D., M AC L EAN , J. S., S AMSON , S. D., D’L EMOS , R. S., A DMOU , H. & H EFFRAN , 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. I RVINE , T. N. & B ARAGAR , W. R. A. 1971. A guide to the chemical classification of common volcanic rocks. Canadian Journal of Earth Sciences, 8, 523– 548. J OUIDER , A. 1997. Etude des granitoı¨des du prote´rozoı¨que supe´rieur du massif d’Ida ou Illoun (Siroua SW, Anti-Atlas central, Maroc): Pe´trographie, ge´ochimie, typologie des zircons et sites ge´odynamiques. Masters Thesis, University Cadi Ayad, Marrakech. K USTER , D. & H ARMS , U. 1998. Post-collisional potassic granitoids from the southern and northwestern parts of the late Neoproterozoic East African orogen: a review. Lithos, 45, 177–195. L AMEYRE , J., 1987. Granites and evolution of the crust. Revisita Brasileira de Geociencias, 17, 349–359. L EAKE , B. E., W OOLLEY , A. R., B IRCH , W. D. ET AL . 1997. Nomenclature of amphibole. Report of the subcommittee on Amphiboles of the International Mineralogical Association Commission on New Minerals and Mineral Names. European Journal of Mineralogy, 9, 623– 651. L EBLANC , M. 1976. A Proterozoic ocean crust at Bou-Azzer. Nature, 216, 34– 35. L EBLANC , M. & L ANCELOT , J. R. 1980. Interpre´tation ge´odynamique du domaine pan-africain (Pre´cambrien terminal) de l’Anti-Atlas (Maroc) a` partir des donne´es ge´ologiques et ge´ochronologiques. 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 ETERRIER , J., M AURY , R. C., T HONON , P., G IRARD , D. & M ARCHAL , M. 1982. Clinopyroxene composition as method of identification of the magmatic affinities of paleo-volcanic series. Earth and Planetary Science Letters, 59, 139–154. L IE´ GEOIS , J. P. & B LACK , R. 1987. Alkaline magmatism subsequent to collision in the Pan-African belt of the Adrar des Iforas. In: F ITTON , J. G. & U PTON , B. G. J. (eds) Alkaline Igneous Rocks. Geological Society, London, Special Publications, 30, 381 –408. L IE´ GEOIS , J. P., B ERTRAND , J. M. & B LACK , R. 1987. The subduction and collision related Panafrican composite batholith of the Adrar des Iforas (Mali): a review. Geological Journal, 22, 158– 211. 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. M ARINI , F. & O UGUIR , H. 1990. Un nouveau jalon dans l’histoire de la distension pre´-panafricaine au Maroc: le pre´cambrien II des boutonnie`res de Jbel Saghro nord oriental (Anti-Atlas, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 310, 577– 582.
M IYASHIRO , A. 1974. Volcanic rocks series in islands arcs and active continental margins. American Journal of Science, 27, 321– 355. M OKHTARI , A., G ASQUET , D. & R OCCI , G. 1995. Les thole´iites de Tagmout (Jbel Saghro, Anti-Atlas, Maroc) te´moins d’un rift au prote´rozoı¨que supe´rieur. Comptes Rendus de l’Acade´mie des Sciences, 320, 381–386. M ORIMOTO , N., F ABRIE` S , J., F ERGUSSON , A. K. ET AL . 1988. Nomenclature of pyroxenes. American Mineralogist, 173, 1123– 1133. M ULLER , D., R OCH , N. M. S. & G ROVES , D. I., 1992. Geochemical discrimination between shoshonitic and potassic setting: a pilot study. Mineralogy and Petrology, 46, 259– 289. N ACHIT , H., R AZAFIMAHEFA , N., S TUSSI , J. M. & C ARRON , J. P., 1985. Composition chimique des biotites et typologie magmatique des granitoı¨des. Comptes Rendus de l’Acade´mie des Sciences, 301, 813–818. P EARCE , J. A. & C ANN , J. R. 1973. Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth and Planetary Science Letters, 19, 290–300. P EARCE , J. A., H ARRIS , N. B. W. & T INDLE , A. G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956–983. P ERFIT , M. R., G UST , D. A., B ENCE , A. E., A RCULUS , R. J. & T AYLOR , S. R. 1980. Chemical characteristics of island arc basalts: implications for mantle sources. Chemical Geology, 30, 227– 256. P UPIN , J. P. 1980. Zircon and granite petrology. Contributions to Mineralogy and Petrology, 73, 207–220. R EGRAGUI , M. 1997. Le magmatisme et les volcanose´dimentaires du prote´rozoı¨que supe´rieur de la re´gion d’Askaoun (Siroua occidentale, Anti-Atlas central, Maroc). Cartographie, lithostratigraphie, pe´trographie et ge´ochimie. Masters Thesis, University Cadi Ayad, Marraikech. R ICKWOOD , P. C. 1989. Boundary lines within petrologic diagrams which use oxides of major and minor elements. Lithos, 22, 247– 263. S AMSON , S. D., I NGLIS , J. D., D’L EMOS , R. S., A DMOU , H., B LICHERT -T OFT , J. & H EFFRAN , K. 2004. 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 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 IZE , W. B. 1984. Polygenetic trondhjemite. Proceeding of the 27th International Geological Congress, Moscow, 9, 543– 559. S UN , S. S. & M C D ONOUGH , W. F. 1989. Chemical and isotopic systematics of oceanic basalts: implications for the 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.
NEOPROTEROZOIC SIROUA GRANITOIDS T HIEBLEMONT , D. & C ABANIS , B. 1990. Utilisation d’un diagramme (Rb/100)– Tb–Ta pour la discrimination ge´ochimique et l’e´tude des roches magmatiques acides. Bulletin de la Socie´te´ Ge´ologique de France, VI, 8, 23–35. T HOMAS , R. J., C HEVALLIER , L. P., D E B EER , C. H., D E K OCK , G. S. & G RESSE , P. G. 2000. Me´moire explicatif de la carte ge´ologique du Maroc au 1/50 000, feuille Assarag. Notes et Me´moires du Service Ge´ologique du Maroc, 392b. T HOMAS , R. J., C HEVALLIER , L. P., G RESSE , P. G. ET AL . 2002. Precambrian evolution of the Siroua 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, 29, 699–713. T HOMPSON , R. N., D ICKIN , A. P., G IBSON , I. L. & M ORRISON , M. A. 1982. Elemental fingerprints of isotopic contamination of hebridean paleocene mantle derived magmas by Archean SiAl. Contributions to Mineralogy and Petrology, 79, 159–168. T OUIL , A. 1999. Pe´trographie, ge´ochimie et contexte de mise en place des granitoı¨des du secteur ouest du massif du Siroua (Anti-Atlas Central, Maroc). PhD Thesis, University Cadi Ayad, Marrakech. T OUIL , A., E L B OUKHARI , A., B ILAL , E. & M OUTTE , J. 1999a. Les tholeiites a` affinite´ alkaline du secteur ouest de Siroua (Anti-Atlas central, Maroc): te´moins
283
d’une distension au Ne´oprote´rozoı¨que. Journal of African Earth Sciences, 39, 217– 226. T OUIL , A., E L B OUKHARI , A. & C HABANE , A. 1999b. The late Pan-African Ida ou Illoun granitoids (Central, Anti-Atlas, Morocco): an igneous province transitional from calc-alkaline to alkaline. Boletı´n de la Sociedad Espan˜ola de Mineralogia, 22, 109– 118. T URNER , S., A RNAUD , N., L IU , J. ET AL . 1996. Postcollision, shoshinitic volcanism on the Tibetan Platau: implications for convective thinning of the lithosphere and the source of ocean island basalts. Journal of Petrology, 37, 45–71. V ILLENEUVE , M. & C ORNE´ E , J. J. 1994. Structure, evolution and paleogeography of the west African craton and bordering belts during the Neoproterozoic. Precambrian Research, 69, 307 –326. W INCHESTER , J. A. & F LOYD , P. A. 1977. Geochemical discrimination of different magmas series and their differentiation products using immobile elements. Chemical Geology, 20, 325 –343. W OOD , B. J. & B ANNO , S. 1973. Garnet–orthopyroxene and orthopyroxene– clinopyroxene relationship in simple and complex systems. Journal of Petrology, 3, 238–243. W OOD , D. A. 1980. The application of a Th–Hf–Ta diagram to problems of tectonomagmatic classification and to establishing the nature of crustal contamination of basaltic lavas of the British Tertiary volcanic province. Earth and Planetary Science Letters, 50, 11–30.
Late Neoproterozoic carbonate productivity in a rifting context: the Adoudou Formation and its associated bimodal volcanism onlapping the western Saghro inlier, Morocco ´ LVARO1,2, ME´LINA MACOUIN3, HASSAN EZZOUHAIRI4, A. CHARIF4, J. JAVIER A N. AIT AYAD4, M. LUISA RIBEIRO5 & MAGALI ADER6 1
Departamento Ciencias de la Tierra, Universidad de Zaragoza, 50009 Zaragoza, Spain 2
Current present address: UMR 8014-LP3, USTL, 59655 Villeneuve d’Ascq, France (e-mail:
[email protected])
3
UMR 5563-LMTG, 14 Av. Edouard Belin, Universite´ Paul Sabatier, 31400 Toulouse, France 4
De´partement de Ge´ologie, Universite´ Chouaib Doukkali, 24000 El Jadida, Morocco
5
INETI—Area de Geocieˆncias, Estrada da Portela, Zambujal, 2721-866 Alfragide, Portugal
6
UMR 7047, Institut de Physique du Globe de Paris, 4 place Jussieu, 75005 Paris, France Abstract: An interval of episodic carbonate productivity, lithostratigraphically recognized as the ‘Calcaires infe´rieurs’ (upper member of the Adoudou Formation), took place across the Neoproterozoic–Cambrian transition onlapping the western Saghro inlier, Morocco. Sedimentation of the ‘Calcaires infe´rieurs’ was highly variable: in relatively stable substrates, a peritidal-dominated mixed platform is recorded where deposition was primarily controlled by autocyclic processes and accommodation space availability, whereas, in unstable substrates, the tectonic activity associated with the inherited block-faulting basement led to deposition of complex slide sheets composed of penecontemporaneous isoclinal folds and disrupted strata. The uppermost part of the ‘Calcaires infe´rieurs’ displays a negative d13C shift reaching values of 26.5‰. This shift may represent the d13C excursion to 26‰ that marks the Neoproterozoic– Cambrian boundary in the western Anti-Atlas. Two volcanic episodes bracketed the carbonate productivity. They consist of lower basaltic flows and an upper rhyolitic ignimbrite, with a SiO2 gap between 52 and 74 wt%. The basic rocks resemble those of tholeiitic magmas in continental rifts. The felsic rocks show high light to heavy rare earth element abundances and negative Nb, Ta, P and Ti anomalies, and were probably generated as a result of either fractional crystallization coupled with relative crustal contamination, or from a different magmatic source. The lower basic flows of tholeiitic affinity predated and geochemically differ from the alkaline magmatism of the Alougoum volcanic complex (Boho jbel) that surrounds the neighbouring Bou-Azzer inlier.
Across the Neoproterozoic–Cambrian transition, the tropical climate, high CO2 content of the atmosphere (Berger 1990) and low bathymetry were favourable for high carbonate productivity patterns in some epeiric seas that bordered West Gondwana. Thus, thick and vast carbonate platforms formed, where a major replacement of benthic communities took place by the beginning of the Atdabanian (c. 20 Ma after the Neoproterozoic–Cambrian boundary): a stromatolite-dominated consortium was replaced by shelly–metazoan and thromboid consortia ´ lvaro et al. 2006a), which led in the Moroccan (A Souss Basin to the occurrence of the so-called Early Cambrian Great Atlasian Reef Complex ´ lvaro & Clausen 2007). (A Another major factor that controlled the establishment and survival of carbonate factories across the Neoproterozoic –Cambrian transition in the
Souss Basin (Morocco) was the configuration of a basement inherited from the Pan-African orogeny and the succeeding onset of an aborted rifting. Volcanoclastic input and tectonically induced subsidence affected dramatically the accommodation space and the geometry of a basin that inherited a complex palaeotopographic basement. The basement of the Taroudant Group was primarily controlled by the superimposed effect of the Pan-African orogeny and the onset of a complex network of volcanic centres and volcanosedimentary deposition (the Ouarzazate Group). Subsequent pulses of rifting reactivation modelled the basin architecture and controlled the nucleation of centres of carbonate productivity (Taroudant Group). To study the interplay between carbonate productivity and the influence of inherited composite palaeotopographies (Pan-African and
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 285–302. DOI: 10.1144/SP297.14 0305-8719/08/$15.00 # The Geological Society of London 2008.
286
´ LVARO ET AL. J. J. A
Ouarzazate deposition), synrift volcanism and tectonic instability, we have focused our study on the characterization of the Adoudou volcanic and volcano-sedimentary complex that onlaps the western margin of the Saghro inlier in the central Anti-Atlas. This area offers a distinct palaeogeographical configuration across the Neoproterozoic– Cambrian transition related to: (1) the active erosion of the source area (the Saghro inlier) followed by the development of a brief episode of carbonate productivity; (2) the proximity to the source area that induced a dramatic reduction in carbonate thickness and the distinct terrigenous character of the overlying Lie-de-vin Formation (Tikirt Member); (3) the episodic instability of the carbonate sea floor; and (4) the record of two volcanic pulses that bracketed the carbonate productivity, different from the coeval alkaline magmatism of the Alougoum volcanic complex that surrounds the neighbouring Bou-Azzer inlier. The purpose of this paper is threefold: (1) to analyse the environmental and palaeogeographical factors that controlled the development of carbonate productivity recorded in the ‘Calcaires infe´rieurs’ that onlap the western Saghro inlier (Fig. 1); (2)
to propose a high-resolution sequence and chemostratigraphic framework for this onlapping, carbonate-dominated, depositional system; (3) to include this stratigraphic sketch in a magmatic context by characterizing, both petrologically and geochemically, the volcanic activity that bracketed this episode of carbonate productivity, and differentiating it from neighbouring volcanic episodes. The volcanic activity in the Ait Saoun area consists of two lower basaltic and andesitic –basalt flows, interbedded in the upper part of the basal massive conglomerates, and an upper acidic tuff capping the Adoudou Formation. Although the lower basic flows have been related to the activity of the Alougoum volcanic complex (flanking the Bou-Azzer inlier), it represents a different volcanic episode as demonstrated by its different geochemical affinity.
Geological setting and stratigraphy The Souss Basin is one of the sedimentary troughs that bordered the western Gondwanan margin during early Palaeozoic times (Geyer & Landing 1995). Infill of the Souss Basin unconformably
Fig. 1. (a) Regional map, and (b) geological map of the Saghro jbel and El Graara massif (central Anti-Atlas), showing locations of studied sections (modified from Service Ge´ologique du Maroc 1970).
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
overlies the Proterozoic basement, and consists of Neoproterozoic to Silurian (Late Carboniferous according to Helg et al. 2004) sediments. The basement of the Souss Basin was consolidated during the Pan-African orogeny (see the synthesis by Gasquet et al. 2005). This was succeeded by deposition of thick volcano-sedimentary complexes, sometimes considered as late Pan-African molasses, the so-called Saghro Supergroup (also named Precambrian II3 or PII3), subsequently overlain by the Ouarzazate (or PIII) Supergroup. The latter is also capped by a major unconformity (the contact of the Ouarzazate –Taroudant groups), which displays both angular discordance and conformable contacts, probably related to normal faulting and the growth and decay of volcanic centres (Soulaimani et al. 2003, 2004). After the Pan-African orogeny, broad shallow epeiric seas occupied extensive areas of the Souss Basin, which recorded a phase of intra-continental extension. This led to development of a multi-step rifting, characterized by magmatism with tholeiitic and alkaline affinities, and a distinct diachroneity related to the rifting propagation from the AntiAtlas to the western High Atlas and the Meseta domains (Pique´ et al. 1995). Finally, activity on this rift ceased in early Middle Cambrian times,
287
associated in the Anti-Atlas with a sharp interval of uplift, erosion and karstification recorded at the top of the Bre`che a` Micmacca, and related to a change from active rifting to more passive regional subsidence (Pique´ et al. 1995; Ait Ayad ´ lvaro & et al. 1998; Soulaimani et al. 2003; A Clausen 2006, 2008). The rifting-related volcanism recorded across the Neoproterozoic–Cambrian transition in the intra-cratonic Souss Basin was accompanied by marked asymmetric changes in sedimentary architecture, development of complex onlapping geometries, and sharp lateral variations in carbonate productivity. After the onset of the Ouarzazate – Taroudant unconformity, the subsequent base-level rise resulted in deposition of a transgressive siliciclastic unit (Soulaimani et al. 2004), composed of fluvial and alluvial conglomerates (the ‘Basal Series’ of the Adoudou Formation; Fig. 2). The unit passes upwards into a thick succession dominated by shallow-marine carbonates bearing scattered evaporitic pseudomorphs, and rich in stromatolitic and thombolitic dolostones (‘Calcaires infe´rieurs’ of the Adoudou Formation and Lie-de-vin Formation; Schmitt 1979). Although the related carbonate productivity attained c. 1000 m thickness in the central and distal parts of
Fig. 2. Stratigraphic summary of the Taroudant Group in the Anti-Atlas; after Boudda et al. (1979), Houzay (1979), Tucker (1986), Buggisch & Flu¨gel (1988), Latham & Riding (1990), Kirshvink et al. (1991), Magaritz et al. (1991), Compston et al. (1992), Geyer & Landing (1995), Landing et al. (1998), Gasquet et al. (2005) and Maalouf et al. (2005); the stratigraphic interval studied here is boxed.
288
´ LVARO ET AL. J. J. A
the Souss Basin (western Anti-Atlas), in the proximal and marginal parts of the basin (central and eastern Anti-Atlas) the carbonate productivity was restricted by the influence of accommodation space, neighbouring topographic relief and terrigenous input, and sometimes is absent (Destombes et al. 1985). The Adoudou Formation (also known as Adoudounian; Choubert 1952) forms the lower part of the Taroudant Group, and represents the oldest volcano-sedimentary complex unconformably covering the Neoproterozoic Ouarzazate Group (Fig. 2). The formation is currently divided into two members: the so-called ‘Basal Series’ or ‘Se´rie de Base’, less than 150 m thick, and the upper member or ‘Calcaires infe´rieurs’, 50–1000 m thick. In the depocentre of the Adoudou Formation, located in the western Anti-Atlas, the ‘Basal Series’ can be subdivided into three units: (1) the lower massive conglomerates (up to 150 m thick), dominated by massive conglomerates and secondary breccias, sandstones and shales, locally interbedded with volcanic ashes and flows, which are widely distributed throughout the Atlas Mountains overlying the Ouarzazate–Taroudant unconformity; (2) a middle unit (up to 50 m thick in jbel Imider and 200 m in the Anezi trough; Choubert 1952; Chazan 1954), named ‘petit calcaire’, composed of dolostones, limestones and shales, and bearing phosphoritic pockets and laminae in the High Atlas (Viland 1977); (3) an upper shale-dominated unit (up to 60 m thick; Boudda et al. 1979; Demange 1980). The upper member of the Adoudou Formation or ‘Calcaires infe´rieurs’ is a monotonous succession of bedded and massive dolostones, locally interrupted by shales and sandstones. The base of the ‘Calcaires infe´rieurs’ is characterized in the western Anti-Atlas by a dolostone unit, 50–200 m thick, partly silicified and locally mineralized by diverse ores, named the Tamjout Bed (Chazan 1954; Demange 1980). The occurrence of thick variegated shale interbeds marks the transition to the overlying Lie-de-vin Formation (also named Taliwinian; Boudda et al. 1979). The palaeogeographical distribution of the carbonate factories related to the ‘Calcaires infe´rieurs’ also depended on the record of a continental volcanism. A striking volcanic activity is reported from the central Anti-Atlas, where several trachyte, trachyandesite and basalt flows and numerous pyroclastic tuffs are intercalated. The centre of volcanism has been conventionally associated with the Alougoum volcanic complex located in the El Gloa area (Choubert 1952; Boudda et al. 1979; Fig. 2), in which volcanic palaeotopographies (such as the Boho volcano) directly cover the Adoudou dolostones and were progressively onlapped by the breccias, dolostones, variegated shales and
´ lvaro sandstones of the Lie-de-vin Formation (A et al. 2006b). An early U/Pb date of 534 + 10 Ma from the Boho volcano (Ducrot & Lancelot 1977) has been recently refined to 531 + 5 Ma by Gasquet et al. (2005). Three U/Pb dates are also available from the uppermost part of the Adoudou Formation and the overlying Lie-de-vin Formation: 525 + 0.46, 522 + 2 and 521 + 7 Ma (Compston et al. 1992; Landing et al. 1998; Maalouf et al. 2005), all of them below the first appearance of biostratigraphically significant fossils, located in the overlying Igoudine Formation and Atdabanian ´ lvaro et al. 2006a). The in age (Sdzuy 1978; A palaeontological record of the Adoudou Formation is represented by the conspicuous presence of microbial mats and stromatolites, and the spotty occurrence of the red alga Kundatia in the upper member (Buggisch et al. 1978; Buggisch & Heinitz 1984; Buggisch & Flu¨gel 1988), and a possible finding of medusoid imprints in the upper shales of the ‘Basal Series’ (Houazy 1979). The Neoproterozoic –Cambrian boundary is tentatively located in the upper part of the Adoudou dolostones (lower Tifnout Member) based on a chemostratigraphic peak of d13C correlatable with coeval successions in Siberia and South China (Tucker 1986; Latham & Riding 1990; Kirshvink et al. 1991; Magaritz et al. 1991). Our study area is located at the westernmost edge of the Saghro inlier (Fig. 1), where two sections of the ‘Calcaires infe´rieurs’ have been studied: they are located in the vicinity of Id Boukhtir (30 km to the WSW of Ouarzazate) and on the western flank of the Tizi n’Tinfifit jbel, in the vicinity of Ait Saoun (38 km to the SSE of Ouarzazate). There, the ‘Basal Series’ is composed exclusively of massive conglomerates and basic lava flows, the ‘Calcaires infe´rieurs’ are only 50 – 60 m thick (disappearing c. 15 km to the NE; Boudda et al. 1979), and the overlying Lie-de-vin Formation is represented by the Tikirt Member, a monotonous succession of reddish sandstones and shales, devoid of carbonates, and up to 300 m thick. This paper documents the facies and sequence arrangement of the dolostones that characterize the ‘Calcaires infe´rieurs’, and the mineralogical and geochemical features of the bimodal volcanism that bracketed this reduced episode of carbonate productivity. This consists of two lower basaltic and andesitic –basalt flows, interbedded in the upper part of the basal massive conglomerates, and an upper acidic tuff capping the Adoudou Formation. Although the lower basic flows have been currently related to the activity of the Alougoum volcanic complex (located flanking the Bou-Azzer inlier, 40 km to the SSW of the study area; Choubert & Faure-Muret 1970; Boudda et al. 1979; Fig. 2), it represents a different
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
volcanic episode, as demonstrated by its different geochemical affinity.
Facies and sequence arrangement of the ‘Calcaires infe´rieurs’ in the Ait Saoun area Peritidal mixed (carbonate – siliciclastic) facies association The lower and upper parts of the ‘Calcaires infe´rieurs’ are characterized, in the Ait Saoun area (Fig. 1), by a heterogeneous lithology composed of carbonate and shale –litharenite on the bedform and bedset scale. Four major facies occur; these are, in order of decreasing abundance: (1) variegated shale; (2) grainy dolostone; (3) fenestral, silty–cherty sparry dolostone; (4) stromatolitic dolostone (Figs 3 and 4a). The intervals of purple and reddish shale are commonly less than 3 m thick, and contain interbedded gravel-to-litharenite beds. The latter are up to 10 cm thick, and can be followed laterally more than 10 m. Texture consists of a mediumsorted, subrounded to subangular, medium-grained litharenite dominated (.60% in volume) by felsic pyroclasts rich in euhedral feldspar (partially to wholly replaced by albite and sericite), engulfed quartz, and accessory mafic phenocrysts (partly chloritized and illitized), and containing also minor proportions of subrounded gravels of sparry dolostone, chert and shale intraclasts (Fig. 4b). This facies is either lithified by a calcareous matrix or poorly cemented. Sedimentary structures are high (c. 258) and low (c. 108) planar cross-lamination, low-amplitude (2–6 cm), wide (c. 50 cm) basal scours, and positive grading. The grainy dolostone consists of thin- to medium-bedded, coarsely crystalline strata, with moderately preserved packstone to grainstone textures (Fig. 4c and d). Skeletal grains are absent, and allochems are peloids, peloidal aggregates, millimetre-sized oncoids, quartz sand grains, and dolomite and shale intraclasts. Peloids are structureless spheres of fine-grained dolomite that occur as single spheres and as flocculent coalesced clots or aggregates. Preservation of peloids and oncoids is locally poor and identification in the field is difficult because of pervasive dolomitization. Cross and planar laminations are present but poorly preserved. The base of the grainy dolostone beds is commonly erosive, and covered by a lag deposit (up to 30 cm thick), composed of contorted and distorted centimetre-thick beds of sparry dolostone, laterally fragmented into angular intraclasts (Fig. 4e and f). Fenestral, silty–cherty dolostone strata are 0.1–0.8 m thick, and display partially silicified
289
dolomicrosparite to dolopseudosparite textures. Chert nodules and layers, up to 10 cm thick, are concentrated along the bedding planes, but rarely form a continuous bed. Thin sections of the dolomite show disseminated, silt-sized, quartz and mica grains. Sedimentary structures include wavy and low-angle lamination, shaly partings, sharp lower contacts, and transitional upper boundaries grading into stromatolitic dolostones. The fenestral fabric is characterized by the presence of thinly disrupted laminae bearing abundant, millimetre-sized, spar-filled fenestrae and vugs, or consists of thick layers disrupted by discontinuously bedded, spar-filled fenestrae. The stromatolitic dolostone consists of a basal unit (0.1 –0.4 m thick), which comprises fenestral and silty dolomite couplets, and an upper unit (up to 1.4 m thick) composed of microbial flat, wavy or crinkled microbial laminites capped by a continuous package of decimetre-thick, dome-shaped stromatolitic mounds. The silty dolomite has sharp bases that are, in some cases, lined by flat pebble intraclastic lags. The stromatolitic fabric is punctuated with stylolitic and argillaceous seams, where chert nodules and laminae also occur (Fig. 4g). Stromatolitic laminae are characterized by subtle changes in the size of dolomite crystals, and are occasionally delineated by either a slight increase in abundance of quartz silt grains or cherty laminae. Stringers of quartz silt, vugs, and cherty laminae also delineate lamination in the intermound fill. Domal stromatolites form unlinked mounds, with up to 50 cm synoptic relief, grading laterally into low-relief domal stromatolites or wavy microbial dolostone. Stromatolitic heads are commonly separated by c. 10 –50 cm wide areas filled with silty dolomitic packstone that forms an anastomosing and laterally migrating pattern between mounds. At the top of the domal stromatolites, they are smoothly laminated, have low synoptic relief, and are irregularly cracked. Graded, crosslaminated dolomitic grainstone (similar to the aforementioned grainy dolostone) fills the relief between the heads of the uppermost dome-shaped stromatolites and the centimetre-scale cracks. The episodic input of litharenite interbeds suggests a local pyroclastic source. The lower peloidal dolopackstone to grainstone represents subtidal to possible intertidal peloidal sheets (or low-angle shoals) with an intermittent pyroclastic flux. Sharp contacts and conglomeratic rip-ups in the base of some beds suggest storm energy sufficient to disrupt the sea floor. The depositional environment of the coarse-crystalline dolostone is not clear but the minor interbeds of the nodular, peloidal mudstone to packstone may suggest a subtidal setting. An agitated shallow subtidal environment is envisaged for the stromatolite-dominated parts.
290
´ LVARO ET AL. J. J. A
Fig. 3. Stratigraphic logs of the ‘Calcaires infe´rieurs’ from the Ait Saoun and Id Boukhtir sections. (a) Legend. (b) d13C‰ PDB isotopic data from calcite and dolomite. (c) d13C/ d18O crossplot of dolostone samples. (d) Arrangement of facies associations, cycles and chemostratigraphic variations.
The high synoptic relief of the stromatolites and the lack of distinct subaerial exposure surfaces within associated sediments indicate that microbial mounds accreted in a wave-dominated subtidal setting, a view supported by the high-energy grainstone fill between mounds.
Offshore dolomite – shale rhythmites This second facies association is located in the middle part of the ‘Calcaires infe´rieurs’ (Fig. 3) and is monotonously laminated. Muddy and microsparitic dolomite laminae are interlaminated with reddish and purple shale. The facies is for the most part monotonously laminated, with local development of cross-laminated siltstone lenses, 2–5 cm thick, and lenticular-to-flaser fabrics lacking erosive contacts (Fig. 4h). Microsparitic dolomite is light red, massive to thinly bedded, and has moderately well-preserved mudstone
textures. Sedimentary structures include plane beds, wavy lamination, thin laminae of clay, and thin beds of siltstone. Because of the scarcity of erosive and current sedimentary structures, these rhythmites were deposited on a calm substrate, which rhythmically recorded a fine-grained siliciclastic influx transported via decantation, and was episodically affected by distal storm events below fair-weather wavebase.
Cycles and controls The aforementioned peritidal facies association is arranged into 1.5–5 m thick cycles or parasequences (Fig. 3). The cycle boundaries of these mixed (carbonate –siliciclastic) cycles are recognized on the basis of erosive surfaces located at sharp lithological contacts followed by sharp deepening. The cycles begin with (1) purple and
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
291
Fig. 3. Continued.
reddish shale with interbedded, gravel-to-litharenite, pyroclastic event beds, (2) grainy (peloidaldominated grainstone to packstone) and fenestral, silty–cherty dolostone, and (3) wavy-to-crinkled microbial dolostone, which passes upwards into domal stromatolites. The top of the stromatolitic
domes commonly contains cracks and fissures (crosscutting a fully rigid sediment), which are filled and covered by (4) peloidal –intraclastic dolograinstone –rudstone and fenestral dolostone, common also in the lows between stromatolitic domes. Larger stromatolites locally pass laterally
292
´ LVARO ET AL. J. J. A
Fig. 4. (a) Panorama of the western flank of the Tizi n’Tinfifit jbel (Ait Saoun) with arrowed base and top of the ‘Calcaires infe´rieurs’ Member; b, basalt; ab, andesitic basalt. (b) Volcanoclastic limestone showing clasts, composed of partly silicified (left rim) mudstone (mud), chert (ch) and felsic debris (fel), embedded in a micritic matrix (mi); scale bar represents 1 mm. (c) Contorted beds of sparry dolostone and breccia intraclasts at the base of the peritidal mixed cycles; erosive surfaces are arrowed, the uppermost marking the contact with the overlying peloidal-dominated grainstone; scale bar represents 4 cm. (d) Lateral transition from the previous view, showing abundant dolostone clips and the same upper contact (arrowed); scale bar represents 10 cm. (e) Peloidal and oncoidal grainstone encrusted by a microbial film (arrowed), subsequently covered by a volcanoclastic lag; scale bar represents 1 mm. (f) Detail of peloidal and oncoidal grainstone showing the perfect preservation of microbial textures despite dolomitization and neomorphism; scale bar represents 500 mm. (g) Stromatolitic dolostone (d) punctuated by silicified strings (s); scale bar represents 1 mm. (h) Reddish, microsparite dolomite–shale rhythmites showing parallel and low-angle lamination, and some contorted contacts where the shales are thicker.
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
into smaller stromatolites, peloid-rich microbial dolomite and intraclastic rudstone. The lower part of the shallowing-upward (coarsening-upward) cycles may record a shoaling depositional sequence, with interbedded shales and dolostones giving way to peloidal-dominated, grainstone low-angle shoals and sheets, indicating environments of increasing energy. This lower part of the cycle is capped by fenestral to stromatolite mounds where microbial laminae pass vertically from flat, crinkled and wavy into dome-shaped. The stromatolitic mounds are capped by cracked surfaces, which may reflect subaerial exposure and tepee structures, filled by dolomitic grainstone and rudstone. Efforts to trace cycles in the available outcrop showed that they are extremely variable, although they cannot be traced laterally more than 200 m, and component facies pinch out or interfinger with the other facies. The poor traceability of these peritidal cycles suggests an autogenic origin. Overall, carbonate production and accumulation in the peritidal complex was sufficient to track sea-level fluctuations, so it can be considered that the environment remained shallow (Pratt et al. 1992).
293
On a larger scale, the cycles are stacked in a transgressive-to-regressive depositional system primarily controlled by the availability of accommodation space. The depositional system is limited, at its bottom and top, by two regressive coarse-grained siliciclastic deposits: the ‘Basal Series’ and the Tikirt members. In addition, the transgressive and regressive trends of the depositional system are separated by the aforementioned mixed rhythmites, which represent a major flooding of the platform (Fig. 3).
Facies and sedimentary geometry of the ‘Calcaires infe´rieurs’ at Id Boukhtir At Id Boukhtir, the strata of the ‘Calcaires infe´rieurs’ (Van Looy 1985) lack distinct cycles, as a result of the presence of numerous erosive truncations (Fig. 5a). A suite of c. 58 m of stromatolitic dolostones (Fig. 5b) and local litharenites is underlain and internally interrupted by successive sedimentary unconformities (Fig. 3). These are interpreted as large-scale, sliding flows where carbonate strata are commonly disrupted and slump
Fig. 5. (a) Panorama of the Id Boukhtir section taken from the Tazenakht– Ouarzazate road; the base and top of the slide sheet described in the text are arrowed. (b) Dome-shaped stromatolite unaffected by sliding. (c) Conical folds made up of wavy-to-crinkled stromatolites; hinge points are arrowed; width of compass is 7 cm. (d) Schmidt net plot of the crest lines of the measured cylindrical and conical synsedimentary folds, some of them bearing crestline culminations and depressions, and their setting after correcting the stratification plane to the horizontal.
294
´ LVARO ET AL. J. J. A
folded, on a small and medium scale, into disharmonic tight and isoclinal folds, with cylindrical and conical shapes (Fig. 5c). Their fold axes and crest lines trend c. 340 –3508, after correction for tectonic dip (Fig. 5d), which is broadly parallel to the associated inlier border. These geometries fit Schlische’s (1992) concept of ‘drag fold’. Lineations and foliations, which are commonly roughly axial planar to folds (Fig. 5c), are also observed in the hinges of slump folds. Models that attempt to explain the genesis of penecontemporaneous foliation involve simple load compaction superimposed on preexisting folds or rotation of grains facilitated by slump folding (Tobisch 1984). The presence of these gravity-related sedimentary structures (abundant in other laterally equivalent strata; Buggsich & Heinitz 1984) suggests that the ‘Calcaires infe´rieurs’ were deposited in this area in an unstable slope environment, where gravity movements disturbed the autochthonous (microbially dominated) reef sedimentation. As a whole, this sedimentary system can be interpreted as a large, plastically deformed olistolith or slide sheet of semi-consolidated dolostones (see the geometry shown in Fig. 5a). According to the average fold axis orientation, the sea floor of the Id Boukhtir section was a SW-facing slope bordering the western Saghro inlier basement. The pronounced difference in facies between the Ait Saoun and Id Boukhtir sections, separated by only c. 50 km, indicates the development of a steep-sloped, possibly faulted-bounded, intra-shelf basin.
Geochemical methods We analysed for chemostratigraphic purposes 24 samples for carbonate d13C and d18O spanning a mixed stratigraphic interval of 50 m from the ‘Calcaires infe´rieurs’, to test the occurrence of correlatable shifts in d13C (Fig. 3). Between 1 and 10 mg of powdered samples were reacted with H3PO4 at 25 8C for 18 h to extract the CO2 from the calcite, and then at 80 8C for 2 h to extract the CO2 from the dolomite (the method has been described in detail by Swart et al. 1991), although a perfect chemical separation of calcite and dolomite is difficult (Al-Assam et al. 1990; Yui & Gong 2003). The amount of extracted CO2 and its carbon and oxygen isotopic compositions were measured using a helium continuous-flow mass spectrometer (AP-2003) at the IPGP laboratory (Paris). The reproducibility of the d13C and the d18O measurements is +0.1‰ and +0.2‰, respectively. Geochemical data from igneous rocks are based on nine chemical analyses of representative samples (Table 1). Major, trace and rare earth elements
were determined using X-ray fluorescence and inductively coupled plasma mass spectrometry (ICP-MS) at the INETI laboratory in Porto. Precision for major and trace elements is usually better than 2% and 5–10%, respectively.
Chemostratigraphic control Recently, Maalouf et al. (2005) have updated and proposed a composite d13C curve for the Neoproterozoic –Cambrian transition in Morocco. They defined two members for the Adoudou Formation in the western Anti-Atlas, named the Tabia and Tifnout members. However, the presence of peritidal carbonates in both members makes difficult to differentiate them in the central and eastern AntiAtlas, where Maalouf et al. identified only the Tabia Member. The stratigraphically condensed character of the ‘Calcaires infe´rieurs’ bordering the western Saghro inlier is probably due to low sedimentation rates and not to sharp erosive surfaces, absent in the facies associations described above. This can prevent the identification of the Tabia– Tifnout contact in the study area, which is a key marker bed for chemostratigraphic correlation. The latter is based on a d13C excursion to 26‰ at the base of the Tifnout Member (Fig. 2) that is correlated with the Neoproterozoic – Cambrian boundary in Siberia and Death Valley, and dated at 542 + 0.6 Ma in Oman by Amthor et al. (2003) and Maalouf et al. (2005). A positive d13C shift of 7‰ located in the upper part of the Tifnout Member was proposed by the same workers as the Nemakit–Daldinian –Tommotian boundary (Early Cambrian; Fig. 2). d18O values vary between 22.3 and 28.7‰ and are in accordance with those obtained by Maalouf et al. (2005). When plotting the d13Ccarbonate as a function of the d18Ocarbonate (Fig. 3c), the lack of correlation between d13C and d18O indicates that these samples have not been strongly modified by meteoric diagenesis. In addition, the profiles of d13Ccalcite and d13Cdolomite are parallel, reflecting the lack of diagenetic processes affecting preferentially the d13C values in calcite or dolomite. The d13C profile is not correlatable with facies or with parasequence trends. d13C background values for this section gradually increase (from 22.4 or 21.3 to 20.6 or þ0.4‰) and decrease to broadly constant values ranging from 24 to 22‰, as shown in the lower and middle parts of the section. However, this trend is interrupted in the upper part of the section by two sudden negative shifts of d13C values to c. 21.3 to 21.8‰ in the lower, and c. 23.2 to 23.6% in the upper part. After both negative shifts, the values sharply decrease to background averages.
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
295
Table 1. Selected chemical analyses from the basalts, andesitic basalts and rhyolites sampled in the Ait Saoun area Sample:
AS-1
AS-2
AS-3
AS-4
AS-5
AS-6
Ouri1
AS-12
AS-13
(wt %) SiO2 Al2O3 Fe2O3* MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total
46.08 15.02 15.84 0.11 8.15 2.93 3.67 0.8 3.53 0.34 3.08 99.55
45.89 15.12 15.6 0.17 5.69 4.53 3.73 1.73 3.36 0.34 3.24 99.4
46.63 14.23 17.25 0.15 6.06 3.39 4.4 1.27 3.26 0.33 2.6 99.57
51.8 13.43 12.19 0.1 4.82 2.81 3.03 3.83 3.1 0.64 4.02 99.77
49.64 13.43 13.19 0.09 5.59 2.97 2.98 3.48 3.11 0.65 4.64 99.77
50.22 13.91 13.35 0.08 8.79 1.39 2.15 3.28 3.31 0.68 2.38 99.54
73.75 10.63 2.99 0.45 0.55 0.4 0.13 8.54 0.45 0.08 1.37 99.36
75.63 9.14 2.62 0.06 0.21 1.61 0.33 7.91 0.49 0.12 1.66 99.78
78.26 8.38 2.35 0.04 0.31 1.26 0.2 7.1 0.46 0.13 1.43 99.92
17 56 38 143 4 101 0.24
42 61 38 144 5 73 0.3
21 96 35 137 5 161 0.3
41 107 63 285 11 617 0.66
37 91 66 288 11 438 0.66
30 58 66 283 11 287 0.66
4.29 13 30.4 4.6 24.4 6.2 2.5 7 1.1 6.6 1.4 3.6 0.5 3.4 0.5
4.32 10.7 26.4 4.2 21.3 5.8 2.3 6.6 1.1 6.5 1.3 3.6 0.5 3.3 0.5
4.11 9.4 24.2 3.9 19.3 5.2 1.9 5.8 1 6.1 1.3 3.4 0.5 3.2 0.5
8.55 21.9 56.1 8.2 39.9 10.2 3.4 11.2 1.8 10.9 2.3 6.2 0.9 5.8 0.9
8.64 25.1 60.9 8.9 42.6 10.7 3.8 11.5 1.9 11.2 2.3 6.3 0.9 5.9 0.9
8.49 23 59.2 8.6 41.4 10.6 3.2 11.9 1.9 11.5 2.4 6.6 1 6.3 1
83 35 23 153 7 1357 0.74 13.8 4.56 20 43.9 5.2 21.1 4.7 1.3 4 0.7 4.1 0.9 2.4 0.4 2.4 0.4
59 19 13 133 5 497 0.3 6 3.99 7 33
55 20 12 129 4 489 0.24 6 3.87 14 21
( ppm) Rb Sr Y Zr Nb Ba Ta Th Hf La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
13 2
,6
7 2
,6
*Total iron as Fe2O3. LOI, loss on ignition.
In the Ait Saoun area, the upper negative d13C shift reaches values of 26.4 or 26.6‰, which may represent the d13C excursion to 26‰ located at the base of the Tifnout Member and proposed as the Neoproterozoic–Cambrian boundary in the western Anti-Atlas (Maalouf et al. 2005). However, the lack of carbonate strata in the overlying regressive Tikirt Member precludes any possible identification of the succeeding evolution of the d13C values in this proximal area. In addition, the associated d18O values (22.3 and 23.7‰ for d18Ocalcite and d18Odolomite, respectively) appear higher than in others samples (23.2 to 28.1‰ and 25.3 to 28.7‰ for d18Ocalcite and
d18Odolomite, respectively) displaying similar lithologies and facies, and might indicate a diagenetic signal. As a result, our chemostratigraphic correlation of the Neoproterozoic –Cambrian boundary in the Ait Saoun area is only tentatively proposed.
Mineralogical and geochemical features of the associated volcanism As explained above, an episodic volcanic activity was coeval in the study area with carbonate productivity across the Neoproterozoic –Cambrian transition. This is documented by the presence of felsic
296
´ LVARO ET AL. J. J. A
ash encased in carbonates. However, the carbonate productivity recorded in the Adoudou volcanosedimentary complex was bracketed between two major episodes of effusive and pyroclastic activity: a lower basic episode interbedded within the uppermost coarse-grained siliciclastic strata of the ‘Basal Series’, and an upper acidic episode located at the Adoudou–Lie-de-vin contact. Field, petrological
and geochemical analyses were undertaken to improve understanding of the eruptive style of these volcanic rocks to provide insight into their environment of deposition. The first basic volcanic episode is represented by two lava flows that occur interbedded in the uppermost part of the ‘Basal Series’ (Figs 4a and 6b). This member displays sharp changes in
Fig. 6. (a) Conformable contact of the basal conglomerates (Cg) and the basal flow (B). (b) Thin-section photomicrograph of the basalt lava showing its microlithic to slightly porphyritic texture; polarized light; scale bar represents 2 mm. (c) Breccia formed at the basaltic lava– red sandstone contact. (d) Thin-section photomicrograph of the andesitic basalt displaying a microlithic texture (note the presence of quartz floating in the groundmass); crosspolarized light; scale bar represents 1 mm. (e) Ignimbritic rhyolite capping the Adoudou Formation. (f) Thin-section photomicrograph of the rhyolitic tuff showing a felsic texture; scale bar represents 0.4 mm. Chl, chlorite; FeMg, ferromagnesian minerals; Pl, plagioclase; Q, quartz; R, rhyolite; stro, stromatolitic dolostone.
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
Fig. 7. Harker diagrams (Harker 1909) (in wt%) for the two volcanic episodes recorded in the Ait Saoun area. S, Basalt; þ, ignimbritic rhyolite.
(Fig. 7). The chondrite-normalized rare earth elements (REE) projection of basalts and andesitic basalts displays parallel profiles suggesting their differentiation from a common source; in contrast, the profile of the acidic rocks crosscuts the others, suggesting their provenance from a different source (Fig. 8). All the rock types display enrichment in light REE (LREE), which is greater in the rhyolites (La/Yb ¼ 8.5) despite their low total amount. The Y/Nb variations between 2.5 and 9.5
200
Sample/C1 Chondrite
thickness, reaching 80 m to the east of Ait Saoun, where the contact is marked by a NNE– SSW-striking fault, which probably acted as a normal fault coevally with deposition of the ‘Basal Series’. In the study area, interflow strata consist of centimetre-bedded, siltstone-to-conglomerate wedge-shaped units. Laterally, the lava flow changes sharply into a matrix-supported breccia. The lower lava flow is vesicular, dark grey basalt, 10–15 m thick. It has a microlithic texture, locally changing to porphyritic. The microphenocrysts are predominantly euhedral plagioclase (2 mm in size), and highly altered femic minerals (mainly olivine and pyroxene), widely replaced by an association of chlorite, epidote, calcite, and iron oxide. The groundmass is rich in plagioclase and iron oxide microliths (Fig. 6b). Flow tops are vesicular, where irregular centimetre-sized amygdales are occluded by chlorite and calcite. This lower lava flow is overlain by a breccia-like bed (1–3 m thick) that displays a mixture of volcanic pockets, with indistinct outlines, and fine- to coarsegrained litharenites (Fig. 6c), reflecting the partial reworking of the lava embedded in a coarse-grained siliciclastic sea floor. The upper lava flow, c. 15 m thick, is grey andesitic basalt flow with a vacuolar microlithic texture. The microphenocrysts consist of plagioclase (c. 1 mm long), femic minerals (less abundant than in the lower basalt flow and generally altered), and microliths of iron oxide, feldspar and rare tiny quartz (Fig. 6d). The top of this lava flow is also overlain by another ‘breccia’ of decimetre scale, composed of magmatic pockets and litharenites. The younger acidic volcanic episode, which marks the Adoudou–Lie-de-vin contact, is only 0.2–0.5 m thick and has up to 200 m lateral extent (Fig. 6e). Although the thickness and lateral extent of this ignimbritic rhyolite, which changes laterally into a pyroclastic tuff, may initially appear relatively insignificant, they may yield clues to the understanding of the formation of siliceous magmas. Petrologically, the ignimbritic rhyolite shows a fine equigranular felsic texture (Fig. 6f), where the phenocrysts are composed of albite with polysynthetic twins, K-feldspar, quartz, intergranular iron oxide, and subsidiary chlorite. Carbonate and siltstone clasts, derived from the underlying strata, are also present. From the geochemical compositions, three rock type were distinguished in the volcanic episodes of the Adoudou volcano–sedimentary complex: basalts (45% , SiO2 , 47% and 15.5% , Fe2O3 , 17%), andesitic basalts (49% , SiO2 , 52% and 12% , Fe2O3 , 13.5%), and rhyolites (SiO2 . 74%) (see Table 1). The Harker diagram representation of the available data displays a hiatus on the intermediate field
297
100
+
+ + + + +
++
++ + + + +
10 La Pr Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb Fig. 8. C1 chondrite-normalized (Sun & McDonough 1989) REE patterns from the basalt flow (S), andesitic basalt flow (S) and rhyolitic ignimbrite (þ) found in the Ait Saoun area.
´ LVARO ET AL. J. J. A
298
(a)
(b)
FeO* Zr / Y
16
5 12
8
4
Tholeiitic . + Domain C. alk Calc-alkaline Domain Th
3 Plate margin basalt FeO*/MgO
2
4
6
8
Within-plate basalt Ti/Y
250
500
750
Fig. 9. Geochemical affinity of the Ait-Saoun basalts according to (a) Miyashiro & Shido (1975) and (b) Pearce & Gale (1977) diagrams.
indicate that these rocks belong to the sub-alkaline type. The projection of the basic rocks on Miyashiro & Shido (1975) and Pearce & Gale (1977) diagrams indicates they are intra-plate tholeiites (Fig. 9). The Zr/Hf, Ti/Zr and Nb/La ratios (33, 64–140 and 0.31 –0.53, respectively) are very close to those reported for continental tholeiites (Dupuy & Dostal 1984; Marsh 1987). The spider diagram of these rocks is also similar to that for the Sabie River and Lesotho (South Africa) continental tholeiites (Marsh 1987) (Fig. 10). The Nb negative anomaly, visible in all the profiles, is interpreted as indicative of crustal contamination (Bertrand 1975; Wood 1980; Dupuy & Dostal 1984; Wilson 1994). The volcanism that brackets the establishment and demise of carbonate factories (‘Calcaires infe´rieurs’) in the study area is bimodal, starting with large volumes of effusive basaltic lava flows and ending with a reduced effusive and explosive, rhyolitic volcanism. Intermediate compositions are absent, as the major element and SiO2 content is dominated by basic and acidic end-members (Fig. 7). The present geochemical data do not permit us to propose any linkage between the basic and acidic volcanic episodes. Crustal contamination is obvious in both episodes, mostly in the acidic volcanism. The two episodes may represent either the extreme poles of a geochemical fractionation derived from a common magma with strong crustal contamination, or may have resulted from different sources with a crustal source for the acidic episode. In this case, the basaltic magmatic chamber could have contributed to melt the crustal source of the acidic episode. The basic-to-acidic rock differentiation is a common process in continental tholeiitic settings (Dostal et al. 1983; Marsh 1987; Wilson 1994).
As explained in the ‘Geological setting’ section, the volcanism was considered as being associated with that preserved in the Boho volcano (from the Alougoum volcanic complex; Fig. 2). However, the latter shows alkaline affinities and evolved, via crystal fractionation, from basic to intermediate and ´ lvaro et al. 2006b). The basic tholeiitic acidic (A volcanism of the Ait Saoun area is different from both the late Neoproterozoic K–calc-alkaline magmatism recorded in the Ouarzazate Supergroup
Fig. 10. MORB-normalized (Pearce 1983) trace element patterns for representative basalts (S) and rhyolites (þ) of the Ait Saoun area, compared with continental tholeiitic basalts from Lesotho (B) and the Sabie River (†), South Africa (after Marsh 1987), and ´ lvaro et al. 2006b). alkaline basalt from Boho Jbel (O) (A
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX
from the Ouarzazate–Agdz region (Ezzouhairi 2001; Ezzouhairi et al. 2008), and the alkaline magmatism of the neighbouring Jbel Boho (Bou-Azzer inlier).
Palaeogeographical and geodynamic implications All along the margins of the western Saghro inlier, and postdating the unconformity located at the Ouarzazate –Taroudant contact, variations in thickness, facies and development of stratigraphic discontinuities within the ‘Basal Series’ reflect the syndepositional activity and along-strike variations in displacement of intrabasinal faults (Schlische 1992). Fault scarps along the basin margins had offsets of tens of metres, and remained significant morphological escarpments on the basin floor during deposition of the ‘Basal Series’. Increased sediment supply outpaced the subsidence rate, and the palaeorelief was finally levelled, favouring nucleation of carbonate factories. Much of this relief was reduced by the accumulation of alluvial and fluvial sediments (‘Basal Series’) in topographic lows prior to transgression and establishment of carbonate factories (‘Calcaires infe´rieurs’). Main tensional and transtensional faults arranged the depocentres in a block-faulting basement system and determined the regional dip of the onlapping carbonate strata. The differential tilted fault-block setting of the Ait Saoun and Id Boukhtir sections that flank the western Saghro inlier had important consequences for sequence-stratigraphic interpretations of the overlying ‘Calcaires infe´rieurs’ because of asymmetric subsidence, which caused drastic changes in bounding surface character and sequence thickness. Sedimentation of the ‘Calcaires infe´rieurs’ in relatively stable margins (e.g. Ait Saoun) is organized in distinct peritidal cycles, separated by surfaces of erosion and sharp deepening, whereas unstable sea floors (e.g. Id Boukthir) recorded non-cyclic intervals dominated by slope-related deposits. Tectonism and sediment supply (both from pyroclastic and source areas) were the main controls on accommodation-space fluctuations and the stratigraphic architecture of the margins surrounding the western Saghro inlier. Other mixed platforms bordering neighbouring inliers (such as the Bou-Azzer inlier; Fig. 1) recorded similar peritidal environments and low thickness of the ‘Calcaires infe´rieurs’, rich in stromatolites, sheet-cracks, fenestrae and evaporitic pseudomorphs (Chbani et al. 1999). Finally, the prograding Tikirt Member that overlies the Adoudou Formation westwards and northwards may be related to either a sea-level fall, not clearly recognized in distal areas (western Anti-Atlas), or
299
tilting and flexuring of proximal blocks of the Anti-Atlas margin (Maalouf et al. 2005). The latter, also related to volcanic acidic activity, would lead to enlargement of the emerged source areas, resulting in increased sediment supply and thereby feeding and favouring the progradation of the Tikirt sedimentary system. The geochemical features of the bimodal volcanism (with large amounts of basic volcanic lava and reduced rhyolite tuffs) described in the study area testify to the initial rifting stages of a thickened crust similar to the continental tholeiites of Lesotho and the Sabie River in South Africa. The beginning of this rifting took place after the Pan-African orogeny, as documented by the K –calc-alkaline volcanism recorded in the Ouarzazate Supergroup (or Precambrian III; Ezzouhairi 2001), and was reactivated after the unconformity that marks the Ouarzazate – Taroudant contact (Pique´ et al. 1995; Soulaimani 2001, 2003, 2004). Discontinuity surfaces, and the thickness and sedimentary character of the ‘Basal Series’ molasses, vary greatly depending on their location within the inherited palaeotopography, because of an asymmetric subsidence pattern of tilted fault-blocks associated with a widespread extensional fault and volcanic activity (e.g. Benssaou & Hamoumi 2001; Ezzouhairi et al. 2003; Benssaou 2005). Across the Neoproterozoic –Cambrian transition, the partial burial of the inherited palaeotopography coeval with a widespread transgression favoured development in the study area of carbonate productivity, which recorded the input of numerous felsic pyroclasts. The abundance of shard-rich pyroclasts indicates that much of the pyroclast debris was derived from a neighbouring felsic subaerial or shallow subaqueous explosive volcanism. The final demise and disappearance of the carbonate productivity is also associated with the accentuation of an intermittent explosive acidic volcanism and the establishment of regressive conditions indicated by the progradation of the overlying Tikirt Member. These chemostratigraphic data can also be compared with Maalouf et al.’s (2005) sections 13– 15 from the El Graara inlier. These are located at similar easterly positions in the basin, and are characterized by similar stratigraphic frameworks (e.g. presence of a bimodal volcanism, autobreccias, and peritidal carbonates). In all of the El Graara sections, d13C values rise upsection from a low of 24‰ and then oscillate between 22‰ and þ4‰, mimicking the isotopic profile of the Tifnout Member in the western Anti-Atlas. However, at Ait Saoun, d13C drops rapidly from 0‰ to 24‰, and then essentially remains at 23‰ to 24‰ for the remaining 40 m of the section. A correlation of this drop in d13C with the Neoproterozoic– Cambrian boundary isotopic
300
´ LVARO ET AL. J. J. A
shift (possibly observed in the upper part of the Tifnout Member in the western Anti-Atlas) would suggest that the basic volcanism that onlaps the western Saghro inlier is older than 542 Ma (see Fig. 2). This strengthens the lithostratigraphic argument (the basic lavas occur embedded in the uppermost part of the Basal Series) that the Saghro and El Graara volcanism are of different age, and that accommodation space in the eastern regions of the Anti-Atlas margin was also controlled by block faulting and thermal subsidence in the vicinity of volcanic sources.
Conclusions The western margin of the Saghro inlier in the central Anti-Atlas is an example of a mixed platform with low carbonate production rates and accumulation space. There, the establishment, demise and style of carbonate facies were directly influenced by the neighbouring post-Pan-African palaeorelief and coeval volcanic activity. Succeeding the Ouarzazate –Taroudant unconformity and sedimentation of stacked alluvial and fluvial conglomerates in the lows of the study area, carbonate deposition started during the following transgression, as evidenced by their related onlapping geometries. The facies associations and strata arrangement of the ‘Calcaires infe´rieurs’ in relatively stable substrates (Ait Saoun section) reflect both autocyclic and accommodationspace controls on a peritidal-dominated mixed platform. Carbonate factories were in part microbially induced, as evidenced by the widespread development of stromatolitic mats and mounds. In contrast, the instability of the platform related to the tectonic activity associated with the inherited block-faulting basement is illustrated at Id Boukhtir, where the stromatolite-dominated ‘Calcaires infe´rieurs’ form complex slide sheets composed of penecontemporaneous isoclinal folds and a disrupted stratification. The uppermost interbedded dolostones of the ‘Calcaires infe´rieurs’ display in the Ait Saoun area a negative d13C shift reaching values of 26.4 to 26.6‰. This shift may represent the d13C excursion to 26‰ located at the base of the Tifnout Member (Adoudou Formation), which marks the Neoproterozoic–Cambrian boundary in the western Anti-Atlas. The episode of carbonate productivity represented by the ‘Calcaires infe´rieurs’ was bracketed between two episodes of tholeiitic volcanism. In the study area, the synrift volcanism was bimodal in character, and comprises two lower basaltic and andesitic basalt flows and upper ignimbritic rhyolites, with a SiO2 gap between 52 and 74 wt%. The basic rocks resemble those of the tholeiitic magmas in continental rifts. The felsic acidic rocks show high large ion lithophite element
abundances and negative Nb, Ta, P and Ti anomalies, and were probably generated either as a result of fractional crystallization coupled with relative crustal contamination, or from a different magmatic source. This bimodal volcanism predated and differs from the alkaline volcanism that occurred bordering the neighbouring Bou-Azzer inlier, indicating the presence of different magmatic sources associated with a rifting that postdated the Pan-African orogeny and ended in Middle Cambrian times. The authors acknowledge the numerous useful remarks made by U. Linnemann and A. C. Maalouf, which have helped to improve the ideas expressed in this paper. Research on the Adoudou volcano-sedimentary complex has been supported by ECLIPSE project ‘Evolution of biogeochemical cycles from Archean to Recent environments’, and GRICES/CNRST project ‘A comparative study of the Neoproterozoic– Cambrian transition between the Portuguese Ossa–Morena Zone and the Moroccan Anti-Atlas and Meseta: geologic and geochemical aspects, and geodynamic model’. This paper is a contribution to CGL2006-13533 Programme and IGCP project 485 ‘Cratons, metacratons and mobile belts: keys from the West African craton boundaries, Eburnean versus Pan-African signature, magmatic, tectonic and metallogenic implications’.
References A IT A YAD , N., R IBEIRO , M. L., M ATA , J., F ERREIRA , P., E ZZOUHAIRI , H., C HARIF , A. & D IAS , R. 1998. Evolution du magmatisme cambrien en deux re´gions pe´rigondwanniennes: Azegour (Haut-Atlas) et Alter do Chao-Elvas (NE Alentejo). Communicac¸aos do Instituto Geologico e Mineiro do Portugal, 84, B154–B157. A L -A SSAM , I., T AYLOR , B. E. & S OUTH , B. 1990. Stable isotope analysis of multiple carbonate samples using selective acid extraction. Chemical Geology, 80, 119–125. ´ LVARO , J. J. & C LAUSEN , S. 2006. Microbial crusts as A indicators of stratigraphic diastems in the Cambrian Micmacca Breccia, Moroccan Atlas. Sedimentary Geology, 185, 255 –265. ´ LVARO , J. J. & C LAUSEN , S. 2007. Botoman A (Lower Cambrian) turbid- and clear-water reefs and associated environments from the High Atlas, ´ LVARO , J. J., A RETZ , M., B OULVAIN , Morocco. In: A F., M UNNECKE , A., V ACHARD , D. & V ENNIN , E. (eds) Palaeozoic Reefs and Bioaccumulations: Climatic and Evolutionary Controls. Geological Society, London, Special Publications, 275, 51–70. ´ LVARO , J. J. & C LAUSEN , S. 2008. Paleoenvironmental A significance of hiatal shelled accumulations in a Cambrian intracratonic aborted rift, Atlas Mountains, Morocco. Geological Association of Canada, Special Paper on Epeiric Seas (in press). ´ LVARO , J. J., C LAUSEN , S., E L A LBANI , A. & A C HELLAI , E. H. 2006a. Facies distribution of Lower Cambrian cryptic microbial and epibenthic archaeocyathan–microbial communities in the
ADOUDOU VOLCANO-SEDIMENTARY COMPLEX western Anti-Atlas, Morocco. Sedimentology, 53, 35– 53. ´ LVARO , J. J., E ZZOUHAIRI , H., V ENNIN , E. ET AL . A 2006b. The Early Cambrian Boho volcano of the El Graara massif, Morocco: petrology, geodynamic setting and coeval sedimentation. Journal of African Earth Sciences, 44, 396–410. A MTHOR , J., G ROTZINGER , J., S CRO¨ DER , S., B OWRING , S., R AMEZANI , J., M ARTIN , M. & M ATTER , A. 2003. Extinction of Cloudina and Namacalathus at the Precambrian– Cambrian boundary in Oman. Geology, 31, 431–434. B ENSSAOU , M. 2005. Le Rift Cambrien infe´rieur de l’Anti-Atlas, Maroc: remplissage se´dimentaire, ge´odynamique et pale´oge´ographie. The`se Doctorat d’Etat, Universite´ Ibn Zohr, Agadir. B ENSSAOU , M. & H AMOUMI , N. 2001. L’Anti-Atlas occidental du Maroc: e´tude se´dimentologique et reconstitutions pale´oge´ographiques au Cambrien infe´rieur. Journal of African Earth Sciences, 32, 351–372. B ERGER , R. A. 1990. Atmospheric carbon dioxide over Phanerozoic time. Science, 249, 1382–1386. B ERTRAND , H. 1975. Les dole´rites marocaines et l’ouverture de l’Atlantique: e´tude pe´trologique et ge´ochimique. The`se 3e`me cycle, Universite´ de Lyon. B OUDDA , A., C HOUBERT , G. & F AURE -M URET , A. 1979. Essai de stratigraphie de la couverture se´dimentaire de l’Anti-Atlas: Adoudounien, Cambrien infe´rieur. Notes et Me´moires du Service Ge´ologique du Maroc, 271, 1 –96. B UGGISCH , W. & F LU¨ GEL , E. 1988. The Precambrian/ Cambrian boundary in the Anti-Atlas (Morocco). Discussion and new results. In: J ACOBSHAGEN , V. H. (ed.) The Atlas System of Morocco. Studies on its Geodynamic Evolution. Lecture Notes in Earth Sciences, 15, 81– 90. B UGGISCH , W. & H EINITZ , W. 1984. Slumpfolds and other early deformations in the early Cambrian of the western and central Anti-Atlas (Morocco). Geologische Rundschau, 73, 809– 818. B UGGISCH , W., M ARZELA , C. & H U¨ GEL , P. 1978. Die fazielle und pala¨ogeographische Entwicklung der infrakambrischen bis ordovizischen Sedimente im Mittleren Antiatlas um Agdz (S-Marokko). Geologische Rundschau, 68, 195–224. C HAZAN , W. 1954. Les gisements stratiformes plombozincife`res de l’Infracambrien de l’Anti-Atlas occidental (Maroc). Notes et Me´moires du Service Ge´ologique du Maroc (VIII), 120, 97– 126. C HBANI , B., B EAUCHAMP , J., A LGOUTI , A. & Z OUHAIR , A. 1999. Un enregistrement se´dimentaire e´ocambrien dans un bassin intracontinental en distension: le cycle ‘conglome´rats de base–unite´ calcaire– gre`s de Tikirt’ de Bou-Azzer El Graara (Anti-Atlas central, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 329, 317– 323. C HOUBERT , G. 1952. Histoire ge´ologique du domaine de l’Anti-Atlas. In: C HOUBERT , G. & M ARC¸ AIS , J. (eds) Ge´ologie du Maroc. XIX Congre`s Ge´ologique International d’Alger. Monographies Re´gulie`res (Se´rie 3, Maroc), 6, 77– 194. C HOUBERT , G. & F AURE -M URET , A. 1970. Livret-guide de l’excursion Anti-Atlas occidental et central. Notes et Me´moires du Service Ge´ologique du Maroc, 299.
301
C OMPSTON , W., W ILLIAMS , J. L., K IRSCHVINK , J. L., Z HANG , Z. & M A , G. 1992. Zircon U–Pb ages for the Early Cambrian time scale. Journal of the Geological Society, London, 127, 319–332. D EMANGE , M. 1980. Stratigraphie, volcanisme et pale´oge´ographie du Pre´cambrien III et de la se´rie de base dans la partie sud de la boutonnie`re d’Ouaoufenrha (Anti-Atlas occidental). Mine´ralisations associe´es. Notes et Me´moires du Service Ge´ologique du Maroc, 285, 7–23. D ESTOMBES , J., H OLLARD , H. & W ILLEFERT , S. 1985. Lower Palaeozoic rocks of Morocco. In: H OLLAND , C. H. (ed.) Lower Palaeozoic Rocks of the World. Vol. 4. Lower Palaeozoic of North-Western and West Central Africa. Wiley, Chichester, 157– 184. D OSTAL , J., D UPUY , C. & B ARAGAR , W. R. A. 1983. Geochemistry and petrogenesis of basaltic rocks from Coppermine River area, Northwest Territories. Canadian Journal of Earth Sciences, 20, 684–698. D UCROT , J. & L ANCELOT , J. R. 1977. Proble`me de la limite Pre´cambrien– Cambrien: e´tude radiochronologique par la me´thode U/Pb sur zircon du volcan du Jbel Boho. Canadian Journal of Earth Sciences, 14, 1771–1777. D UPUY , C. & D OSTAL , J. 1984. Trace element geochemistry of some continental tholeiites. Earth and Planetary Science Letters, 67, 61–69. E ZZOUHAIRI , H. 2001. Le magmatisme post-collisionnel panafricain (tardi a` post-oroge´nique) des re´gions d’Aghbalou, Sidi Flah-Bouskour et Oued Imini (Ouarzazate, Anti-Atlas central, Maroc). Lithostratigraphie, ge´ochimie, pe´trogene`se et contexte ge´odynamique. PhD thesis, University Chouaı¨b Doukkali, El Jadida. E ZZOUHAIRI , H., R IBEIRO , M. L., C HARIF , A., A IT A YAD , N., R AMOS , F., M OREIRA , M. E. & C OKE , C. 2003. Contribution a` la caracte´risation pe´trographique et ge´ochimique du volcanisme cambrien moyen du moˆle coˆtier (Maroc). In: REMAL , T. (ed.) 3rd International Colloquium on Magmatism, Metamorphism and Associated Mineralization, University Hassan II - Mohammadia, Casablanca, 8 –10 Mai, Re´sume´s, 49–50. E ZZOUHAIRI , H., R IBEIRO , M. L., A IT A YAD , N. ET AL . 2008. The magmatic evolution at the Moroccan outboard of the West African Craton between the Late Neoproterozoic and the Early Palaeozoic. In: E NNIH , N. & L IE´ GEOIS , J. -P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 329– 343. G ASQUET , D., L EVRESSE , G., C HEILLETZ , A., R ACHID A ZIZI -S AMIZ , M. & 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 EYER , G. & L ANDING , E. (eds) 1995. Morocco’95. The Lower–Middle Cambrian standard of Gondwana. Beringeria, Special Issue, 2, 1– 171. H ARKER , A. 1909. The Natural History of Igneous Rocks. Hafner, New York (1965 facsimile of original edition). H ELG , U., B UKHARD , M., C ARITG , S. & T OBERT -C HARRUE , C. 2004. Folding and inversion tectonics in the Anti-Atlas of Morocco. Tectonics, 21, TC4006, doi:10.1029/2003TC001576.
302
´ LVARO ET AL. J. J. A
H OUZAY , J. P. 1979. Empreintes attribuables a` des me´duses dans la se´rie de base de l’Adoudounien (Pre´cambrien terminal de l’Anti-Atlas, Maroc). Ge´ologie Me´diterrane´enne, 6, 379–384. K IRSHVINK , J. L., M AGARITZ , M., R IPPERDAN , R. L. & Z HURAVLEV , A. YU . 1991. The Precambrian/ Cambrian boundary: Magnetostratigraphy and carbon isotopes resolve correlation problems between Siberia, Morocco, and South China. GSA Today, 1, 69–71. L ANDING , E., B OWRING , S. A., D AVIDEK , K. L., W ETROP , S. R., G EYER , G. & H ELDMAIER , W. 1998. Duration of the Early Cambrian: U– Pb ages of volcanic ashes from Avalon and Gondwana. Canadian Journal of Earth Sciences, 35, 329– 338. L ATHAM , A. & R IDING , R. 1990. Fossil evidence of the location of the Precambrian/Cambrian boundary in Morocco. Nature, 344, 752 –754. M AALOUF , A. C., S CHRAG , D. P., C ROWLEY , J. L. & B OWRING , S. A. 2005. An expanded record of Early Cambrian carbon cycling for the Anti-Atlas margin, Morocco. Canadian Journal of Earth Sciences, 42, 2195–2216. M AGARITZ , M., K IRSHVINK , J. L., L ATHAM , A. J., Z HURAVLEV , A. YU . & R OZANOV , A. YU . 1991. Precambrian/Cambrian boundary problem: Carbon isotope correlations for Vendian and Tommotian time between Siberia and Morocco. Geology, 19, 847–850. M ARSH , J. S. 1987. Basalt geochemistry and tectonic discrimination within continental flood basalt provinces. Journal of Volcanology and Geothermal Research, 32, 35– 49. M IYASHIRO , A. & S HIDO , F. 1975. Tholeiitic and calc-alkaline series in relation to the behaviors of Ti– Cr–Ni. American Journal of Science, 275, 265– 277. P EARCE , J. A. 1983. Role of the sub-continental lithosphere in magma genesis at active continental margins. In: H AWKESWORTH , C. J. & N ORRY , M. J. (eds) Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, 230– 249. P EARCE , J. A. & G ALE , G. H. 1977. Identification of ore-deposition environment from trace element geochemistry of associated igneous host rocks. In: FYFE , W. S. (ed.) Volcanic Processes in Ore Deposits. Geological Society, London, Special Publications, 7, 14–24. P IQUE´ , A., B OUABDELLI , M. & D ARBOUX , J. R. 1995. Le rift cambrien du Maroc occidental. Comptes Rendus de l’Acade´mie des Sciences, 320, 1017–1024. P RATT , B. R., J AMES , N. P. & C LINTON , A. C. 1992. Peritidal carbonates. In: W ALKER , R. G. & J AMES , N. P. (eds) Facies Models: Response to Sea Level Change. Geological Association of Canada, Ottawa, Ont., 303–322. S CHLISCHE , R. W. 1992. Structural and stratigraphic development of the Newark extensional basin, eastern North America: Evidence for the growth of the basin and its bounding structures. Geological Society of American Bulletin, 104, 1246–1263. S CHMITT , M. 1979. The section of Tiout (Precambrian/ Cambrian boundary beds, Anti-Atlas, Morocco): stromatolites and their biostratigraphy. Arbeiten aus dem Pala¨ontologischen Institut Wu¨rzburg, 2, 1 –188.
S DZUY , K. 1978. The Precambrian –Cambrian boundary beds in Morocco (preliminary report). Geological Magazine, 115, 83–94. SERVICE GE´ OLOGIQUE DU M AROC 1970. Carte ge´ologique de l’Anti-Atlas central et de la zone synclinale de Ouarzazate. Feuilles Ouarzazate, Alougoum et Telouet Sud, 1/200 000. Notes et Me´moires du Service ge´ologique du Maroc, 138. S OULAIMANI , A., P IQUE´ , A. L. & B OUABDELLI , M. 2001. La se´rie du PII– PIII de l’Anti-Atlas occidental (Sud marocain): un olistostrome a` la base de la couverture post-panafricaine (PIII) du Prote´rozoı¨que supe´rieur. Comptes Rendus de l’Acade´mie des Sciences, 332, 121– 127. S OULAIMANI , A., B OUABDELLI , M. & P IQUE´ , A. L. 2003. L’extension continentale au Ne´oprote´rozoı¨que supe´rieur-Cambrien infe´rieur dans l’Anti-Atlas (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, 174, 83–92. 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 (Anti-Atlas occidental, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 336, 1433– 1441. 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 Oceanic Basins. Geological Society, London, Special Publications, 42, 313–345. S WART , P. K., B URNS , S. J. & L EDER , J. J. 1991. Fractionation of the stable isotopes of oxygen and carbon in carbon dioxide during the reaction of calcite with phosphoric acid as a function of temperature and technique. Chemical Geology, Isotopic Geosciences Section, 86, 89– 96. T OBISCH , O. 1984. Development of foliation and fold interference patterns produced by sedimentary processes. Geology, 12, 441– 444. T UCKER , E. 1986. Carbon isotope excursions in Precambrian/Cambrian boundary beds, Morocco. Nature, 319, 48–50. V AN L OOY , J. 1985. Het kaartblad Tazenakht 1/100.000, Anti-Atlas, Marokko. Kartiering, Lithostratigrafie, Biostratigrafie van Precambrium tot Tremadoc. PhD thesis, Katholieke Universiteit, Leuven. V ILAND , J. C. 1977. Pre´sence d’horizons phosphate´s a` la base de l’Infracambrien supe´rieur du Haut Atlas de Marrakech (Maroc). Notes et Me´moires du Service ge´ologique du Maroc, 268, 13–22. W ILSON , M. 1994. Igneous Petrogenesis. A Global Tectonic Approach. Harper Collins, London. W OOD , D. A. 1980. The application of the Th– Hf–Ta diagram to problems of tectono-magmatic classification, and to establishing the nature of the crustal contamination of basaltic lavas of the British Tertiary Province. Earth and Planetary Science Letters, 50, 11–30. Y UI , T. F. & G ONG , S. Y. 2003. Stoichiometry effect on stable isotope analysis of dolomite. Chemical Geology, 201, 359 –368.
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 A. POUCLET1, H. OUAZZANI2 & A. FEKKAK3 1
Institut des Sciences de la Terre d’Orle´ans, UMR 6113, Universite´, B.P. 6759, 45067 Orle´ans Cedex 2, France (e-mail:
[email protected])
2
Faculte´ des Sciences, Universite´ Moulay-Ismail, B.P. 4010, Mekne`s, Morocco
3
Faculte´ des Sciences, Universite´ Chouaı¨b Doukhali, B.P. 20, 24000 El Jadida, Morocco Abstract: In the westernmost part of the High Atlas, two Palaeozoic formations, rich in mafic volcanic rocks, are distinguished. They belong to different structural blocks created during the Variscan orogeny. New U– Pb dating yields an Early Cambrian age. The basaltic lavas have the composition of continental tholeiites and the magmatic signature of an initial rifting tectonic setting. They are related to the western Moroccan Cambrian rift. Their geodynamical context could be a passive margin initiated from an active rift that aborted in the Middle Cambrian.
The western High Atlas is a mountainous region composed of Neoproterozoic and Palaeozoic formations folded during the Variscan orogeny and raised by the Atlas (Alpine) tectonic event (Fig. 1). The oldest formations, consisting of volcano-plutonic and volcano-sedimentary rocks dated from the Late Neoproterozoic to the Early Cambrian period, crop out in the southeastern part, between the northern and southern branches of the Tizi-n’Test Fault. The High Atlas westernmost part exhibits Early Cambrian to Early Carboniferous marine sedimentary and volcanic formations unconformably overlain by Late Carboniferous to Triassic continental detrital formations. In this area, with the exception of the granodioritic intrusion of Wirgane, along the N’Fis Fault and dated at 598 + 5 Ma (Eddif 2002), no Neoproterozoic rocks are known. The Early Cambrian formations are dated by very few occurrences of palaeontological remains (Archaeocyatha), but most of them are totally azoic. Neoproterozoic ages have been suggested for the Ifri –Azegour formation (Ouazzani et al. 1998, 2001) (Fig. 2). However, this formation, as well as the Ouzaga formation located to the south, includes volcanic rocks that can be analysed and dated. The aim of this paper is to describe the lithological nature of these sedimentary piles related to the Cambrian history, to determine their distinct tectonic features relative to Variscan events, and to date and analyse the volcanic rocks. In the literature, very few data are available for these volcanic-rich formations, which are undifferentiated throughout the middle and higher part of the western High Atlas (Schaer et al. 1981; Corne´e 1989). On the basis of the
magmatic signature of the lavas, the geodynamical context is determined and placed in the Neoproterozoic to Cambrian evolution of the Atlas region.
Lithology and tectonic of the Cambrian western High Atlas In the middle part of the western High Atlas, we distinguish two main Cambrian structural and stratigraphic formations (Fig. 2): the Ifri–Azegour formation to the north and NE and the Ouzaga – Tizzirt formation to the south. These formations are separated by a major fault system, the Middle Western High Atlas Fault. This fault system strikes N60–708, from south of Talmakant and north of the Ida ou Zal area to the N’Fis Fault, south of Wirgane. It consists of reverse faults dipping 50–608 to the south.
The Ifri – Azegour formation The Ifri–Azegour formation crops out in two areas, Ifri and Azegour, separated by a Cretaceous graben (Fig. 2). It includes fine-grained terrigenous sediments, shales, siltstones and sandstones, alternating with limestones and volcanic products. The two areas show the same structural features, with an initial interfolial micro-folding developing a schistosity subparallel to the stratification under greenschist-facies conditions, followed by folding with SSW –NNE-striking axes that developed in upper-crustal conditions and shows a typical knee-fold pattern. As a result of the highly dissected
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 303–327. DOI: 10.1144/SP297.15 0305-8719/08/$15.00 # The Geological Society of London 2008.
304
A. POUCLET ET AL.
Fig. 1. Distinction of the main Atlas regions. WHA, Western High Atlas; MHA, Middle High Atlas. Bold lines, main faults: MWHAF, Middle Western High Atlas Fault; TNTF, Tizi-n’Test Fault; SB, southern branch; NB, northern branch; SAF, South Atlas Fault; AAMF, Anti-Atlas Major Fault. TNTF, SB and SAF separate High Atlas and Anti-Atlas. Mk, Marrakech; Ag, Agadir; Oz, Ouarzazate.
Fig. 2. Geological map of the Palaeozoic part of the Western High Atlas. MWHAF, Middle Western High Atlas Fault; TNTF, Tizi-n’Test Fault; SB, southern branch; NB, northern branch.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
topography, and mining and coring studies, the stratigraphic colum may be established in the Ifri area. The Ifri area: lithology. In the Ifri area, we distinguish three successive units, I-U1, I-U2 and I-U3 (Figs 3 and 4). The whole recovered sequence consists of 1400 m of shallow basin sediments. The lower unit, unit 1, is an alternation of shales and siltstones with a few limestone beds and decimetreto metre-thick lava flows or sills injected into semiconsolidated sediments, at a shallow level of basin infilling. Five to six main flows are recognized. This unit is named the ‘lower volcanic sequence’ (LVS). The second unit begins with a black shale deposit (graphite weight content 1.1–1.9%) averaging 1 m in thickness and indicative of a major change in the sedimentary conditions. The black shale is overlain by a metre-thick massive grey dolostone layer showing more or less intense hydrothermal recrystallization and by a thin (,10 cm) siliceous cherty bed. This triple association (black shale –grey dolostone –chert) is the regional
305
marker of the base of unit 2 and a guide for the copper mineralization (Barbanson et al. 2003). Above, unit 2 is composed of alternating shales, siltsones, limestones and metre-sized lava flows (8– 10 main flows). The intensity of the schistosity is moderate and all the sedimentary features are well preserved. Progressively, the volcanic occurrences decrease whereas the carbonated beds increase in number and size. Unit 2 can be subdivided into a lower and volcanic member (I-U2a), also named the ‘upper volcanic sequence’ (UVS) and an upper and carbonate member (I-U2b). Analyses of a metre-sized limestone bed indicate a calcareous composition with less than 2% of dolomite and variable amounts of quartz (4– 8%), clay (3– 4%) and ferrous hydroxides (5– 6%). In the northern sector of Amerdoul (Fig. 3), a more important detrital supply resulted in the deposition of metre-sized greywackes, but also of a few 10 m thick scoria debris flow showing typical slump structures of gravity-driven sliding. These observations allow us to locate the edge of
Fig. 3. Detailed geological map of the Ifri and Ouzaga area for showing the stratigraphic units. I-U1 to I-U3, Ifri units; O-U1 and O-U2-7, Ouzaga units.
306
A. POUCLET ET AL.
Fig. 4. Stratigraphic column of the Ifri and Azegour formations. LVS, lower volcanic sequence; UVS, upper volcanic sequence.
the basin to the NW. The overlying unit 3 is characterized by the vanishing of the volcanic activity and by the increasing contribution of the argillaceous component. Some microconglomerates, greywackes and sandstones form the base of this unit whereas shales dominate the upper and thicker part, the ‘Sembal sequence’.
The Ifri area: tectonics. Detailed structural investigations (Gaouzi et al. 2001; Chauvet et al. 2002; this work) indicate that the Ifri formation recorded three successive phases of folding: P1, P2 and P3, and two main faulting events. The phase P1 yielded a regional schistosity S1 subparallel to the S0 stratification, associated with interfolial isoclinal
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
micro-folds and a stretching lineation L1. Taking into account the subsequent P2 and P3 deformations, this lineation strikes SW –NE and plunges to the SW, whereas the micro-fold axes trend NW–SE (Fig. 5). A weak metamorphism caused S1-controlled recrystallization of tiny flakes of muscovite, quartz and calcite in the felsic rocks. In the mafic lavas, plagioclase is replaced by albite, pyroxene and amphibole are replaced by actinolite and chlorite, and the groundmass by an association of chlorite, actinolite, epidote, albite, quartz and calcite. This metamorphism took place in the upper greenschist facies. The structural features can be explained by tilting to the SW of the rock pile and by a SW –NE compression at an upper crustal level, under ductile conditions. The phase P2 is characterized by mesoscopic ESE-verging knee-folds whose axes trend NNE– SSW and plunge at 108 to the SSW (Fig. 5). A widely spaced crenulation cleavage S2 developed along the axial planes, which caused an intersection lineation L2 with S0 – 1. A NNE –SSW faulting commonly appeared along the fold hinges with fault planes steeply plunging to the west. In the Amerdoul zone, folding is more intense and the metre-sized folds are almost isoclinal and inclined to the east. The asymmetric geometry of the slightly overturned folds, in the Ifri zone, is emphasized in the WNW –ESE cross-section of Figure 5. This deformation is due to a WNW–ESE-oriented compressional stress field and occurred under conditions of lower ductility than for the P1 phase. It may be explained by rising of the rock pile, combined with a moderate clockwise rotation of the P1 stress field direction. An alternative explanation could be an anti-clockwise rotation of the rock pile during its rise. This rotation is consistent with the dextral pull-apart structure of the Cambrian basin in the model of Corne´e et al. (1987a) documented by the late Variscan right-lateral motion along the Tizi-n-Test Fault (Proust et al. 1977; Houari & Hoepffner 2003). In this assumption, the P1 stress field direction could have been WNW– ESE, as for the P2 direction. The phase P3 is mainly developed in the southern area. It is related to the Ouzaga thrust (Fig. 5). It corresponds to the reverse folding of the Sembal sequence and to a tectonic reworking of the more incompetent rock units (shales and pelitic shales) at the contact of these rocks with the competent ones (sandstones and dololimestones). This is particularly the case for the black shale –grey dolostone contact, at the base of the unit 2. The thrust ramps (S3 planes) weakly dip to the SSE and exhibit a discrete stretching lineation L3 plunging at 10 –208, with an average orientation of 1508. During this tectonic phase, the Sembal fault system registered a reverse movement.
307
This phase is due to a SSE– NNW major compression that mainly affected the Ouzaga –Tizzirt formation, at a supracrustal level, as shown in SSE –NNW cross-section of Figure 5. The P3 event was associated with intrusions of dioritic to granitic plutons (Fig. 2), particularly the huge complex of Tichka dated at 291 + 5 Ma (Gasquet et al. 1992). The Tichka bodies were emplaced under the control of a NW–SE shortening (Lagarde & Roddaz 1983), which was a continuation of the P3 compression. The same tectonic context prevailed in the Azegour granite emplacement (Prost et al. 1989) dated at 271 + 3 Ma (Mrini et al. 1992). The composition of the Ifri– Azegour granitoids (Ouazzani 2000) and of the Tichka complex (Gasquet 1991; Gasquet et al. 1992) is calc-alkaline with a late orogenic magmatic signature. According to the ages of the granite intrusions, the P3 event occurred in the Late Carboniferous. The copper mineralization of Ifri, also controlled by the P3 tectonic phase, has a minimum age of 270 Ma (Chauvet et al. 2002) and can be related to the regional granitic intrusions. After the three major compressional events, at least two extensional tectonic phases can be distinguished and are dated to the Early Jurassic and the Middle Jurassic, respectively. The first phase is characterized by the intrusion of numerous dykes and laccoliths of dolerites and microgabbros that cross-cut all the Palaeozoic formations in the Ifri– Azegour, Ouzaga –Tizzirt, Tanout and Talmakant areas. They show the compositional features of tholeiitic to transitional dolerites and the chemical signature of continental tholeiites (Ouazzani 2000). The dykes strike N108–N508 throughout the area, and to N808 in the Ouzaga area. They indicate a NW–SE extension that may be attributed to the Atlas rifting, which occurred in Late Triassic to Early Jurassic time, contemporaneously with the central Atlantic rifting (Laville & Pique´ 1991; Pique´ & Laville 1995). This extension is documented by the post-Permian left-lateral motion of the Tizi-n’Test Fault (Proust et al. 1977). The second extensional phase is characterized by the emplacement of barite veins, which cross-cut the dolerite dykes. In the Ifri area, these veins strike WNW – ESE. Their emplacement is associated with a north –south or NW– SE extensional stress that reworked the SW –NE faults (e.g. Sembal fault system) with a normal motion (Fig. 5). Such a context prevailed in the Middle to Late Jurassic and is responsible, in the central High Atlas, for the emplacement of alkaline to transitional magmatic rocks (Pique´ et al. 1998). The Azegour area: lithology. The Azegour area exhibits two stratigraphic units, a lower volcanic
308
A. POUCLET ET AL.
Fig. 5. Structural map of the Ifri and Ouzaga area. P2 and P3, tectonic phases. CS1, WNW– ESE cross-section in the Ifri formation to show the P2 folding. CS2, SSE–NNW cross-section in the Ouzaga and Ifri formations to show the P3 folding.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
and carbonate unit (A-U1) and an upper pelitic unit (A-U2) (Figs 4 and 6). The thickness of the recovered sequence averages 1000 m. The lower part of unit 1 is an alternation of silty shales and lava flows or sills showing the same features as those of the Ifri area. This sequence, named V1, averages 100 m in thickness. It is overlain by a
309
100 m thick pile of limestone beds, C1. Renewal of the volcanic activity yielded a second volcanic sequence, V2, which consists of interbedded lava flows, epiclastites and silty shales, reaching 200 m in thickness. In the western Tizgui area (Fig. 6), the upper part of the V2 sequence includes a 30 –50 m thick protrusion of dacite, intercalated
Fig. 6. Geological map of the Azegour area and WSW– ENE cross-section to show the P2 folding.
310
A. POUCLET ET AL.
with epiclastite beds and overlain by a fine ignimbritic layer. The V2 phase ended with deposition of impure calcarenites mixed with chlorite layers and pods, then the carbonate deposition prevailed as a 300 m thick pile of massive limestone C2. No structures or biological remains were observed in the C1 and C2 limestones. The unit 2 is characterized by a monotonous deposition of shale with some interbedded silty shale and fine sandstone layers. Its thickness exceeds 400 m. The base of this unit includes a few intercalated lava flows indicating a late and third volcanic activity, V3. There are some resemblances between the stratigraphic logs of Ifri and Azegour, particularly concerning the volcanic products (Ouazzani et al. 2001) and the two successive occurrences of limestone deposits (I-U1 and A-U1 C1; I-U2 and A-U1 C2), all being overlain by a thick pile of sandstones, siltstones and shales, although the Ifri limestones are very less massive and are interbedded with more abundant detrital sediments. This is explained by a lateral lithological variation. To the south of the Cretaceous graben, the Azegour formation shows a rapid decrease in the thickness of the limestone beds and an increase in the thickness of the siltstones and shales. The overall lithology indicates a setting in a epicontinental marine basin that deepens to the south. The Azegour area: tectonics. The Azegour formations record two main phases of deformation. The first phase yielded a cleavage schistosity S1 subparallel to the stratification S0, associated with interfolial micro-folds and a stretching lineation L1. The second phase produced mesoscopic ESEverging folds, with axes striking NNE –SSW. The cross-section of Figure 6 shows the style of this folding. These two deformations are clearly related to the P1 and P2 phases of the Ifri area. A third phase caused a weak strain-slip cleavage striking WSW–ENE and could be a distant effect of the P3 phase. As in the Ifri area, many postfolding bodies of diorites and granites intrude the Azegour formations. The largest one, the Azegour granitoid complex dated at 271+3 Ma, is located in the southern border of the Palaeozoic terrane of Azegour, and extends to the south below the sediments of the Cretaceous graben (Fig. 2). It generated a prominent contact metamorphism in the Azegour rocks, which are transformed into hornfels and skarns enriched with a molybdenum mineralization. Extending from this pluton, numerous granitic to rhyolitic dykes cross-cut the Azegour strata, and trend north– south to NNW –SSE (Fig. 6). Compared with the Ifri area, the Mesozoic doleritic dykes are scarce, and the barite veins are poorly known.
The Ouzaga –Tizzirt formation The Ouzaga –Tizzirt formation is located between the Middle Western High Atlas Fault and the Tizi-n’Test Fault, to the NW of the Tichka igneous massif. The sedimentary pile shows, from south to north, an antiformal part dominated by the syntectonic setting of the Tichka pluton (Lagarde & Roddaz 1983), a middle synformal part with an ESE –WNW-striking axis, and a northern antiformal part (Fig. 2). We investigated only the northern part. Lithology. The oldest sediments crop out in the northernmost area, ENE of Ouzaga. The youngest sediments occur in the syncline structure, at the northwestern border of the Tichka intrusive complex (Fig. 2). In the study area, we distinguish seven stratigraphic units on the basis of the deposit gradation and/or contrasting lithology (Figs 3 and 7). The total sedimentary record reaches 3100 m. The lowest unit, unit 1 (300 m), begins with detrital deposits of sandstones, conglomerates and siltstones. Then, volcanic activity resulted in numerous metre-sized basaltic flows intercalated with hyaloclastites, epiclastites and clayey siltstones. Some flows exhibit a pillow lava structure with typical quenched margins and interpillow cavities filled with formerly palagonitized glass. The pillows average 30–60 cm in diameter. The sequence includes two main decametre-sized volcanic conglomerates. In the upper part, the number of lava flows decreases, whereas silty shales and shales become thicker. Unit 2 (500 m) is characterized by a lack of volcanic deposits and the deposition of limestones. These last layers include some isolated dome-shaped stromatolites averaging 10 –20 cm in diameter. They are associated with shales, fine sandstones and calcirudites. The limestones disappear in the upper part of the unit, which consists of an alternation of shales, cherts, sandstones and arkosic sandstones. A thick conglomerate initiates unit 3 (750 m). It is overlain by few metre-sized sequences of conglomerates, arkoses, sandstones and shales. The slump structures, the sole casts of the detrital strata, and the subsequent sedimentation belong to typical Bouma sequences of turbidites; after this, the sedimentation conditions become rapidly quieer. Then, most of the unit consists of an alternation of shales, siltstones and fine sandstones, with episodic centimetre-sized layers of limestones and cherts. Unit 4 (800 m) is dominated by a new turbiditic deposition of slump-structured metre-sized sequences. Fining- and thinning-upward deposition ends with a limestone bed. Again, two turbiditic sequences form the main part of the unit. The
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
311
Fig. 7. Stratigraphic column of the Ouzaga– Tizzirt formation.
upper part is limited to siltstones, shales and few limestones. Unit 5 (250 m) is distinguished by the presence of new volcanic activity, represented by spilitic flows, hyaloclastites, shales and black cherts. Unit 6 (200 m) is a pile of brown dololimestone. This is overlain by a .300 m thick deposit of black shales, unit 7, which ended the sedimentary record in the synformal structure. Compared with the Ifri –Azegour formation, the Ouzaga –Tizzirt formation begins with a similar volcanic activity, but the lava flows exhibit pillow features and are interbedded with numerous conglomerates. This indicates a steeper substratum. Above, the lithology is an alternation of basinal detrital and calcareous sediments and of turbiditic sequences probably due to episodic tectonic activity of a cliff. The setting is the continental slope of a deep basin. Tectonics. The Ouzaga –Tizzirt formation mainly recorded the P3 event. Axes of the P3 folds strike WSW–ENE (Fig. 5). The deformation yielded decametre-sized and NNW-verging inclined folds dipping 60 –708 to the ESE. The fold angle is less than 208. In the Ouzaga area, the northern border fault (the Middle Western High Atlas Fault) is divided into two thrusts plunging 608 to the SSE (Fig. 3). Several granitic intrusions, dykes and apophyses cross-cut the sedimentary strata (Figs 2 and 3). They caused a thermal effect and local static recrystallization of biotite. The intrusion of the Tichka complex generated a strongly dipping
foliation and a large thermal metamorphism with hornfelses, spotted slates and marbles at its contact. Below the sedimentary pile, there are igneous bodies, because transformation of limestones into marbles is common in the mountainous area of Moulay Ali and Mtdadene. As in the Ifri area, Mesozoic dykes of dolerite are common.
Regional interpretation of the tectonic phases The three major compressional events, P1, P2 and P3, are related to the Variscan orogeny. The inner zone of the Atlas Variscan chain (central western High Atlas) recorded a first and important compressional effect in the Late Devonian (Hoepffner et al. 2005). Subsequently, all the western High Atlas underwent a WNW –ESE major compressional event responsible for the NNE–SSW-trending structural patterns. This event took place in the Late Vise´an (Corne´e et al. 1987b; Hoepffner et al. 2005), around 330 Ma (Essaifi et al. 2003). These data are consistent with the P1 and P2 deformations of the Ifri–Azegour formation. The sense of shear is east-vergent, as in the western Anti-Atlas (Belfoul et al. 2001), whereas it is west-vergent in the Jebilet and Rehmna areas, north of the Atlas (Hoepffner et al. 2005). The easterly overthrusting P2 folds can be observed in the formation of Tanout and in the formation of Talmakant (Fig. 2), which we date to the Early Cambrian by its Archaeocyatha fauna. The same stress direction caused the reverse faulting of the formation of
312
A. POUCLET ET AL.
Fig. 8. Structural sketch of the Variscan belt of Morocco after Houari & Hoepffner (2003), modified in the Western High Atlas region. MWHAF, Middle Western High Atlas Fault; TNTF, Tizi-n’Test Fault; SAF, South Atlas Fault; APDTZ, Atlas Palaeozoic Dextral Transform Fault. P2, Carboniferous transpressional convergence along the North African plate margin; open arrows, P2 associated fold axes; P3, Late Carboniferous–Early Permian compression linked to the clockwise motion of the West African craton; filled arrows, P3 associated fold axes.
Talmakant over that of Tanout. The P2 compressional stress affected all the Early to Middle Palaeozoic formations of the western and northwestern part of the western High Atlas. These formations are unconformably overlain by the weakly folded late Carboniferous formations and by the unfolded Permo-Triassic continental detrital deposits. A major consequence of the P2 compression is an important right-lateral strike-slip movement along the southern border of the High Atlas, the Atlas Palaeozoic dextral transform zone of Houari & Hoepffner (2003). The NNW-verging P3 tectonic event affected the southern part of the Ifri formation. It mainly involved the Ouzaga–Tizzirt formation and the area south of the Middle Western High Atlas Fault (Fig. 2). The P3 event is dated to the Late Carboniferous by the granitic intrusions. According to the structural sketch proposed by Houari & Hoepffner (2003) for the late Carboniferous evolution of the Variscan belt of Morocco, the WSW–ENE-trending fold belt (P3 deformation) is limited to the southern side of the Tizi-n’Test Fault, which is a part of the Atlas Palaeozoic dextral transform zone (APDTZ). In this study, we show that the P3 deformation prevailed in the southern part of the western High Atlas. Consequently, the APDTZ has to be located along the Middle Western High Atlas Fault. This fault registered the right-lateral motion of the APDTZ during the P2 phase (Fig. 2), before becoming a thrust as a result of the P3 compression. In
Figure 8 we represent the study area in the Variscan belt of Morocco, with regard to the P2 and P3 tectonic events, in the structural sketch of Houari & Hoepffner (2003), modified by the location of the APDTZ along the Middle Western High Atlas Fault. The P2 phase is due to the Carboniferous transpressional convergence along the North African plate margin. The P3 phase is a Late Carboniferous– Early Permian compression linked to the clockwise motion of the West African craton (Houari & Hoepffner 2003; Hoepffner et al. 2005, 2006).
Geochronology The chronostratigraphic age of the Ifri–Azegour formations cannot be well established. Because the central part of the westernmost High Atlas is overlain by Ordovician quartzites, this area was attributed to the Middle Cambrian by Corne´e (1989), in the absence of paleontological records. The Cambrian age of the Ouzaga– Tizzirt formation is based on rare Archaeocyatha remains in the southern part (Schaer 1964) and on a lithological correlation with the Ida ou Zal terrane, to the west, where Early and Middle Cambrian fossils have been described (De Koning 1957). Except for the stromatolites of unit 2, no fossils were found. A search for microfossils in the finer sediments was negative.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
We performed a U –Th– Pb analytical programme on zircons. Many mafic lavas of the Ifri, Azegour and Ouzaga area were investigated. All these lavas are devoid of analysable zircons. Zircons were finally extracted from two rocks: a dacite of unit 1 of the Azegour area, in the village of Tizgui, and a greywacke of unit 2 of the Ifri area, in the Amerdoul sector. The analytical results are listed in Table 1 and plotted in Figure 9. Concordia calculations were made using the ISOPLOT/EX program of Ludwig (2003).
The dacite of Azegour (Tizgui) The dacite is a microlitic and porphyritic lava. Zircons are rare, very fine, prismatic and clear. They belong to the S18 –S24 types of Pupin (1980), high-temperature zircons of calc-alkaline volcanic rocks. Seventeen grains have been analysed (Table 1). The data points are concordant to subconcordant (Fig. 9a). Two discordia yield intercepts at 532.5 + 4.2 and 529 + 15 Ma by the normal and inverse methods, respectively. The mean 206Pb/238U and 207Pb/206Pb ages are 532 + 12 and 537+9 Ma. All these data indicate a Nemakit –Daldynian to Tommotian age, according to the Early Palaeozoic chronology of Tucker & McKerrow (1995), and they date the upper part of the V2 volcanic phase (Fig. 4). In the central Anti-Atlas, a volcanic activity dated at 531+5 Ma (Gasquet et al. 2005) seems to be contemporaneous with the Azegour V2 phase. It occurred during the deposition of the upper part of the Tamjout dolostone, a noteworthy lithological marker in the southern High Atlas and in the Anti-Atlas. This dolostone conformably overlies the Adoudounian detrital basal formation and a thick pile of continental tholeiites in the Agoundis – Ounein and Toubkal areas of the southern High Atlas (Pouclet et al. 2007). These lavas overlie, with angular unconformity, the Ediacaran rhyolitic complex, of which the youngest lavas are dated at 543+9 Ma (Gasquet et al. 2005). Their ages are bracketed in very late Ediacaran to Early Cambrian time. The Azegour V1 phase (and the equivalent Ifri activity) could be contemporaneous with the southern High Atlas tholeiites, or slightly younger (Nemakit–Daldynian), because, in the Azegour and Ifri areas, there is no equivalent formation to the Adoudounian basal formation. This is consistent with the Nemakit–Daldynian to Tommotian dating of the V2 phase.
313
populations of detrital zircons. The first group consists of coarse, prismatic corroded and dark brown grains strongly zoned with darker inherited cores. These features are characteristic of granitic zircons. Five grains were analysed. They are moderately to strongly discordant (Fig. 9c), which indicates an origin from igneous rocks that underwent thermal disturbance. A discordia gives an upper intercept at 2075 + 290 Ma. This imprecise result is due to the varying nature of the host-rocks of the zircons, which were not cogenetic. However, a Palaeoproterozoic age can be admitted. The second group of zircons is more frequent. The grains are medium- to fine-sized, prismatic and light brown to clear, without zoning. They may have originated from acid volcanic rocks. A total of 13 grains were analysed. They are concordant to subconcordant (Fig. 9d), with 207Pb/235U ages ranging from 682 to 439 Ma (Table 1). No histogram can be made and no discordia is available. Averages and standard deviations are 546 + 77 Ma and 596 + 127 Ma for the 207Pb/235U and 207 Pb/206Pb age calculations, respectively. One may conclude that these zircons came from Late Neoproterozoic to very Early Cambrian rocks. According to the petrographical features of the greywacke and the characteristics of zircons, the parent rocks were Palaeoproterozoic granitic rocks and Neoproterozoic volcanic formations. A Neoproterozoic volcanic substratum is suspected in the Cambrian basin of Ida ou Zal (De Koning 1957; Corne´e 1989), 15 km to the SW. The 598 Ma granitic outcrop of Wirgane is 70 km to the ENE (Fig. 2). The Late Neoproterozoic rhyolitic basement below the Late Ediacaran to Cambrian basalts of the Adoudounian basal formation is 80 km to the east (Pouclet et al. 2007). The nearest Palaeoproterozoic granitic outcrop that belongs to the West African craton border is more than 100 km to the SSE, in the Western Anti-Atlas. It is possible that Neoproterozoic and Palaeoproterozoic rocks were close to the Ifri basin before the Variscan orogeny, and probably occur below the Ifri formation.
Petrography and geochemistry The volcanic rocks of the Ifri –Azegour and Ouzaga –Tizzirt formations were investigated for their petrographic and chemical compositions (Table 2).
The greywacke of Ifri
Compositions of the volcanic rocks
The greywacke is an arkosic sandstone rich in angular lithic fragments of granitic rocks and in subhedral quartz crystals. We separated two
The volcanic rocks are located in three areas: the Ifri and Azegour areas, suspected to belong to the same stratigraphic formation, and the Ouzaga
Table 1. Ion microprobe U– Pb analytical data Pb (ppm)
U (ppm)
Th (ppm)
204
314
Grain spot
Pb/ Pb
Radiogenic ratios
206
207
Pb/ Pb
206
207
Pb/ U
235
+
206
Pb/ U
238
+
Correlation error
206
Pb/ U
238
+
207
Pb/ U
235
+
207
Pb/ Pb
206
+
146.2 15.8
0.000050 0.000130 0.000008 0.000153 0.001190
0.1330 0.1224 0.1282 0.1363 0.1239
0.0591 0.0086 0.0002 0.0006 0.0024
3.166 5.282 6.098 6.346 4.335
0.311 0.147 0.201 0.446 0.635
0.1726 0.3130 0.3448 0.3407 0.2790
0.0136 0.0083 0.0034 0.0043 0.0048
0.898 0.951 0.995 0.996 0.989
1026 1756 1910 1890 1586
74 41 16 20 24
1449 1866 1990 2025 1700
73 24 28 60 208
2138 1991 2074 2165 1844
103 14 3 9 41
28.9 88.7 308.3 55.5
0.000166 0.000030 0.000060 0.000014 0.000805 0.000875 0.000503 0.000380 0.000038 0.000350 0.000427 0.000400 0.001390
0.0628 0.0610 0.0602 0.0636 0.0631 0.0634 0.0613 0.0629 0.0665 0.0630 0.0635 0.0640 0.0791
0.0004 0.0002 0.0004 0.0001 0.0004 0.0005 0.0002 0.0003 0.0009 0.0002 0.0003 0.0006 0.0003
0.805 0.661 0.646 0.916 0.541 0.589 0.649 0.659 0.627 0.782 0.868 0.622 0.958
0.024 0.018 0.018 0.024 0.051 0.068 0.036 0.045 0.084 0.039 0.072 0.092 0.075
0.0967 0.0791 0.0789 0.1046 0.0713 0.0783 0.0836 0.0808 0.0687 0.0953 0.1063 0.0752 0.1067
0.0026 0.0021 0.0022 0.0028 0.0009 0.0005 0.0005 0.0008 0.0004 0.0008 0.0024 0.0021 0.0022
0.883 0.991 0.970 0.995 0.989 0.996 0.988 0.997 0.995 0.991 0.969 0.998 0.983
595 491 489 641 444 486 518 501 428 587 651 467 653
15 13 13 16 5 3 3 5 3 5 14 13 13
600 515 506 660 439 470 508 514 494 587 635 491 682
14 11 11 13 33 43 22 27 113 22 38 56 38
619 624 580 725 411 392 463 571 813 586 576 605 778
27 7 15 5 27 37 14 17 53 15 20 32 42
116.1 180.6 127.2 120.5 95.5 48.2 56.9 35.1 26.84 37.86 37.96 25.81 32.36 28.36 31.26 17.33 21.68
0.000030 0.000060 0.000010 0.000010 0.000010 0.000100 0.000010 0.000110 0.000090 0.000043 0.000027 0.000069 0.000076 0.000057 0.000110 0.000028 0.000055
0.0578 0.0584 0.0582 0.0578 0.0576 0.0580 0.0577 0.0585 0.0590 0.0584 0.0583 0.0587 0.0584 0.0581 0.0581 0.0585 0.0579
0.0001 0.0003 0.0001 0.0002 0.0001 0.0004 0.0002 0.0006 0.0005 0.0003 0.0002 0.0004 0.0003 0.0003 0.0003 0.0006 0.0002
0.721 0.678 0.709 0.713 0.700 0.735 0.694 0.720 0.665 0.690 0.692 0.707 0.657 0.654 0.652 0.676 0.658
0.024 0.024 0.024 0.024 0.026 0.025 0.025 0.025 0.014 0.013 0.011 0.013 0.011 0.011 0.011 0.018 0.010
0.0905 0.0841 0.0884 0.0896 0.0882 0.0920 0.0872 0.0893 0.0818 0.0858 0.0861 0.0873 0.0816 0.0817 0.0813 0.0838 0.0824
0.0030 0.0029 0.0030 0.0030 0.0032 0.0031 0.0031 0.0030 0.0010 0.0012 0.0011 0.0011 0.0010 0.0010 0.0011 0.0014 0.0010
0.997 0.991 0.998 0.996 0.998 0.981 0.996 0.960 0.991 0.990 0.998 0.987 0.988 0.992 0.995 0.997 0.987
558 521 546 553 545 567 539 551 507 530 532 540 506 506 504 519 510
18 17 18 18 19 18 18 18 6 7 6 7 6 6 6 9 6
551 526 544 547 539 560 535 551 518 533 534 543 513 511 509 524 513
14 14 14 14 15 15 15 15 8 8 6 8 7 7 7 11 6
521 546 537 520 514 528 519 547 566 544 540 556 544 532 535 549 526
5 10 4 6 5 14 7 20 14 8 5 10 8 8 8 16 7
AM, detrital zircons of the Ifri greywacke; Gr.1, Palaeoproterozoic zircons: Gr. 2, Neoproterozoic zircons. T and TIZ, magmatic zircons of the Tizgui dacite (Azegour area). Analyses were carried out with the CAMECA 1270-SIMS of the centre de Recherches Pe´trographiques et Ge´ochimiques of Nancy. Calculation method is according to Deloule et al. (2001). Absolute errors are 2s. Pb, U and Th values were not determined for some samples.
A. POUCLET ET AL.
Ifri greywacke AM-Gr1 AM1 57.1 385.2 AM18 4.3 15.8 AM1-8 AM1-13 AM1-12 AM-Gr2 AM17 1.9 22.6 AM19 26.9 396.0 AM20 26.4 388.9 AM25 13.4 148.9 AM1-1 AM1-2 AM1-3 AM1-4 AM1-5 AM1-6 AM1-7 AM1-10 AM1-11 Azegour (Tizgui) dacite T1 28.6 367.9 T2 31.2 431.8 T3 27.5 362.3 T4 29.6 384.8 T6 24.6 325.1 T7 15.8 200.2 T8 14.5 193 T10 6.9 90.1 TIZ-07 18.42 262 TIZ-08 16.23 220.2 TIZ-10 31.05 419.7 TIZ-12 8.57 114.2 TIZ-17 12.56 179.2 TIZ-20 10.48 149.3 TIZ-21 6.98 99.96 TIZ-23 20.86 289.7 TIZ-28 12.43 175.6
+
Ages (Ma)
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
(b) 0.102
Tizgui dacite (n 17)
0.098
Mean206Pb/238U age 532 ± 12 560
0.086
0.0595 570
0.0585
520
Intercept at 529 ± 15 Ma MSWD = 4.7
0.0605
207
540
550
500
0.074 0.58
0.62
0.66
0.70
207
0.74
0.78
0.0565 10.4
0.82
0.32 0.28
2100
1700
206
1300
0.20
1100
Intercepts at 199 ± 1400 & 2075 ± 290 Ma MSWD = 2.1
0.16 0.12 0
1
2
3
11.2
4
11.6
12.0
5
0.12 0.11
1900
1500
0.24
10.8
12.4
238U/206Pb
(d)
0.36
510
Pb/235U
0.40 Ifri greywacke Group 1 old detrital zircons (n 5)
530
0.0575
Intercepts at –386 ± 260 & 532.5 ± 4.2 Ma MSWD = 0.91
0.078 480
206Pb/238U
580
0.090
0.082
(c)
600
Pb/238U
0.094
Tizgui dacite (n 17)
0.0615
Pb/238U
206Pb/238U
(a)
315
6
7
8
207Pb/235U
Ifri greywacke Group 2 young detrital zircons 640 (n 13)
0.10
720 680
600 560
0.09
520
0.08 0.07
480 440 400
0.06 0.3
0.5
0.7
0.9
1.1
207Pb/235U
Fig. 9. (a, b) U–Pb concordia diagrams for magmatic zircons from the dacite of Tizgui of the Azegour formation: (a) conventional plot; (b) Tera –Wasserburg plot. (c, d) U –Pb concordia diagrams for detrital zircons from the Ifri greywacke: (c) Palaeoproterozoic zircons; (d) Neoproterozoic zircons. Data point error ellipses are 2s.
area. Previous data are available from Ouazzani et al. (1998, 2001) and Ouazzani (2000). The Ifri lavas. The Ifri area exhibits two volcanic sequences, LVS and UVS in units 1 and 2, respectively (Fig. 4). Both sequences have the same lithological and petrographical features of underwater basaltic flow or sills. However, the spilitization effect is more important in the LVS. The rocks are massive or, sometimes, vesicular with fine amygdales of calcite. The texture is hyalomicrolitic and more or less porphyritic with seriate microphenocrysts of plagioclase (10 –30%) and subordinate amphiboles and clinopyroxenes replaced by secondary amphiboles. A static epizonal metamorphism caused replacement of the groundmass glass and minerals by an association of ripidolite, clinozoisite, actinolite, albite, muscovite, calcite, leucoxene, Fe-hydroxides and quartz. Feldspars are partly albitized. Their preserved cores range from An66 to An50. Ferro-magnesian
minerals are recrystallized to actinolite (Mg/ (Mg þ Fe2þ) ¼ 0.65–0.72; Si 7.6–7.7 a.p.f.u.) and Fe-oxides. Pseudomorphs of clinopyroxenes and amphiboles are observed, and Mg-hornblende was analysed (Mg/(Mg þ Fe2þ) ¼ 0.68 –0.72; Si 6.5 –7.1 a.p.f.u.; TiO2 0.2 –0.6%). Magmatic oxides consist of microcrysts of Ti-magnetite (TiO2 12 –16%) and ilmenite. No pseudomorphs of olivine were seen. Tectonic cleavages do not significantly modify the texture. Local increase of stress along hinges or shear zones is responsible for foliation-guided crystallization of chlorite and muscovite. The vicinity of granitic intrusions is revealed by the appearance of static biotite (Mg/ (Mg þ Fe2þ) ¼ 0.55–0.79; TiO2 1.1 –1.9%). A total of 25 lava flows were analysed, from outcrops and from four bore-holes of the Compagnie Minie`re de Seksaoua. Because of the contamination by neighbouring lodes and by granites, some analyses were discarded. Seventeen analyses remain: five for the lower volcanic sequence (LVS) and 12
316
A. POUCLET ET AL.
Table 2. Geochemical analyses of the Cambrian volcanic rocks of the western block of the Western High Atlas, Ifri –Azegour and Ouzaga – Tizzirt areas Location:
Ifri LVS
Ifri UVS
Sample no:
SKD4k
SKD49f
SKD49e
IF-284
SKD4m
SKD-25k
SKD25o
AM-8
AM355
SKD2c
SKD2a
AM319
IT-8
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2 O P2O5 LOI Total V Cr Ni Ga Rb Sr Y Zr Nb Ba Hf Ta Pb Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
47.43 1.90 16.86 11.30 0.16 6.92 7.53 3.57 0.75 0.32 3.14 99.88 249.7 248.73 94.1 20.87 24.46 412.2 34.09 181 6.43 155 4.11 0.48 8.80 0.80 0.31 12.12 31.39 4.42 20.75 5.41 1.89 5.82 0.93 5.71 1.15 3.07 0.48 2.99 0.42
50.76 1.11 19.41 8.17 0.08 5.42 6.62 5.62 0.43 0.19 2.07 99.88 162.6 107.40 78.5 19.85 6.40 465.0 18.63 90 5.21 198 2.47 0.39 9.25 1.69 1.27 13.07 29.45 4.16 18.91 4.37 1.51 3.94 0.58 3.34 0.64 1.83 0.27 1.81 0.28
51.18 1.28 15.06 9.27 0.14 5.51 6.17 4.31 0.06 0.20 6.85 100.03 198.3 108.60 50.2 19.98 4.31 226.0 37.66 208 4.66 98 5.18 0.38 8.01 1.39 0.74 13.09 31.24 4.34 19.60 5.22 1.52 5.86 0.99 6.24 1.28 3.74 0.57 3.91 0.60
55.52 1.02 17.72 8.47 0.06 4.74 2.24 5.28 0.66 0.32 3.61 99.64 158.7 47.09 43.0 22.32 23.20 415.1 11.68 155 13.02 231 3.71 0.75 3.25 1.63 0.61 13.78 31.07 3.95 16.79 3.50 0.96 2.44 0.39 1.98 0.40 1.17 0.19 1.17 0.19
58.32 1.81 15.99 8.44 0.09 3.38 1.97 6.58 1.56 0.35 1.87 100.36 85.0 7.61 18.0 20.02 66.73 213.0 40.90 237 11.30 234 5.69 0.89 1.89 3.53 1.52 25.62 59.68 8.59 33.18 8.55 2.48 7.87 1.16 6.87 1.29 3.60 0.56 3.52 0.52
44.52 1.60 17.63 15.02 0.14 5.80 2.63 2.48 1.50 0.41 8.72 100.45 159.0 104.00 67.6 15.70 65.10 62.3 31.00 172 9.64 89 3.18 0.67 4.89 0.74 0.36 19.00 41.30 4.99 19.60 4.54 1.51 4.93 0.82 5.10 1.10 3.01 0.47 2.70 0.40
44.81 1.54 17.09 9.89 0.16 6.62 8.47 3.46 1.33 0.41 6.31 100.09 152.0 60.00 64.3 17.50 31.00 536.0 23.70 167 9.66 274 3.39 0.71 8.38 1.25 0.38 17.60 38.90 5.14 19.90 4.33 1.45 4.31 0.66 4.01 0.87 2.24 0.35 2.20 0.35
45.24 2.44 15.28 12.91 0.09 9.24 3.62 3.54 0.17 0.42 7.03 99.98 265.0 274.00 122.0 20.20 10.14 78.0 39.90 212 6.19 67 5.10 0.49 4.56 0.45 0.20 13.30 34.19 5.21 22.95 7.39 2.24 7.48 1.26 6.88 1.43 3.71 0.54 3.74 0.51
46.02 2.45 16.03 14.07 0.07 9.13 2.71 3.85 0.20 0.29 5.01 99.83 283.9 288.03 111.0 19.61 6.64 53.2 36.94 177 4.63 50 4.55 0.37 1.00 0.37 0.15 8.70 23.09 3.76 19.31 5.12 1.93 6.44 0.98 6.25 1.24 3.59 0.58 3.50 0.60
46.22 2.11 17.02 10.29 0.12 4.93 5.88 4.87 4.14 0.83 3.79 100.20 157.8 9.33 20.9 20.01 386.79 510.4 35.49 348 25.12 535 6.47 1.63 3.29 2.87 1.13 58.93 121.51 14.21 54.02 9.58 2.59 7.89 1.20 6.75 1.29 3.36 0.51 3.40 0.48
46.64 2.17 17.41 10.80 0.11 5.17 5.66 4.50 3.95 0.87 2.79 100.07 165.7 10.74 23.4 20.08 271.81 732.1 35.84 343 25.15 725 6.43 1.60 2.97 2.78 1.17 62.47 124.87 14.40 55.41 9.62 2.96 7.75 1.26 6.94 1.27 3.56 0.56 3.45 0.48
48.80 2.75 14.77 16.05 0.05 6.33 2.54 4.13 1.23 0.35 3.48 100.48 419.0 182.00 71.1 24.00 37.00 91.3 51.70 188 5.21 42 4.83 0.41 2.15 0.41 0.25 12.50 34.40 5.48 26.00 7.64 2.66 8.13 1.40 8.59 1.95 5.27 0.81 5.07 0.70
50.52 1.03 16.00 8.07 0.09 5.84 5.45 4.45 1.42 0.28 6.22 99.37 187.0 102.00 63.5 20.30 7.82 288.0 14.00 129 10.00 399 3.07 0.80 2.03 1.67 0.63 14.84 32.65 3.98 16.77 3.66 1.19 3.27 0.51 2.56 0.49 1.33 0.18 1.04 0.15
for the upper one (UVS). Compositional ranges are, for the LVS and UVS lavas, respectively (Table 2): SiO2 47.4 –58.3 and 44.5 –54.6%; MgO 3.4–6.9 and 2.3–9.2%; Mg-number (molar 100 Mg/ (Mg þ Fe2þ) ¼ 44.2– 56.8 and 29.4–58.6; TiO2 1.0–1.9 and 1.0–2.8%. These lavas are basic and moderately evolved, the most evolved being in the UVS. The UVS lavas are the richest in titanium. The alkali contents are highly variable, related to the spilitization effect (Na2O high) and weathering (K2O low), which induce high loss on ignition. These processes are also responsible for anomalous variations of the mobile elements, Sr, Rb and Ba. Because of the mobility of alkalis, it is not possible to use the silica v. alkalis diagram to give a
conventional name to the rocks, or to use the normative composition to determine a saturation index. The incompatible trace element patterns (Fig. 10) show a low to moderate rare earth element (REE) fractionation, (La/Yb)N ¼ 2.4– 8.4 and 1.8–13.1 for the LVS and UVS, respectively; moderate negative to absent Nb and Ta anomalies, (Nb/La)N ¼ 0.4 –1.0 and 0.3–0.7; moderate negative to no Th anomalies, (Th/La)N ¼ 0.5–1.1 and 0.3–0.9; and variable contents of Rb, Ba and Sr. These last variations could be due to the spilitization process. However, the moderate positive and negative Sr anomalies could be explained by cumulation or fractionation of plagioclases.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
Ifri UVS
317
Azegour
AM-7-13 AM-408 SKD-25n AOS-422 AA-10 AA-34 AZ-273 AA-66 AZ-321 AZ-324 AZ-329 AA-13 AA-14 AA-26a 51.25 1.81 16.63 9.96 0.06 3.96 5.79 5.65 0.98 0.45 3.35 99.89 258.0 48.80 25.7 29.50 30.54 193.0 44.80 245 7.87 135 6.11 0.57 5.58 1.80 1.30 27.83 64.96 8.78 40.53 9.50 3.77 9.19 1.30 7.59 1.59 4.41 0.65 4.29 0.59
51.61 1.74 17.47 8.71 0.09 4.11 3.06 8.15 0.06 0.90 3.95 99.85 85.9 19.28 22.0 19.62 9.65 448.5 29.12 400 34.49 125 7.70 2.46 2.74 4.18 1.57 51.60 103.90 11.96 45.27 8.13 2.35 6.80 0.98 5.57 1.06 2.92 0.42 2.79 0.44
53.96 1.50 16.26 8.19 0.08 4.29 2.72 6.84 0.00 0.34 5.60 99.78 140.0 80.60 61.7 19.40 2.64 84.5 30.20 269 8.16 37 6.25 0.65 2.42 0.82 1.63 14.00 33.10 4.56 19.90 4.51 1.51 5.34 0.78 4.93 1.04 2.94 0.47 3.15 0.53
54.63 1.77 14.05 10.92 0.11 2.30 4.54 4.35 1.54 0.25 4.34 98.80 150.3 22.17 10.5 22.01 47.06 116.4 43.08 260 8.34 515 6.41 0.72 6.33 1.73 0.74 22.54 47.17 7.11 31.50 7.63 2.72 7.73 1.23 7.51 1.51 4.20 0.62 3.97 0.61
42.36 1.85 20.06 11.38 0.25 2.09 7.69 4.74 1.65 0.30 7.41 99.78 242.0 136.00 51.6 24.40 29.90 232.0 31.80 154 5.40 420 3.89 0.41 1.57 0.92 0.32 12.73 29.47 3.96 18.85 5.01 1.73 4.94 0.84 4.66 1.29 2.95 0.44 3.10 0.43
44.96 1.35 15.26 10.09 0.16 7.13 10.10 2.17 0.92 0.24 7.46 99.84 255.0 166.00 65.2 17.50 28.90 349.0 27.10 96 2.00 136 2.35 0.16 3.94 0.25 0.12 4.40 12.50 1.97 10.16 3.33 1.25 4.09 0.67 4.19 0.90 2.42 0.40 2.55 0.39
44.97 1.26 19.85 14.32 0.29 2.06 4.09 5.66 0.64 0.36 5.97 99.47 201.8 151.60 47.7 20.75 11.84 222.4 20.97 133 5.05 154 3.37 0.39 1.77 0.86 0.36 7.50 19.16 2.67 11.93 2.95 0.66 3.33 0.53 3.18 0.75 2.05 0.37 2.24 0.37
46.03 1.11 18.62 9.49 0.12 8.55 10.67 2.12 0.43 0.26 2.49 99.89 197.4 334.10 161.2 16.65 9.80 201.3 20.44 76 1.09 62 1.56 0.08 3.98 0.11 0.04 2.75 8.58 1.36 7.72 2.15 0.90 2.92 0.42 2.99 0.71 1.78 0.28 1.79 0.27
47.63 2.33 17.74 13.68 0.20 4.42 3.79 4.89 0.00 0.28 4.84 99.80 306.9 80.86 6.9 25.00 2.00 184.2 38.58 189 5.15 29 4.74 0.39 2.96 1.05 0.58 11.14 28.38 4.31 20.29 6.16 2.08 6.89 1.06 6.81 1.41 3.85 0.58 3.93 0.60
47.99 1.45 15.88 9.52 0.13 6.63 9.93 3.31 0.29 0.20 4.52 99.85 243.4 235.65 67.8 17.22 6.57 253.2 29.93 127 3.08 106 3.03 0.23 4.01 0.79 0.35 7.32 19.02 2.83 13.14 4.00 1.21 4.52 0.74 4.94 1.04 2.89 0.44 2.87 0.45
48.97 1.36 16.53 9.58 0.22 7.37 12.72 1.94 0.32 0.21 0.65 99.87 229.9 217.40 79.9 17.77 5.63 425.9 30.28 113 3.83 85 2.66 0.25 2.78 0.68 0.33 10.00 24.14 3.50 16.91 4.69 1.85 5.09 0.82 5.36 1.10 2.90 0.44 2.91 0.42
53.60 56.10 0.95 0.62 18.40 18.80 7.55 5.81 0.07 0.08 4.40 3.45 3.28 4.89 5.74 7.02 0.83 0.23 0.21 0.20 4.64 2.58 99.67 99.78 136.0 102.0 47.40 84.20 20.1 31.0 20.50 20.40 13.30 3.00 537.0 616.0 17.00 9.60 106 91 7.70 5.80 352 151 3.04 2.15 0.64 0.46 3.80 5.83 1.30 1.38 0.72 0.59 12.20 9.26 27.10 20.80 3.55 2.52 14.50 10.40 3.19 2.39 1.19 0.89 3.11 2.10 0.55 0.33 3.05 1.72 0.71 0.37 1.56 0.96 0.23 0.15 1.43 0.92 0.24 0.15
56.52 0.64 18.65 6.12 0.06 3.53 5.30 5.25 0.52 0.18 3.11 99.88 100.0 71.30 32.0 19.40 8.27 724.0 10.60 93 6.22 147 2.32 0.49 1.45 1.20 0.63 9.19 21.50 2.78 10.70 2.45 0.76 2.12 0.32 1.87 0.40 1.01 0.14 1.11 0.17
(Continued )
There is no europium anomaly (Eu/Eu* ¼ 0.8–1.1 and 0.9–1.1 for LVS and UVS). The positive Ti anomalies may be a magmatic feature, whereas the negative Ti anomalies in the more evolved lavas could be explained by oxide fractionation. The Azegour lavas. The lavas of the Azegour area occur as three volcanic members, V1, V2 and V3 (Fig. 4). They have the same lithological features as the Ifri lavas. However, two additional petrographical facies are present: plagioclase cumulate mafic lavas in the V1 and V2 members and a dacitic lava in the upper part of the V2 member, the protrusion of Tizgui (Fig. 6). The common
mafic lavas are microlitic and porphyritic, with phenocrysts of An60 – 82 plagioclase and of amphibole Mg-hornblende (Mg/(Mg þ Fe2þ) ¼ 0.77 –0.85; Si 6.8–7.4 a.p.f.u.; TiO2 0.2–0.7%). Most, if not all, of these amphibole phenocrysts were former pyroxenes, according their shapes. The magmatic oxide consists of microcrysts of Ti-magnetite (TiO2 15 –29%) and ilmenite. The cumulate facies is characterized by high amounts of plagioclase, forming up to 63 –69% of the rock. The dacite is hyalo-microlitic porphyritic with phenocrysts of An31 – 40 plagioclase, Or86 – 89 sanidine, embayed quartz, amphibole totally transformed to actinolite, and biotite (Mg/(Mg þ Feþ2) ¼ 0.39–0.42; TiO2
318
A. POUCLET ET AL.
Table 2. Continued Location: Sample no: SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2 O P2O5 LOI Total V Cr Ni Ga Rb Sr Y Zr Nb Ba Hf Ta Pb Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Tizgui
T-18
T-14
T-6
T-2
62.82 63.25 65.40 67.80 0.83 0.64 0.55 0.51 17.17 18.34 16.60 17.30 8.27 5.86 3.74 1.59 0.02 0.04 0.05 0.02 2.92 3.09 2.22 0.89 0.38 2.08 2.32 1.58 1.77 1.00 5.51 5.86 3.85 3.60 2.32 3.14 0.20 0.22 0.17 0.14 1.06 1.42 0.69 1.10 99.29 99.54 99.57 99.93 107.0 88.0 62.0 57.0 92.80 87.40 40.40 31.20 53.5 46.4 29.8 19.6 24.60 26.30 22.70 20.70 104.50 92.50 56.90 59.50 50.0 90.0 137.0 166.0 23.50 20.50 5.00 5.30 178 173 148 118 13.00 12.50 7.80 8.40 548 1851 2655 970 4.68 4.83 3.66 3.21 1.07 1.14 0.69 0.75 1.71 5.08 7.98 79.10 9.20 12.80 2.83 2.83 2.26 2.82 1.37 0.94 39.42 42.79 16.60 15.90 81.81 81.29 34.40 31.30 9.05 9.28 4.13 3.80 33.59 35.12 16.10 13.70 6.61 6.24 2.76 2.29 1.52 1.67 0.84 0.54 4.50 4.76 2.01 1.65 0.79 0.70 0.22 0.22 4.48 3.65 1.04 1.16 0.83 0.72 0.19 0.23 2.47 2.08 0.49 0.53 0.37 0.31 0.08 0.08 2.42 2.22 0.48 0.51 0.37 0.39 0.07 0.08
Tifirt
Assif Al Mal
Iberdaten
TI-1
AS-3
IB-7
44.90 1.34 17.80 10.10 0.14 7.58 8.14 2.26 0.68 0.19 6.45 99.58 198.0 85.30 91.7 18.40 29.80 270.0 25.10 111 2.10 110 2.64 0.18 9.81 0.15 0.08 4.69 13.40 2.10 10.70 3.46 1.33 3.87 0.67 4.18 0.98 2.42 0.36 2.35 0.37
45.91 2.84 15.93 13.41 0.15 5.05 6.73 4.39 0.18 0.45 4.91 99.95 287.0 36.40 49.6 22.30 10.02 383.0 40.40 232 8.15 1139 5.79 0.63 7.26 1.31 0.60 16.76 41.46 5.51 25.00 6.93 2.48 8.41 1.29 7.37 1.70 4.21 0.62 4.26 0.67
57.30 0.99 18.30 5.68 0.10 2.20 2.41 7.41 0.87 0.21 4.42 99.89 123.0 91.60 32.0 16.80 24.10 289.0 18.90 113 4.20 169 3.05 0.35 10.20 1.32 2.02 7.33 21.40 2.77 12.50 3.20 0.81 3.10 0.49 3.01 0.72 1.70 0.32 1.82 0.30
2.8–3.8%). According to its MgO, FeO, Al2O3 and TiO2 contents, the biotite is magmatic and belongs to a calc-alkaline magma (Nachit et al. 1985; Rossi & Che`vremont 1987). The groundmass contains actinolite, albite, sanidine, quartz, Fe-oxides, ilmenite, rutile and secondary biotite. A total of 14 lavas were analysed from the three volcanic phases; 10 for the mafic lavas and four for the evolved lavas. In addition, three mafic lavas, located west of the Azegour granite, were selected, in the northernmost continuation of the Ifri area (Tifirt, Assif Al Mal, Iberdaten) (Table 2). Compositional ranges are SiO2 42.4– 67.8%, MgO 0.9– 8.6%, Mg-number ¼ 22.2 –64.1 and TiO2 0.5– 2.8%. The plagioclase cumulates are characterized by high alumina contents compared with low
Ouzaga AOS425
OZ-50
AOS427
OUZ354
AOS436
OUZ353
45.42 45.63 1.62 3.32 17.72 17.73 10.84 13.91 0.04 0.15 11.96 4.30 0.62 3.47 4.05 5.17 0.29 1.08 0.29 0.47 6.38 4.70 99.23 99.93 160.9 418.0 234.60 24.30 180.7 12.8 17.81 27.30 6.97 21.90 120.0 104.0 25.36 36.40 136 225 7.93 8.40 72 2251 2.94 5.29 0.57 0.70 1.54 8.52 0.54 0.84 0.24 0.28 11.76 15.20 25.88 33.77 3.55 4.86 15.44 23.17 3.96 6.71 1.21 2.65 4.41 7.07 0.72 1.04 4.46 6.24 0.89 1.32 2.50 3.26 0.37 0.51 2.51 3.31 0.39 0.54
45.86 1.67 16.72 10.12 0.13 7.02 7.40 3.45 1.22 0.37 5.19 99.15 158.1 116.60 111.6 17.27 28.60 482.0 26.07 152 9.01 248 3.32 0.64 3.22 0.80 0.34 16.61 35.64 4.73 19.85 4.45 1.58 4.58 0.73 4.57 0.93 2.58 0.39 2.55 0.40
48.65 1.68 19.58 11.45 0.11 6.38 0.93 6.14 0.00 0.27 4.58 99.77 205.8 95.84 69.5 20.52 1.33 117.4 26.24 128 4.52 49 3.28 0.40 4.28 1.03 0.65 12.08 31.51 4.61 20.97 5.22 1.53 5.93 0.89 5.21 0.98 2.50 0.35 2.29 0.39
48.96 1.60 15.84 9.12 0.12 4.67 5.65 4.02 2.78 0.62 5.40 98.78 123.4 97.80 60.0 19.02 149.90 338.4 30.83 302 20.73 1296 5.90 1.60 3.12 4.39 6.43 36.44 72.00 8.72 33.42 6.37 1.75 5.86 0.89 5.29 1.06 3.02 0.45 3.11 0.49
49.22 1.61 19.25 9.71 0.10 4.16 4.02 5.95 0.74 0.29 4.79 99.84 205.9 44.60 29.2 20.46 16.96 242.6 22.14 129 4.70 236 3.38 0.39 4.96 1.02 0.74 11.54 29.23 4.25 19.95 4.76 1.67 4.89 0.68 4.29 0.83 2.26 0.31 2.27 0.38
magnesia contents. The dacitic rocks have a silica content up to 60%, and show the lowest Mgnumbers and titanium contents. As in the Ifri lavas, the amount of alkalis was modified by postmagmatic processes, as well as Ba, Rb and, pro parte, Sr. The incompatible trace element patterns of the mafic lavas (Fig. 10) resemble to those of the Ifri lavas, with low to moderate REE fractionation, (La/Yb)N ¼ 1.8 –7.2. However, three more basic samples show no REE fractionation ((La/ Yb)N ¼ 1.1–1.4), but are depleted in Ta, Nb and Th, ((Nb/La)N ¼ 0.4; (Th/La)N ¼ 0.3–0.4). The common basic rocks show moderate negative Nb and Ta anomalies, (Nb/La)N ¼ 0.4–0.7, and moderate negative to positive Th anomalies,
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
Ouzaga
319
Tizzirt
Talmakant
OZ-29
OZ-298
OZ-7
OZ-10
OZ-374
OUZ-373
OUZ-379
TZ-404
TZ-404a
TM-365
TM-364
49.49 2.35 18.03 12.61 0.10 4.06 2.72 6.56 0.43 0.51 2.99 99.85 278.0 39.20 18.0 26.10 7.10 195.0 36.30 276 9.60 189 6.35 0.70 4.64 1.41 1.55 12.85 32.90 4.75 22.70 6.40 1.36 6.10 0.97 6.27 1.37 3.34 0.59 3.44 0.55
49.70 2.97 16.47 13.01 0.18 3.94 3.70 6.26 0.13 0.45 3.10 99.91 384.7 25.28 11.0 22.68 1.70 234.8 31.21 197 6.97 398 4.30 0.58 4.58 0.84 0.43 13.68 35.01 5.08 22.85 5.41 2.07 6.38 0.88 5.47 1.09 3.00 0.47 2.50 0.43
49.80 1.69 17.80 9.51 0.13 2.85 7.42 6.10 0.10 0.38 3.98 99.76 183.0 30.70 18.2 26.90 3.60 671.0 35.50 198 7.40 41 4.43 0.58 4.99 1.32 1.10 17.80 43.50 5.81 27.30 6.85 3.16 6.12 0.98 5.73 1.30 3.17 0.47 2.96 0.43
50.00 2.07 18.40 10.00 0.10 3.80 3.28 6.64 0.15 0.50 4.82 99.76 169.0 6.20 6.0 22.20 5.30 354.0 31.00 228 11.20 72 6.19 0.81 6.58 1.59 0.79 16.70 39.50 5.55 23.00 6.12 1.62 6.90 1.09 6.41 1.41 3.33 0.52 3.66 0.59
53.25 1.12 19.02 9.60 0.06 5.19 0.46 6.79 0.51 0.24 3.56 99.80 165.3 124.20 61.1 21.17 9.86 71.6 18.63 119 4.48 96 3.26 0.38 3.20 1.61 0.93 13.43 26.97 3.61 14.84 3.15 1.04 3.16 0.56 3.72 0.78 2.20 0.32 2.01 0.31
53.41 1.17 19.69 10.98 0.06 3.19 0.37 7.69 0.17 0.21 2.84 99.78 96.4 66.18 24.2 22.73 3.80 82.0 16.26 109 4.17 69 2.82 0.32 2.58 1.55 0.73 9.67 22.89 3.17 13.69 2.92 0.80 2.77 0.48 3.18 0.69 1.98 0.29 1.80 0.24
53.90 1.27 18.40 8.74 0.08 4.06 4.34 6.00 0.00 0.22 2.78 99.79 153.70 64.00 28.43 23.42 3.56 253.40 25.68 200.50 6.91 52.00 4.81 0.57 3.89 1.72 2.16 16.79 37.76 5.29 23.24 5.43 1.65 5.17 0.79 4.60 0.90 2.50 0.38 2.48 0.38
45.67 1.84 17.02 10.23 0.13 6.44 5.47 4.17 1.19 0.62 7.14 99.92 154.1 123.20 84.5 17.85 25.95 452.5 31.05 194 13.79 736 4.21 0.98 3.51 1.19 0.49 31.61 64.77 8.23 32.83 6.55 2.23 6.13 0.96 5.76 1.13 3.19 0.46 3.07 0.47
47.52 1.78 17.91 12.17 0.07 6.65 2.15 5.08 0.41 0.60 5.49 99.83 161.0 122.60 102.6 16.63 10.00 103.2 30.81 193 12.83 116 4.02 0.92 2.43 1.16 0.86 27.07 49.98 6.24 23.50 4.54 1.74 4.91 0.86 5.67 1.17 3.16 0.42 2.56 0.39
51.77 1.32 15.80 11.98 0.18 5.48 4.02 4.14 0.00 0.14 5.06 99.89 176.9 174.00 60.7 21.10 3.99 535.9 10.43 55 4.26 120 1.00 0.36 67.33 0.44 0.14 3.73 8.63 1.40 7.49 2.41 1.30 2.55 0.39 2.20 1.65 0.98 0.13 0.80 0.12
52.09 1.24 15.88 11.80 0.15 5.32 3.29 4.30 0.14 0.15 5.38 99.74 176.1 177.60 65.9 21.52 9.70 362.0 10.52 56 3.95 173 1.37 0.34 15.23 0.43 0.14 3.45 8.22 1.31 6.90 2.27 1.04 2.38 0.37 2.15 1.54 0.98 0.14 0.83 0.12
Analyses were by inductively coupled emission and mass spectrometry at the Centre de Recherches Pe´trographiques et Ge´ochimiques of Nancy, and by inductively coupled emission spectrometry at the Institut des Sciences de la Terre d’Orle´ans. Analysis errors are around 2% for major elements and 5–10% for minor elements.
(Th/La)N ¼ 0.5–1.4. The negative anomalies are limited to the less basic rocks. A positive Sr anomaly characterizes the cumulative facies. The dacites are more REE fractionated, (La/ Yb)N ¼ 11.6 –24.6. The highest ratio is due to the heavy REE (HREE) depletion in two samples, which could be explained by mineral fractionation, as also shown by Sr, Eu and Ti negative anomalies. The Ouzaga–Tizzirt lavas. The volcanic rocks of the Ouzaga –Tizzirt formation are abundant in unit 1, as lava flows and hyaloclastites. Some spilites
and hyaloclastites are also present in unit 5, in the Tizzirt sector. Volcanic rocks were sampled in the neighbouring formations of Tanout and Talmakant (Figs 2 and 3). They consist of highly altered flowsills and of a laccolith, 50 m in thickness, stratigraphically emplaced a few hundred metres below the Archaeocyatha bioherm of Talmakant. The Ouzaga lavas have a microlitic porphyritic texture rich in plagioclases, ferro-magnesian minerals and ilmenite. A quench texture characterizes the flow margins. The lavas recorded a static greenschist-facies metamorphism. All the
320
A. POUCLET ET AL.
magmatic minerals are pseudomorphosed into albite, actinolite, ripidolite, clinozoisite, calcite, and quartz. The order of crystallization shows syncrystallization of plagioclases and ferro-magnesian minerals. No pseudomorphs of olivine were identified. The Tizzirt lavas are totally spilitized. The Talmakant laccolith rock is a doleritic microgabbro with an intersertal and cumulate texture rich in plagioclase. From the analytical data, we retain 13 analyses of the Ouzaga lavas, two of the Tizzirt spilites, and two of the Talmakant laccolith. Compositional
Fig. 10. Primitive mantle normalized diagrams of the Western High Atlas Early Cambrian volcanic rocks. (a) Ifri area; (b) Azegour area; (c) Ouzaga area; (d) comparative patterns of the basic lavas of the Ifri, Azegour and Ouzaga areas. Normalization values after Sun & McDonough (1989).
Fig. 11. (a) La v. La/Yb diagram of the Western High Atlas Early Cambrian volcanic rocks, for determination of the source features and the differentiation process. Same symbols as for Figure 10. OIB, ocean island basalt; UC, upper crust; CC, crustal contamination process; FC, fractional crystallization process. Reference values after Sun & McDonough (1989) and Taylor & McLennan (1985). (b) Ta/Yb v. Th/Yb diagram of the Western High Atlas Early Cambrian volcanic rocks after Pearce (1983). Same symbols as for Figure 10, plus: PM, primitive mantle (Sun & McDonough 1989); IRT, initial rift tholeiite (Pouclet et al. 1995); CT, continental tholeiite (Holm 1985); SZ, subduction zone enrichment process; IAT, island arc tholeiite; CAB, calc-alkaline basalt; SHO, shoshonite.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
ranges are: (1) for the Ouzaga lavas: SiO2 45.4– 53.9%; MgO 2.9– 12.0%; Mg-number ¼ 36.5– 68.6; TiO2 1.1– 3.3%; (2) for the Tizzirt spilite: SiO2 45.7– 47.5%; MgO 6.4– 6.7%; Mg-number ¼ 52.0– 55.5; TiO2 1.8%; (3) for the Talmakant dolerite: SiO2 51.8–52.1%; MgO 5.3– 5.5%; Mg-number ¼ 47.2 –47.5%; TiO2 1.2– 1.3%. We note the high TiO2 content for some Ouzaga basalts and the high MgO content for a ferro-magnesian mineral-rich basalt. The alkali contents were modified during post-magmatic processes, with Na enrichment and K depletion. The incompatible trace element patterns of the Ouzaga and Tizzirt lavas (Fig. 10) are close to those of the Ifri lavas with low to moderate REE fractionation, (La/Yb)N ¼ 2.7– 8.3. The negative Ta and Nb anomalies are moderate to low, (Nb/La)N ¼ 0.3–0.7. The Th behaviour is variable, (Th/La)N ¼ 0.3–1.2. Some positive Ti anomalies are noticeable, Ti/Ti* ¼ 1.5. There are no europium anomalies. Variable contents of Ba, Rb and Sr, which are mobile elements during the metamorphic and alteration processes, are not significant. The depletion in incompatible elements of the Talmakant microgabbro is a normal feature of cumulates, caused by the dilution effect of the cumulative minerals poor in these elements. Indeed, the patterns are characterized by a weak REE fractionation, (La/Yb)N ¼ 3.0–3.3, and the lack of Nb and Th anomalies. The positive Sr and Eu anomalies are symptomatic of the plagioclase
321
cumulation, whereas the positive Ti anomaly could be due to the early crystallization of ilmenite.
Magmatological aspects and geotectonic relationships The basic lavas of the three areas, Ifri, Azegour and Ouzaga, show similar petrographical and chemical compositions. The trace element patterns (Fig. 10) are more depleted in the Azegour basic lavas, but the patterns of the Azegour intermediate lavas are similar to those of the Ifri lavas. The Ouzaga patterns are the least fractionated. Calculation of the magmatic source compositions and differentiation processes is severely handicapped by the postmagmatic variations of the alkaline, Mg, Rb, Ba, and Sr contents. The magmatic compositions resulted from various processes: magma mixing of depleted to enriched sources, partial melting of the sources, fractional crystallization and crustal contamination–assimilation. The La v. La/Yb diagram (Fig. 11a), excluding the cumulative rocks, indicates a moderately depleted to undepleted source for the basic lavas and a differentiation by fractional crystallization. Crustal contamination cannot be excluded. This is confirmed by the Ta/Yb v. Th/Yb diagram of Pearce (1983) and Pearce et al. (1990) (Fig. 11b). The magmas were generated from more or less depleted sources, possibly with slight variations in the degree
Fig. 12. Comparative primitive mantle normalized diagrams of the Western High Atlas Early Cambrian volcanic rocks with typical profiles of basalts from different geotectonic settings. N-MORB and OIB (ocean island basalt) after Sun & McDonough (1989); CT (continental tholeiite) after Holm (1985); IRT (initial rift tholeiite) and BABB (back-arc basin basalt) after Pouclet et al. (1995). Normalization values of Sun & McDonough (1989).
322
A. POUCLET ET AL.
of partial melting. No subduction zone enrichment is indicated. The relationships between the geochemical signatures and the geotectonic setting is illustrated by a comparison of the incompatible element patterns of the mafic lavas (excluding the cumulates) with typical profiles of lavas from different geodynamic sites (Fig. 12). Because of their erratic variations, Rb, Ba and Sr are not considered. We selected the normal mid-ocean ridge basalt (N-MORB) and ocean island basalt (OIB) recommended compositions of Sun & McDonough (1989), and
(a)
Tb x 3
N-
IAT
RB
MO
BABB
IRT LC CT
CA
B
OIB
UC
Ta x 2
Th (b)
Nb/3
OIB
the continental tholeiite (CT), initial rift tholeiite (IRT) and back-arc basin basalt (BABB) average compositions of Holm (1985) and Pouclet et al. (1995). No fit can be made with the active margin lavas. The best fit is obtained for the basalts of Ouzaga with the CT having weak negative Nb, Ta and Ti anomalies. A different CT, with no Ti anomaly, could provide a fit for the Ti-rich basalts of Ouzaga. The Ifri and Azegour lavas share the compositional features of CT and IRT, in terms of their Nb and Ta anomalies and their heavy rare earth element (HREE) depletion, respectively. A BABB-type pattern characterizes the most depleted lavas of Azegour in terms of their light rare earth elements (LREE) and large ion lithophile elements (LILE), but not their HREE. The Th–Tb –Ta diagram of Cabanis & Thie´blemont (1988) provides a synthesis of the prominent magmatic affinities corresponding to within-plate continental setting, rifting and marginal basin context (Fig. 13a). The significance of the Ti behaviour is tested in the TiO2 –Nb–Th diagram of Holm (1985) (Fig. 13b). This diagram separates within-plate from platemargin continental tholeiites. The plate-margin domain includes the IRT, which are characterized by a fractionated pattern close to that of the OIB, at lower trace element contents, by the lack of Nb and Ta anomalies, and by a weak positive anomaly of Ti (Fig. 12). These features are explained by a contribution of an asthenospheric source, which is added to the lithospheric source, the common source of the CT, during continental break-up, as a consequence of the upwelling of the asthenospheric mantle (Hawkesworth & Gallagher 1993). In the three areas, the two plate-margin tholeiite and within-plate tholeiite signatures are associated. This is symptomatic of an initial rift context characterized by a magmatic source contribution from both the continental lithosphere and the asthenosphere.
Plate-margin
Geodynamic evolution of the West African northern margin of the Palaeo-Gondwana
IRT CT Within-plate
TiO2
Th
Fig. 13. (a) Th– Tb 3– Ta 2 diagram of the Western High Atlas Early Cambrian volcanic rocks, after Cabanis & Thie´blemont (1988). Same symbols as for Figure 10. (b) TiO2 – Nb/3 –Th diagram of the Western High Atlas Early Cambrian volcanic rocks, after Holm (1985). Same symbols as for Figure 10.
During the Neoproterozoic, the West African northern margin of Palaeo-Gondwana was the site of a Pan-African orogenic belt. This belt resulted from the opening, around 750 Ma, and the closure, around 580 Ma, of an oceanic domain, at the northwestern border of the West African craton (Thomas et al. 2002). Northern and western Moroccan continental blocks belonging to the Carolina–Iberia terranes (Pique´ 2003), and Avalonian– Cadomian terranes were accreted to the West African craton along the Tizi n’Test and South Atlas Fault (Ennih & Lie´geois 2001, 2003) and along the Anti-Atlas Major Fault (Fig. 1; Bouougri 2003;
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS
Bouougri & Saquaque 2004; Soulaimani et al. 2006). Subsequently, an intracontinental calc-alkaline volcanic chain was formed, between 580 and 550 Ma, from the Atlantic coast to Algeria (Seddiki et al. 2004), in a transpressional tectonic context (Ennih & Lie´geois 2001). Then a post-orogenic extensional regime developed, related to the opening of a new ocean to the west, between the Avalonian terranes and the Mesetan– Atlas terranes (Keppie et al. 2003; Pique´ 2003; Soulaimani et al. 2003). This new regime is associated with a marine transgression and the formation of the western Morocco ‘Cambrian rift’, which is a deep graben trending SSW–NNE across the western Anti-Atlas, the westernmost part of High Atlas, and the western central Meseta (Fig. 14; Bernardin et al. 1988; Pique´ et al. 1990, 1995; Benssaou & Hamoumi 2003). The Ifri –Azegour and Ouzaga– Tizzirt formations belong to this tectonic and sedimentary event. According to their sedimentary patterns, the Ifri –Azegour formation was deposited at the edge of the basin, whereas the Ouzaga–Tizzirt formation was deposited in the middle part of the graben. Taking into account the Variscan right-lateral and NNW thrust motion along the MWHAF, the Ifri –Azegour deposits were located at the western border of the graben (Fig. 14). The geodynamic context of this basin formation could be (1) a back-arc basin of an active margin, (2) a passive margin ocean basin, or (3) an intra-continental rift. To choose between these models, we consider the nature of the volcanic rocks. The Ifri –Azegour and Ouzaga–Tizzirt lavas are classified as continental tholeiites. Their geochemical features show an association of within-plate and plate-margin magmatic signature. Similar volcanic rocks are known throughout the western High Atlas (Jouhari et al. 2001; El Archi et al. 2004). They display continental tholeiite and pseudo calc-alkaline features resulted from crustal contamination. Some other records of Cambrian volcanic activity occur throughout western Morocco (sites A –G in Fig. 14). Site A. In the Agoundis–Ounein region, east of the Ouzaga–Tizzirt area, the Early Cambrian sedimentary pile includes several volcanic rocks. The base of the pile consists of a 400 m thick stack of continental tholeiites, which unconformably overlie the rhyolitic complex of the Ediacaran AntiAtlas volcanic chain (Aarab et al. 2005; Pouclet et al. 2007). This volcanic activity can be dated to the latest Ediacaran Period. The lavas poured out from a N308-trending fissure system located at the edge of the rift, which indicates a WNW –ESE extensional regime. Later, the volcanic activity was almost continuous, resulting in flows and sills that are interbedded with the sedimentary
323
Fig. 14. Location of the Cambrian volcanic rocks in the Atlas and the western Meseta. Cambrian rift pattern adapted from Bernardin et al. (1988) and Benssaou & Hamoumi (2003). 1, Structural blocks: (I) northwestern region of the Ifri–Azegour formation; (II) southwestern region for the Ouzaga–Tizzirt formation; (III) southern region of the Agoundis–Ounein and Toubkal formation of Adoudounian type; (IV) Anti-Atlas region of the other Adoudounian formations. 2, outcrops of Late Ediacaran to Middle Cambrian volcanic rocks, in addition to the Ifri–Azegour and Ouzaga–Tizzirt areas: A, Agoundis-Ounein and Toubkal; B, Djbel Boho; C, Waoufengha– Igherm; D, Kerdous; E, Sidi-Saı¨d Maaˆchou; F, Bou Acila; G, oued Rhebar. 3, Domain of the Ediacaran Anti-Atlas volcanic chain (EAVC). 4, High Atlas and western Meseta outcrops of the Late Precambrian volcano-plutonic rocks after Morin (1962c), Corne´e et al. (1984), Corsini et al. (1988), Corne´e (1989), Eddif (2002) and Baudin et al. (2003); these rocks are the continental substratum of the Late Ediacaran and Early Cambrian formations. MWHAF, Middle Western High Atlas Fault; TNTF, Tizi-n’Test Fault; SAF, South Atlas Fault; AAMF, Anti-Atlas Major Fault.
sequences, from the Early Cambrian to the early Middle Cambrian. Site B. Farther to the east, the Djbel Boho volcanic activity located in the lower part of the Adoudounian sedimentary formations is dated at 531+5 Ma (Gasquet et al. 2005). The volcano provided alkaline lavas exhibiting an ocean island ´ lvaro et al. 2006; basalt chemical signature (A Pouclet et al. 2007). This intra-plate signature and the location of the volcano in the Adoudounian shelf of the Anti-Atlas region are consistent
324
A. POUCLET ET AL.
with a distant effect of the Cambrian rift volcano-tectonic activity. Sites C and D. To the west of the Anti-Atlas, basaltic flows are located at the base of the Adoudounian sequence and above the acid volcanic and detrital formations of the Ediacaran Anti-Atlas volcanic chain (Leblanc 1977; Demange 1980; Algouti et al. 2001; Soulaimani et al. 2004). They display the same continental tholeiite composition as the lavas of the Agoundis–Ounein region. Sites E–G. Cambrian terranes crop out in the central and northwestern Meseta. They overlie the late Precambrian volcano-plutonic rocks of the continental substratum. At the base of the Cambrian formations, basaltic lavas are known in the Sidi-Saı¨d Maaˆchou, Bou Acila and oued Rhebar areas (Gigout 1951; Corne´e et al. 1984; Morin 1962a–c). They have the composition of continental tholeiites and within-plate alkali basalts (a pseudo ‘orogenic’ signature is due to crustal contamination) and are dated from the Early to the Middle Cambrian (El Attari et al. 1997; Ouali et al. 2000, 2003; El Hadi et al. 2006). To sum up, all the Cambrian Atlas and Mesetan lavas have the composition of intra-continental tholeiitic to alkaline basalts. They are very different from the calc-alkaline andesitic to rhyolitic lavas of the Ediacaran Anti-Atlas volcanic chain. They cannot be related to any active margin volcanic arc. Consequently, the Cambrian rift is not a back-arc basin. The intra-plate and plate-margin signatures of the Cambrian lavas are characteristics of active rifting. Because of the marine sedimentation, a volcanic passive margin context may be suggested, instead of a true continental rift. However, in middle Cambrian time, the volcanic activity decreased (Pouclet et al. 2007), and there was no production of MORB-type lavas. It is concluded that the active rift aborted.
Conclusion In the western part of High Atlas, two distinct volcano-sedimentary formations are investigated: the Ifri –Azegour formation and the Ouzaga– Tizzirt formation. They belong to two structural blocks separated by the Middle Western High Atlas Fault (Figs 1 and 2), which was a right-lateral shear zone and a NNW– SSE thrust during the P2 and P3 tectonic phases of the Variscan orogeny (Figs 5 and 8). The Ifri –Azegour formation consists of siliciclastic and carbonate sediments of a shallow basin. It includes interbedded basaltic flows with the chemical composition of continental tholeiites. A dacite protusion of the upper part of the formation is dated at 533+4 Ma by the U –Pb ion microprobe method (Fig. 9). U –Pb dating of detrital zircons
(Fig. 9) indicates an origin from Palaeoproterozoic granitic rocks and from Neoproterozoic acid volcano –plutonic rocks. The Ouzaga –Tizzirt formation consists of a thick pile of siliciclastic and carbonate sediments with turbiditic sequences of a deep basin. It is attributed to Early to Middle Cambrian time, by correlations with palaeontologically dated similar formations located in the southern and western parts of the area. Interbedded flows at the base of the pile display a continental tholeiite composition. The lavas of both the Ifri–Azegour and Ouzaga –Tizzirt formations are characterized by an association of a plate-margin and a withinplate magmatic signature. This is indicative of an initial continental rift tectonic setting. Some other volcanic rocks are known from the Cambrian formations of the Atlas region and of the western Meseta (Fig. 14). They have compositions of continental tholeiites and intra-plate alkaline basalts. All these lavas are related to the volcano-tectonic activity of the western Moroccan Cambrian rift. An active margin context is discarded. The geodynamical context may be a passive margin related to a former active rift. The rifting aborted in the Middle Cambrian, with the cessation of the volcanic activity. This study was supported by the Institut des Sciences de la Terre d’Orle´ans (France) and by the French–Moroccan Action inte´gre´e 222/STU/00. We thank E. Deloule, M. Champenois and D. Mangin for the ion microprobe facilities at Nancy (France), and J.-P. Lie´geois, J.-Y. Cottin and J. M. Ugidos for critical reviews of the first draft. We also thank H. Lazrak and the SNAREMA Team for providing accommodation at the Ifri Mine, fieldwork facilities, and access to the sub-surface data.
References A ARAB , A., P OUCLET , A. & B OUABDELLI , M. 2005. Riftogene`se au Cambrien infe´rieur dans le nord-ouest du Pale´ogondwana. Exemple de la marge sud-est du Haut-Atlas occidental du Maroc. Annales de la Socie´te´ Ge´ologique du Nord, 12, 77– 85. A LGOUTI , A., A LGOUTI , A., C HBANI , B. & Z AIM , M. 2001. Se´dimentation et volcanisme synse´dimentaire de la se´rie de base de l’adoudounien infra-cambrien a` travers deux exemples de l’Anti-Atlas du Maroc. Journal of African Earth Sciences, 32, 541–556. ´ LVARO , J. J., E ZZOUHAIRI , H., V ENNIN , E. ET AL . A 2006. The Early-Cambrian Boho volcano of the El Graara massif, Morocco: Petrology, geodynamic setting and coeval sedimentation. Journal of African Earth Sciences, 44, 396– 410. B ARBANSON , L., C HAUVET , A., G AOUZI , A., B ADRA , L., M ECHICHE , M., T OURAY , J.-C. & O UKAROU , S. 2003. Les mine´ralisations Cu– (Ni–Bi– U– Au– Ag) d’Ifri (district du Haut Seksaoua, Maroc); apport de l’e´tude texturale au de´bat syngene`se versus e´pigene`se. Comptes Rendus Ge´oscience, 335, 1021– 1029.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS B AUDIN , T., C HE` VREMONT , P. & R AZIN , P. ET AL . 2003. Carte ge´ologique du Maroc au 1/50 000, feuille Skhour des Rehamna. Me´moire explicatif. Notes et Me´moires du Service Ge´ologique du Maroc, 435 bis. B ELFOUL , M. A., F AIK , F. & H ASSENFORDER , B. 2001. Mise en e´vidence d’une tectonique tangentielle ante´rieure au plissement majeur dans la chaıˆne hercynienne de l’Anti-Atlas occidental, Maroc. Journal of African Earth Sciences, 32, 723–739. B ENSSAOU , M. & H AMOUMI , N. 2003. Le graben de l’Anti-Atlas occidental (Maroc): controˆle tectonique de la pale´oge´ographie et des se´quences au Cambrien infe´rieur. Comptes Rendus Ge´oscience, 335, 297– 305. B ERNARDIN , C., C ORNE´ E , J.-J., C ORSINI , M., M AYOL , S., M ULLER , J. & T AYEBI , M. 1988. Variation d’e´paisseur du Cambrien moyen en Meseta marocaine occidentale: signification ge´odynamique des donne´es de surface et de subsurface. Canadian Journal of Earth Sciences, 25, 2104– 2117. B OUOUGRI , E. H. 2003. The Moroccan Anti-Atlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton. Precambrian Research, 120, 179–183. 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. C ABANIS , B. & T HIE´ BLEMONT , D. 1988. La discrimination des thole´iites continentales et des basaltes arrie`re-arc. Proposition d’un nouveau diagramme, le triangle Th–3 Tb–2 Ta. Bulletin de la Socie´te´ Ge´ologique de France, IV, 927– 935. C HAUVET , A., B ARBANSON , L., G AOUZI , A., B ADRA , L., T OURAY , J. C. & O UKAROU , S. 2002. Example of a structurally controlled copper deposit from the Hercynian western High Atlas (Morocco): the High Seksaoua mining district. In: B LUNDELL , D. J., N EUBAUER , F. & V ON Q UADT , A. (eds) The Timing and Location of Major Ore Deposits in an Evolving Orogen. Geological Society, London, Special Publications, 204, 247 –271. C ORNE´ E , J.-J. 1989. Le Haut Atlas occidental pale´ozoı¨que : un reflet de l’histoire hercynienne du Maroc occidental, stratigraphie, se´dimentologie et tectonique. PhD thesis, Universite´ de Marseille St. Jeroˆme. C ORNE´ E , J.-J., C OSTAGLIOLA , C. & L EGLISE , H. 1984. Lithostratigraphie et tectonique des terrains Ante´ce´nomaniens d’El Jadida, Me´seta Marocaine Hercynienne. Bulletin de la Faculte´ des Sciences de Marrakech, 2, 23–42. C ORNE´ E , J.-J., F ERRANDINI , J., M ULLER , J. & S IMON , B. 1987a. Le Haut-Atlas occidental pale´ozoı¨que: un graben cambrien moyen entre deux de´crochements dextres N 608E hercyniens (Maroc). Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 305, 499–503. C ORNE´ E , J.-J., D ESTOMBES , J. & W ILLEFERT , S. 1987b. Stratigraphie du Pale´ozoı¨que de l’extre´mite´ nord-ouest du Haut-Atlas occidental (Maroc hercynien); interpre´tation du cadre se´dimentaire du Maroc occidental.
325
Bulletin de la Socie´te´ Ge´ologique de France, III, 327– 335. C ORSINI , M., M ULLER , J., C ORNE´ E , J.-J. & D IOT , H. 1988. De´couverte de la se´rie basale du Cambrien et de son substratum dans les Rehamna Centraux, hautfond au Cambrien (Me´se´ta marocaine). Pre´mices de l’orogene`se hercynienne. Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 306, 63–68. D E K ONING , G. 1957. Ge´ologie des Ida ou Zal (Maroc). Stratigraphie, pe´trographie et tectonique de la partie SW du bloc occidental du massif ancien du Haut Atlas. Leidse Geologische Mededelingen, 23. D ELOULE , E., C HAUSSIDON , M., G LASS , B. P. & K OEBERL , C. 2001. U– Pb isotopic study of relict zircon inclusions recovered from Muong Nong-type tektites. Geochimica et Cosmochimica Acta, 65, 1833– 1838. D EMANGE , M. 1980. Stratigraphie, volcanisme et pale´oge´ographie du Pre´cambien III et de la se´rie de base dans la partie sud de la boutonnie`re d’Ouaoufengha (Anti-Atlas occidental) et mine´ralisation associe´e. Notes et Me´moires du Service Ge´ologique du Maroc, 41, 7 –23. E DDIF , A. 2002. Ge´ochronologie, pe´trologie, ge´ochimie et structure des intrusions tardi panafricaines de Wirgane et de leur couverture ne´oprote´rozoı¨que a` pale´ozoı¨que (Haut Atlas occidental, Maroc). PhD thesis, University of Rabat. E L A RCHI , A., E L H OUICHA , M., J OUHARI , A. & B OUABDELLI , M. 2004. Is the Cambrian basin of the Western High Atlas (Morocco) related either to a subduction zone or a major shear zone? Journal of African Earth Sciences, 39, 311– 318. E L A TTARI , A., H OEPPFNER , C. & J OUHARI , A. 1997. Nouvelles donne´es magmatiques et structurales en relation avec la cine´matique de l’ouverture du bassin cambrien de la Me´seta occidentale (Maroc). Gaia, 14, 11–21. E L H ADI , H., T AHIRI , A., S IMANCAS C ABRERA , F., G ONZALEZ L ODEIRO , F., A ZOR P E´ REZ , A. & M ARTINEZ P OYATOS , D. J. 2006. Un exemple de volcanisme calco-alcalin de type oroge´nique mis en place en contexte de rifting (Cambrien de l’oued Rhebar, Meseta occidentale, Maroc). Comptes Rendus Ge´oscience, 338, 229–236. 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 NNIH , N. & L IE´ GEOIS , J.-P. 2003. The Moroccan AntiAtlas: the West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton: reply to comments by E. H. Bouougri. Precambrian Research, 120, 185– 189. E SSAIFI , A., P OTREL , A., C APDEVILA , R. & L AGARDE , J.-L. 2003. Datation U –Pb: aˆge de mise en place du magmatisme bimodal des Jebilet centrales (chaıˆne Varisque, Maroc). Implications ge´odynamiques. Comptes Rendus Ge´oscience, 335, 193– 203. G AOUZI , A., C HAUVET , A., B ARBANSON , L., B ADRA , L., T OURAY , J.-C., O UKAROU , S. & E L W ARTITI , M. 2001. Mise en place syntectonique des
326
A. POUCLET ET AL.
mine´ralisations cuprife`res du gıˆte d’Ifri (district du Haut Seksaoua, Haut Atlas occidental, Maroc). Comptes Rendus de l’Acade´mie des Sciences, Sciences de la Terre et des Plane`tes, 333, 277–284. G ASQUET , D. 1991. Gene`se d’un pluton composite tardihercynien. Le massif de Tichka, Haut-Atlas occidental (Maroc). PhD thesis, Universite´ de Nancy 1. G ASQUET , D., L ETERRIER , J., M RINI , Z. & V IDAL , P. 1992. Petrogenesis of the Hercynian Tichka plutonic complex (Western High Atlas, Morocco): Trace element and Rb– Sr and Sm– Nd isotopic constraints. Earth and Planetary Science Letters, 108, 29–44. 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 IGOUT , M. 1951. Etudes ge´ologiques sur la Me´se´ta marocaine occidentale (arrie`re-pays de Casablanca, Mazagan et Safi). Notes et Me´moires du Service Ge´ologique du Maroc, 86. H AWKESWORTH , C. J. & G ALLAGHER , K. 1993. Mantle hotspots, plumes and regional tectonics as causes of intraplate magmatism. Terra Nova, 5– 6, 552– 559. H OEPFFNER , C., S OULAIMANI , A. & P IQUE´ , A. 2005. The Moroccan Hercynides. Journal of African Earth Sciences, 43, 144 –165. H OEPFFNER , C., H OUARI , M. R. & B OUABDELLI , M. 2006. Tectonics of the North African Variscides (Morocco, western Algeria): an outline. Comptes Rendus Ge´oscience, 338, 25–40. H OLM , P. E. 1985. The geochemical fingerprints of different tectonomagmatic environments using hygromagmatophile element abundances of tholeiitic basalts and basaltic andesites. Chemical Geology, 51, 303– 323. H OUARI , M.-R. & H OEPFFNER , C. 2003. Late Carboniferous dextral wrench-dominated transpression along the North African craton margin (Eastern High-Atlas, Morocco). Journal of African Earth Sciences, 37, 11–24. J OUHARI , A., E L -A RCHI , A., A ARAB , M., E L -A TTARI , A., E NNIH , N. & L ADURON , D. 2001. Ge´ochimie et cadre ge´odynamique du volcanisme ne´oprote´rozoı¨que terminal (vendien) du haut Atlas occidental, Maroc. Journal of African Earth Sciences, 32, 695– 705. K EPPIE , J. D., N ANCE , R. D., M URPHY , J. B. & D OSTAL , J. 2003. Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic –Paleozoic birth and development of peri-Gondwana terranes and their transfer to Laurentia and Laurussia. Tectonophysics, 365, 195–219. L AGARDE , J.-L. & R ODDAZ , B. 1983. Le massif plutonique du Tichka (Haut Atlas Occidental, Maroc): un diapir syntectonique. Bulletin de la Socie´te´ Ge´ologique de France, XXV, 389–395. L AVILLE , E. & P IQUE´ , A. 1991. La distension crustale atlantique et atlasique au Maroc au de´but du Me´sozoı¨que: le rejeu des structures hercyniennes.
Bulletin de la Socie´te´ Ge´ologique de France, 162, 1161– 1171. L EBLANC , M. 1977. Synchronisme des facie`s volcaniques (Pre´cambrien III) et se´dimentaires (Adoudounien) dans l’Infracambrien d’Alous (Anti-Atlas, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 284, 879–881. L UDWIG , K. R. 2003. Users’ manual for ISOPLOT/EX, version 3. A geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center, Special Publications, 4. M ORIN , P. 1962a. Premie`re preuve pale´ontologique de l’existence du Cambrien dans le Maroc central. Comptes Rendus de l’Acade´mie des Sciences, 254, 2198–2199. M ORIN , P. 1962b. Les se´ries volcano-se´dimentaires cambriennes du Maroc central. Comptes Rendus de l’Acade´mie des Sciences, 254, 2396– 2398. M ORIN , P. 1962c. Une vue d’ensemble nouvelle des formations ante´vise´ennes du pays des Zaı¨an (anticlinorium de Kasba–Tadla–Azrou, Maroc central). Comptes Rendus de l’Acade´mie des Sciences, 254, 3385–3387. M RINI , Z., R AFI , A., D UTHOU , J.-L. & V IDAL , P. 1992. Chronologie Rb–Sr des granitoı¨des hercyniens du Maroc: conse´quences. Bulletin de la Socie´te´ Ge´ologique de France, 163, 281–291. N ACHIT , H., R AZAFIMAHEFA , N., S TUSSI , J.-M. & C ARON , J.-P. 1985. Composition chimique des biotites et typologie magmatique des granitoı¨des. Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 301, 813–818. O UALI , H., B IAND , B., B OUCHARDON , J.-L. & M AAˆ TAOUI , M. 2000. Mise en e´vidence d’un volcanisme alcalin intraplaque d’aˆge Acadien dans la Meseta nord-occidentale (Maroc). Comptes Rendus de l’Acade´mie des Sciences, Sciences de la Terre et des Plane`tes, 330, 611– 616. O UALI , H., B IAND , B., B OUCHARDON , J.-L. & C APIEZ , P. 2003. Le volcanisme cambrien du Maroc central: implications ge´odynamiques. Comptes Rendus Ge´oscience, 335, 425– 433. O UAZZANI , H. 2000. Le pale´ovolcanisme des secteurs de Guedmioua et du Haut-Seksaoua (massif ancien du Haut-Atlas occidental, Maroc): te´moin d’un contexte convergent. PhD thesis, University of Mekne`s. O UAZZANI , H., B ADRA , L., P OUCLET , A. & P ROST , A. E. 1998. Mise en e´vidence d’un volcanisme d’arc ne´oprote´rozoı¨que dans le Haut-Atlas occidental (Maroc). Comptes Rendus de l’Acade´mie des Sciences, Sciences de la Terre et des Plane`tes, 327, 449– 456. O UAZZANI , H., P OUCLET , A., B ADRA , L. & P ROST , A. 2001. Le volcanisme d’arc du massif ancien de l’ouest du Haut-Atlas occidental (Maroc), un te´moin de la convergence de la branche occidentale de l’oce´an panafricain. Bulletin de la Socie´te´ Ge´ologique de France, 172, 587–602. P EARCE , J. A. 1983. Role of subcontinental lithosphere in magma genesis at active continental margin. In: H AWKESWORTH , C. J. & N ORRY , J. (eds) Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, 230–249. P EARCE , J. A., B ENDER , J. F., D E L ONG , S. E. ET AL . 1990. Genesis of collision volcanism in eastern Anatolia Turkey. Journal of Volcanology and Geothermal Research, 44, 189– 229.
EARLY CAMBRIAN VOLCANISM, HIGH ATLAS P IQUE´ , A. 2003. Evidence for an important extensional event during the Latest Proterozoic and Earliest Paleozoic in Morocco. Comptes Rendus Ge´oscience, 335, 865–868. P IQUE´ , A. & L AVILLE , E. 1995. L’ouverture initiale de l’Atlantique central. Bulletin de la Socie´te´ Ge´ologique de France, 166, 725–738. P IQUE´ , A., O’B RIEN , S., K ING , A. F., S CHENK , P. E., S KEHAN , J. W. & H ON , R. 1990. La marge nord-occidentale du Pale´o-Gondwana (Maroc occidental et zones orientales des Appalaches); Rifting au Pre´cambrien terminal et au Pale´ozoı¨que infe´rieur, et compression hercynienne et alle´ghanienne au Pale´ozoı¨que supe´rieur. Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 310, 411–416. P IQUE´ , A., B OUABDELLI , M. & D ARBOUX , J.-R. 1995. Le rift cambrien du Maroc occidental. Comptes Rendus de l’Acade´mie des Sciences, Se´rie IIa, 320, 1017–1024. P IQUE´ , A., A¨I T B RAHIM , L., A¨I T O UALI , R. ET AL . 1998. Evolution structurale des domaines atlasiques du Maghreb au Me´so-Ce´nozoı¨que; le roˆle des structures he´rite´es dans la de´formation du domaine atlasique de l’Afrique du Nord. Bulletin de la Socie´te´ Ge´ologique de France, 169, 797–810. P OUCLET , A., L EE , J.-S., V IDAL , Ph., C OUSENS , B. & B ELLON , H. 1995. Cretaceous to Cenozoic volcanism in South Korea and in the Sea of Japan: magmatic constraints on the opening of the back-arc basin. In: S MELLIE , J. L. (ed.) Volcanism Associated with Extension at Consuming Plate Margins. Geological Society, London, Special Publications, 81, 169 –191. P OUCLET , A., A ARAB , A., F EKKAK , A. & B ENHARREF , M. 2007. Geodynamic evolution of the north-western Paleo-Gondwanan margin in the Moroccan Atlas at the Precambrian– Cambrian boundary. In: L INNEMANN , U., N ANCE , R. D., K RAFT , P. & Z ULAUF , G. (eds) The evolution of the Rheic Ocean: From Avalonian– Cadomian Active Margin to Alleghanian –Variscan Collision. Geological Society of America, Special Papers, 423, 27– 60. P ROST , A. E., B ADRA , L. & E L H ASNAOUI , H. 1989. Superposition de trois de´formations ductiles hercyniennes dans le Haut Atlas (re´gion d’Azegour– Erdouz, Maroc). Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, 309, 627–632. P ROUST , F., P ETIT , J.-P. & T APPONNIER , P. 1977. L’accident du Tizi n’Test et le roˆle des de´crochements dans la tectonique du Haut Atlas occidental (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, XIX, 541–551. P UPIN , J.-P. 1980. Zircon and granite petrology. Contributions to Mineralogy and Petrology, 73, 207–220.
327
R OSSI , P. & C HE` VREMONT , P. 1987. Classification des associations magmatiques des granitoı¨des dans le cadre de la carte ge´ologique a` 1/50 000 de la France. Ge´ochronique, 21, 14–18. S CHAER , J. P. 1964. Volcanisme cambrien dans le massif ancien du Haut Atlas occidental. Comptes Rendus de l’Acade´mie des Sciences, 258, 2114– 2117. S CHAER , J. P., D UFFAUD , F., T IXERONT , M. & H OLLARD , H. 1981. Carte ge´ologique du Maroc, feuille d’Imi-n’Tanout au 1/100 000. Notes et Me´moires du Service Ge´ologique du Maroc, 203. S EDDIKI , A., R EMACI -B ENAOUDA , N., C OTTIN , J.-Y., M OINE , B. N., M E´ NOT , R.-P. & P ERRACHE , C. 2004. The volcano-sedimentary Boukaı¨s inlier (southwestern Algeria): evidence for lithospheric thinning during the Late Neoproterozoic. Journal of African Earth Sciences, 39, 257 –266. S OULAIMANI , A., B OUABDELLI , M. & P IQUE´ , A. 2003. L’extension continentale au Ne´o-Prote´rozoı¨que supe´rieur–Cambrien infe´rieur dans l’Anti-Atlas (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, 174, 83– 92. 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 (Anti-Atlas occidental, Maroc). Comptes Rendus Ge´osciences, 336, 1433– 1441. S OULAIMANI , A., J AFFAL , M., M AACHA , L., K CHIKAH , A., N AJINE , A. & S AIDI , A. 2006. Mode´lisation magne´tique de la suture ophiolitique de Bou Azzer– El Graara (Anti-Atlas central, Maroc). Implications sur la reconstitution ge´odynamique panafricaine. Comptes Rendus Ge´oscience, 338, 153– 160. 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 Scientific, Oxford. 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 UCKER , R. D. & M C K ERROW , W. S. 1995. Early Paleozoic chronology: a review in light of new U– Pb zircon ages from Newfoundland and Britain. Canadian Journal of Earth Sciences, 32, 368– 679.
The magmatic evolution at the Moroccan outboard of the West African craton between the Late Neoproterozoic and the Early Palaeozoic H. EZZOUHAIRI1, M. L. RIBEIRO2, N. AIT AYAD1, M. E. MOREIRA3, A. CHARIF1, J. M. F. RAMOS3, D. P. S. DE OLIVEIRA2,4 & C. COKE5 1
Universite´ Chouaı¨b Doukkali, Faculte´ des Sciences, Department de Ge´ologie, BP. 20, 24 000, El Jadida, Morocco (e-mail:
[email protected]) 2 ´ INETI—Area de Geocieˆncias, Estrada da Portela, Zambujal, 2721-866 Alfragide, Portugal 3
Laboratoire de INETI, R. da Amieira, Ap. 89, 4465, S, Mamede de Infesta Codex, Portugal 4
CREMINER, Centro de Recursos Minerais, Mineral. e Cristalografia, Ed. C6, piso 3, 1749-016, Lisbon, Portugal 5
Universidade de Tras-Os-Montes e Alto Deuro, Apartada 1013, 5001-911 Vila Real codex, Portugal
Abstract: The Late Neoproterozoic Ouarzazate Group crops out on the north margin of the West African craton (WAC). In this group an important post-collisional magmatism is characterized by a great diversity in plutonic and volcanic rock types of the high-K calc-alkaline series. This series evolved mainly by crystal fractionation and by an important crustal contamination from an anomalous mantle source. The Early Cambrian magmatism began at the same time on both sides of the Anti-Atlas Major Fault, the southwestern side (Kerdous region) and northeastern side (Ouarzazate-Agdz region), interbedded in the Early Cambrian Basal Series and spread later to the Western High Atlas of the Morocco northern WAC outboard areas. This magmatism changes from a continental tholeiitic series (HPT and LPT) at the beginning to an alkaline series at the top (Adoudou and ‘Lie de vin’ formations). Fractional crystallization and pelagic or crustal contamination were the most important processes in the magma differentiation. The geochemical inversion from calc-alkaline to tholeiitic magmatism between the Late Neoproterozoic and the Early Cambrian is documented, as is the major extension of the tholeiitic activity on both sides of the South Atlas Fault. This geochemical variation indicates a transition of the tectonic regime from compressive to extensional. The late local Jbel Boho alkaline magmatism indicates the sink of the source and the mitigation or closure of the extensional cycle at this time.
The precise boundary of the Moroccan outboard of the West African craton (WAC) is controversial. For some workers, the Anti-Atlas Major Fault (AAMF) marks this boundary (Fig. 1): on the southwestern side of the AAMF the Eburnean deformed and metamorphosed Precambrian basement (Choubert 1963; Charlot 1982; Ait Malek et al. 1998; Walsh et al. 2002) crops out, with a Neoproterozoic cover deformed by the Pan-African orogeny (Clauer 1976; Clauer et al. 1982); on the northeastern side of the AAMF the ‘Pan-African Anti-Atlas mobile zone’ crops out; this corresponds to the Pan-African deformed Neoproterozoic formations (Leblanc & Lancelot 1980; Saquaque et al. 1989; Hefferan et al. 2000). However, the findings of Eburnean material in Saghro sediments, as well as the old crustal isotopic signatures in some Neoproterozoic granitoids on both sides of the AAMF have led other workers to place
the WAC northwestern margin boundary on the South Atlas Fault (SAF) (Fekkak et al. 2000; Ennih & Lie´geois 2001; Errami et al. 2002). Furthermore, the recent aeromagnetic data of Soulaimani et al. (2006) suggest that the AAMF is an ophiolitic suture zone whose subduction plane dips to the south (Levresse 2001). The Bou-Azzer El Graara blueschist metamorphism seems to reinforce this hypothesis (Levresse 2001; Hefferan et al. 2002). The oblique collision of the WAC north margin and the Saghro volcanic arc (Saquaque et al. 1989; Hefferan et al. 1992; Thomas et al. 2002) produced the Pan-African heterogeneous deformation and metamorphism. These events have been dated respectively at 750–650 Ma (Inglis et al. 2005; D’Lemos et al. 2006) and 685 Ma (Clauer 1974). Late Neoproterozoic formations overlapping the Pan-African deformed Early Neoproterozoic are widespread in geological windows of the Anti-Atlas
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 329–343. DOI: 10.1144/SP297.16 0305-8719/08/$15.00 # The Geological Society of London 2008.
330
H. EZZOUHAIRI ET AL.
Fig. 1. Simplified geological map showing the Early Cambrian formations in the High Atlas and Anti-Atlas Domains. 1, Fault; 2, Lower Cambrian; 3, Late Neoproterozoic; 4, Palaeoproterozoic and Early Neoproterozoic.
domain and outcrop, and also in some windows of the High Atlas domain (Fig. 1). The Anti-Atlas and High Atlas Late Neoproterozoic formations, locally known as the Ouarzazate Group (or ‘Precambrian III’), correspond to fluvial and lacustrine sequences with interbedded magmatic rocks unaffected by the Pan-African orogeny. The magmatic rocks of the Ouarzazate Group formations display geochronological ages of between 574 + 5 Ma (U/Pb) for the upper ignimbrite (Mifdal & Peucat 1985) and 543 + 9 Ma for the rhyolitic dyke in the eastern Anti-Atlas (Gasquet et al. 2005), and correspond to the late and post-orogenic magmatism (Youbi 1998; Ezzouhairi 2001; Thomas et al. 2002; Inglis et al. 2004). The Anti-Atlas and the High Atlas Late Neoproterozoic magmatic rocks display large petrographical varieties ranging from basalts, andesites, dacites, rhyolites–ignimbrites and pyroclastic tuffs to intrusive rocks such as gabbro–diorites and granites (Ezzouhairi et al. 1998; Youbi 1998; Ezzouhairi 2001; Zahour 2001; Thomas et al. 2002). The Adoudou Formation overlaps the Ouarzazate Group of those regions. The Adoudou Formation consists of a lower conglomeratic–siliciclastic Basal Member and an upper Limestone Member or (‘Calcaires infe´rieurs’), essentially composed of limestone and dolostones (Fig. 2a). The mafic volcanic rocks
interbedded in this formation have geochronological ages on the Precambrian–Cambrian boundary (e.g. Jbel Boho U–Pb data of Ducrot & Lancelot (1977) and Gasquet et al. (2005) respectively 534 + 10 Ma and 531+5 Ma) (Fig. 2). The Adoudou Formation is overlain by the ‘Lie-de-vin’ Formation (or Taliwine Fm, dated at 522+2 Ma and 521+7 Ma (Compston et al. 1992; Landing et al. 1998; Maalouf et al. 2005), and the Upper Limestone Formation (‘Calcaires supe´rieurs’ Fm) (or Igoudine Fm) (Choubert 1963; Destombes et al. 1985; Buggisch & Flu¨gel 1988; Geyer 1990; Benssaou & Hamoumi 2001; Thomas et al. 2002; Alvaro et al. 2006). These formations, the Adoudou Fm and ‘Lie-de-vin’ formations, which are about 1000 m thick in the Western Anti-Atlas but only a few hundred metres thick in the Eastern Anti-Atlas and High Atlas, are known to mark the Neoprotero´ lvaro zoic–Cambrian transition (Benssaou 2005; A et al. 2006). The palaeontological discoveries in the lower limestone (Buggisch & Flu¨gel 1988), and the geochronological dating (U–Pb) of the Aghbar trachyte (531 + 5 Ma) (in the Bou-Azzer region) as well as the intercalated Jbel Boho syenite at the boundary of the Adoudou–Lie-de-Vin formations (Leblanc & Lancelot 1980; Gasquet et al. 2005), confirm that the entire Adoudou Formation could
MAGMATIC EVOLUTION OF NORTHERN WAC
331
Fig. 2. Location of the representative stratigraphic columns and geological profiles (a– f) of the study areas. Data sources: b, Ezzouhairi (2001); e, Ait-Ayad et al. (1998); f, Soulaimani et al. (2004); g, Choubert (1952) and Choubert & Faure-Muret (1970). A.F., Adoudou Formation; L.V.F., ‘Lie-de-vin’ Formation; U.L.F, Upper Limestone Formation; sd, sandstone; s, syenite; b, basalt; t.t., trachyandesite and trachyte; LN, Late Neoproterozoic; PP, Palaeoproterozoic.
extend to the Early Cambrian. Benssaou & Hamoumi (2001), Soulaimani et al. (2003) and Benssaou (2005) share this opinion. This paper presents a summary of the research carried out at several of the main outcrops of magmatic rocks interbedded in the Late Neoproterozoic and Early Cambrian formations of the Anti Atlas and Western High Atlas of the northern WAC margin. The most significant geological and geochemical signatures found are presented and a palaeogeological environment is suggested.
Stratigraphical framework and general petrographic features The simplified regional geological map provides a general view of the areas of the Late Neoproterozoic
and Early Cambrian formations that crop out in the northern WAC (Fig. 1). The stratigraphical columns presented in Figure 2 present information about the extent of the magmatic activity, its abundance, variety and relative stratigraphical positions, during the Neoproterozoic–Cambrian transition in the most representative regions of the Anti-Atlas and High Atlas domains. A general succession across the Late Neoproterozoic–Cambrian terranes is also presented in Figure 2a. The studied magmatic rocks (extrusive and intrusive) of the Late Neoproterozoic Ouarzazate Group and the overlying Adoudou Formation were collected in all the indicated regions, except for the data for the Jbel Kerkar – Ida Ougnidif (Figs 1 and 2), which were taken from geological literature (Soulaimani et al. 2004).
332
H. EZZOUHAIRI ET AL.
The Late Neoproterozoic The Late Neoproterozoic magmatism of the central–eastern Anti-Atlas displays a great diversity of rock types, both extrusive and intrusive. The extrusive rock types include andesitic basalts and andesites with microlithic or porphyritic textures (Fig. 3b), porphyritic dacites with immiscible amygdaloidal bodies of quartz and potassium feldspars, and rhyolites with ignimbritic textures. The basaltic rocks contain plagioclase, clinopyroxene and rare altered olivine phenocrysts in a groundmass of plagioclase, clinopyroxene and iron oxide microliths. The clinopyroxene of the basalts has complex oscillatory zoning as a result of the oxygen fugacity variation (Ezzouhairi 2001). The andesitic rocks usually contain plagioclase and hornblende phenocrysts. Microliths of plagioclase and hornblende with iron oxides are abundant in their groundmass. The dacites have the same mineralogy as the andesites but with some interstitial quartz in the groundmass. The plagioclase phenocrysts, as well as the alkali feldspars and quartz of the acidic volcanic rocks (rhyolite and ignimbrite), are occasionally accompanied by biotite. The quartz and alkali feldspar grains of these rocks are most common in the groundmass. This magmatic mineralogy was partially transformed by low-grade metamorphism into secondary mineral associations of chlorite, epidote, calcite, actinolite and iron oxides. These rocks and the associated pyroclastic tuffs form a sequence that was intruded by small bodies and dykes of gabbro–dioritic rocks (Fig. 2a) and by some intrusive granites and microgranites. The local gabbro–diorite (dated by U –Pb at 561 + 2 Ma, Chebbaa 1996) is composed by plagioclase, clinopyroxene, amphibole, biotite and rare xenocrysts of interstitial quartz and alkali feldspar. The plagioclase, usually the most abundant mineral phase, generally displays normal zoning. The locally intrusive granites (dated by U –Pb at 580 + 5 Ma (Bouskour granite), 561 + 3 Ma (Ouarzazate microgranite) and 562 + 5 Ma (Siroua granite) by Mrini (1993), Chebbaa (1996) and Thomas et al. (2002), respectively) have the usual mineralogy of this rock type: plagioclase, alkali feldspar, quartz, biotite (locally transformed into chlorite) and zircon (the most common accessory mineral). Early Cambrian conglomerates, sandstones, siltstones and limestones overlie all these lithotypes (Ezzouhairi 2001).
The Early Cambrian The most common igneous Early Cambrian lithotypes cropping out in the Anti-Atlas and High Atlas domains of the northern WAC, except the
Alougoum region, are the aphyric to sub-aphyric basalts. Microlithic or porphyritic textures can also be found. The original igneous associations are usually entirely replaced by a low-grade metamorphic paragenesis. Albite, chlorite, actinolite, epidote, calcite, quartz and iron oxides are the main mineral components of these rocks. In the Anti-Atlas domain, we have considered two sub-domains, the central –eastern Anti-Atlas and the western Anti-Atlas. In the central –eastern Anti-Atlas, three regions were studied: Ouarzazate –Agdz, NE Bou Azzer and Alougoum (Jbel Boho). The lithostratigraphical column of the Ouarzazate –Agdz region begins with several magmatic events (volcanic and plutonic) emplaced during the Late Neoproterozoic and Early Cambrian (Figs 2b,c and 3a,b). The Basal Series that unconformably overlaps the Late Neoproterozoic mostly corresponds to a conglomerate of about 30 m thickness, although it is usually thicker near the fault zones. This conglomerate is polymictic and poorly sorted (clasts size varies from a few centimetres to over 3.5 dm; Fig. 3e). One or two levels of basaltic flows, cropping out over several square kilometres, lie on top of this conglomerate and under the Lower Limestone Unit (Figs 2c and 3c–f ). In contrast, in the NE of the Bou Azzer region surrounding the AAMF, where the Basal Series conglomerate is thinner, the Cambrian basalts are interbedded with the Lower Limestone Unit (Fig. 2d). In the Alougoum (Jbel Boho) region, Late Neoproterozoic sedimentary and magmatic formations are overlain by the Basal Series conglomerate, sometimes with local angular unconformities (Fig. 2g). These volcanic flows are intercalated at the boundary of the Lower Limestone and the ‘Lie-de-vin’ Formations (Fig. 2a). The Alougoum volcano consists of a syenitic–trachytic core surrounded by a great variety of volcanic rocks, which are essentially of basaltic composition on the northern flank and trachytic–rhyolitic with associated pyroclastic tuffs and breccias on the southern flank. The original mineralogy of the basalts (plagioclase, clinopyroxene and olivine) was almost completely replaced as a result of low greenschist-facies metamorphism. However, in the most differentiated rocks, igneous feldspars and small relics of pyroxenes, amphiboles and biotites are frequently preserved. For the western Anti-Atlas the data correspond to the Jbel Kerkar region (JK), for outcrops in the geological window of Kerdous (Fig. 1). This region has a general stratigraphical column equivalent to that presented in Figure 2f. From the base to the top, the Tazeroualt Palaeoproterozoic granite is overlapped by the Basal Series conglomerate
MAGMATIC EVOLUTION OF NORTHERN WAC
333
Fig. 3. (a) Panorama of the western Ait Saoun region (central Anti-Atlas), showing the thickening of the basal conglomerate level near the fault zone. (b) Late Neoproterozoic porphyritic andesite (western Ait Saoun region, Agdz). (c) Massive basaltic flow overlapping basal conglomerate in the central Anti-Atlas (Ouarzazate –Agdz region), scale bar represents 15 m. (d) Basaltic pillow lavas in the Azegour– Wirgane region (High Atlas). (e) Heterogeneous conglomerate of Adoudou basal series near Ait Saoun village (central Anti-Atlas). (f) Red Lie-de-vin Formation overlapping the lower limestone (50 m thick) of the Adoudou Formation in the central Anti-Atlas. C, conglomerates; B, basalt; L, limestone; L-V-F, Lie-de-vin Formation.
334
H. EZZOUHAIRI ET AL.
Table 1. Selected chemical analyses of the Late Neoproterozoic magmatic rocks Sample:
HA13
HA13a
HA4
153b
137b
177a
167e
HA2
131c
162a
SiO2 Al2O3 Fe2OT3 MnO CaO MgO Na2O K2O TiO2 P2O5 LOI Total Rb Sr Y Zr Nb Ba Ta Th Hf V La Ce Pr Nd Sm Eu Yb
49.27 18.18 9.03 0.3 6.8 3.71 2.93 4.01 1.51 0.32 2.53 98.59 130 478 25.2 183 19 1266 0.8 3 4.2 248 23.1 51.5 5.9 27.2 6 1.79 2.51
49.13 18.39 9.1 0.29 6.85 3.8 3.2 4.27 1.54 0.34 2.79 99.7 127 517 34 193 18 1364 0.8 4 4 236
51.12 16.25 8.78 0.19 5.30 5.69 4.58 2.58 1.35 0.33 2.78 98.95 75 563 28.0 182 13 722 1.0 3.7 4.0 171 20.0 39.0 0.0 20.0 4.48 1.40 2.01
51.10 15.71 8.94 0.20 5.68 5.86 3.99 2.56 1.35 0.33 3.38 99.10 79 553 26.6 204 14 1632 1.6 3.0 4.8 301 25.0 55.9 6.2 29.0 6.30 1.71 2.68
51.81 15.87 9.11 0.18 5.12 5.88 3.83 2.20 1.41 0.34 3.08 98.83 54 371 28.2 203 15 667 0.0 3.0 4.6 153 27.4 59.9 7.2 30.9 6.70 1.75 2.76
55.36 17.29 7.71 0.20 4.59 3.47 3.82 3.11 0.77 0.17 2.17 98.66 87 431 17.8 91 6 861 0.0 0.0 2.1 187 14.1 31.1 3.6 15.7 3.60 1.07 1.92
55.14 15.10 7.47 0.15 5.27 5.73 3.81 3.23 0.73 0.17 2.13 98.93 80 597 16.1 92 5 1214 0.0 0.0 2.2 159 16.0 36.5 4.1 19.1 4.00 1.09 1.67
49.37 17.00 10.25 0.25 6.28 3.71 4.12 3.01 1.77 0.39 3.08 99.23 91 472 30.0 198 16 1386 0.8 3.0 4.0 228 23.0 50.0 0.0 24.0 5.03 1.60 2.28
51.01 17.20 9.29 0.19 6.90 3.50 3.20 3.75 1.70 0.42 1.69 98.85 128 532 22.0 152 17 1253 0.0 3.0 3.1 182 23.3 50.2 5.6 24.5 5.20 1.43 2.12
51.75 17.60 8.35 0.15 5.65 3.17 4.09 4.00 1.53 0.44 1.97 98.70 131 608 29.3 243 17 1197 1.5 4.0 5.5 140 35.6 77.1 8.7 38.6 8.00 2.12 2.86
(20 m thick), then basaltic flows of about 30 m thick, which are capped by pyroclastic rocks. Finally, overlapping this conglomerate there are some levels of sandstones, superimposed by the Lower Limestone Unit, which completes the stratigraphical column. The JK lava flows, which crop out over a distance of 9 km, have fine-grained, microlithic and vacuolar textures, from the base to the top. They become pyroclastic in the highest few centimetres. These pyroclastic levels include old fossilized mud cracks, suggesting contemporaneous exposure to erosion. The JK lava flows are metabasaltic rocks that have undergone lowgrade greenschist facies metamorphism, and the original doleritic, sub-ophitic and microlithic textures are often preserved. The same lithostratigraphy as at JK can be observed in the Ida Ougnidif region, 60 km to the NE (Fig. 1). In this region the basaltic flows are thicker and have pillow lava structures. In the High Atlas domain, the northern WAC includes the western part of the domain. The two studied regions of this domain, Azegour and Wirgane, have similar lithostratigraphical columns
with slight variations in the thickness and relative position of the geological units (Fig. 2e). Over the acidic volcano-sedimentary beds attributed to the Late Neoproterozoic a massive level of basalt and pillow lavas crops out (Fig. 3d). These rocks are overlain by the sandstones of the top of the Basal Series, which, in turn, are overlain by the Lower Limestone Unit (Fig. 2e) (Ait Ayad et al. 1998). Microscopically, these are low-grade greenschistfacies metabasalts. Microlithic porphyritic textures, locally doleritic, are often preserved. Some relics of plagioclase phenocrysts (1–3 mm) have anorthite contents ranging from 25 to 42%. The matrix, which corresponds to about 60% of the total volume, display microlites of sericitized plagioclase and small grains of retrograded iron–magnesium minerals (chlorite, clinoamphiboles, epidote and iron oxides).
Geochemistry and petrological data Geochemical data are presented in chronostratigraphical order from Neoproterozoic to Cambrian. Analyses were carried out using X-ray fluorescence
MAGMATIC EVOLUTION OF NORTHERN WAC
335
HA9
HA1
137a
152a
HA6
169e
132a
180d
RBS2
RBS1
54.72 17.60 8.47 0.12 0.50 2.87 2.73 8.60 0.85 0.21 2.51 99.18 173 64 16.0 122 6 2429 0.3 3.4 3.0 150 15.0 27.0 0.0 10.0 1.85 0.50 1.75
61.89 14.89 6.93 0.09 1.72 1.76 6.42 1.90 0.65 0.17 2.44 98.86 37 68 18.0 98 6 366 0.7 2.5 2.0 155 5.0 15.0 0.0 10.0 2.46 0.60 1.48
57.54 15.18 8.36 0.1 1.04 2.64 3.88 5.22 1.63 0.44 2.29 98.32 99 143 37.2 302 23 1016 1.4 11 6.2 186 37.1 87.5 10.4 43.6 8.9 1.98 3.84
58.98 15.11 9.30 0.07 0.91 0.28 2.29 8.50 1.76 0.52 0.96 98.68 161 51 42.3 353 26 1805 1.7 9.0 7.2 322 22.5 56.6 9.3 45.7 10.40 2.48 4.49
63.59 13.30 7.29 0.05 0.70 0.80 1.01 8.76 1.33 0.38 1.52 98.73 149 35 44.0 345 15 1504 1.3 11.2 7.0 102 42.0 96.0 0.0 46.0 8.12 2.20 3.47
70.09 14.20 3.36 0.36 0.24 3.95 6.01 0.09 0.41 0.02 0.70 99.44 126 69 23.7 235 11 1283 0.00 10.0 4.7 31 17.5 35.5 3.9 18.1 4.40 0.80 2.74
70.12 12.87 2.45 0.06 0.45 1.33 0.46 9.20 0.37 0.08 1.60 98.99 165 40 21.6 216 11 1252 0.0 19.0 5.8 191 25.6 58.5 5.2 19.7 3.30 0.63 2.58
65.56 15.93 3.54 0.04 0.46 1.18 4.72 5.67 0.59 0.21 1.30 99.20 89 50 23.1 221 9 1130 0.0 5.0 5.5 97 21.4 50.8 5.5 23.9 5.00 0.94 2.74
75.13 11.40 1.67 0.03 0.13 0.59 0.07 8.92 0.10 0.01 0.91 98.96 207 27 47.7 287 16 1308 0.0 14.0 8.9 0 6.8 15.7 1.7 9.4 3.40 0.29 6.28
75.76 12.08 2.13 0.04 2.58 0.93 4.47 0.41 0.15 0.03 0.89 99.47 5 180 49.6 182 10 144 0.0 9.0 4.6 27 41.6 95.3 10.3 42.9 9.10 1.45 6.00
Samples HA13, HA13A, 177A, 167e, HA2, 131c, 162a and HA9 are basalt; samples HA4, 153b and 137b are gabbro–diorite; samples HA1, 137a, 152a and HA6 are andesite –dacite; samples 169e, 132a, 180d, RBS2 and RBS1 are rhyolite– ignimbrite. The Lambert coordinates of these rocks are respectively: HA13 (368.2; 427.4); HA13a (368; 428); 177a (362.1; 426.8); 167e (361.4; 426.4); HA2 (363.9; 427.6); 131c (366.5; 428.8); 162a (362.3; 426.3); HA9 (367.7; 427); HA4 (364.1; 425.9); 153b (364.2; 426); 137b (364.4; 426.2); HA1 (364; 427.2); 137a (364.5; 425.6); 152a (363.3; 426.1); HA6 (363; 426.3); 169e (361.5; 425.4); 132a (366.6; 428.4); 180d (369.4; 427.5); RBS2 (412.6; 440.6); RBS1 (413.8; 439.9).
(XRF) at the certified laboratories of INETI, Porto, Portugal and Ontario, Canada. Precision for major and trace elements is usually better than 2% and 5–10%, respectively. Twenty representative geochemical analyses carried out on samples from the Ouarzazate –Agdz Late Neoproterozoic formations are presented in Table 1. The relatively high percentages of loss on ignition (LOI) in this table, between 3.38 and 0.58, indicate the presence of some hydrous mineral associations (with chlorite and actinolite) produced by the regional low-grade metamorphism. In ancient volcanic rocks such as these, which have undergone post-magmatic processes such as low-grade metamorphism, weathering and, possibly some spilitization, high field strength elements have to be used (Winchester & Floyd 1976). However, the large ion lithophile elements (LILE; K, Rb and Ba) which are useful to constrain the geological environment, were also used with caution. In fact,
the multi-element patterns obtained display enrichment of these elements, which suggests that their mobility was not so extensive that they could not be used. The analytical results (Table 1) show compositions ranging from basic (49% , SiO2 , 55%) to acidic (SiO2 . 75%). Their projection on the SiO2 v. Zr/TiO2 0.0001 diagram (Winchester & Floyd 1976) shows a great variety of rock types (basalts, andesites, dacites and rhyolites) according to the petrographical determination (Fig. 4a). The Nb/Y values ,1 indicate a sub-alkaline character (Fig. 4a). The K2O/Na2O ratios of basalt are variable from 0.7 to 1.4, suggesting their potassic affinity (Ezzouhairi 2001). The very conspicuous Ba concentration observed in the analyses in Table 1 is usually associated with these rock types. Such K and Ba enrichments are usually interpreted as the result of metasomatic processes in the source region (Wilson 1994).
336
H. EZZOUHAIRI ET AL.
(a)
Zr/TiO2*0.0001
1
(b)
Ouarzazate-Agdz Rhyolite/ignimbrite Andesite/dacite basalt gabbro-diorite
Com/Pant
Phonolite
Rhyolite 0.1
Basalt, Ait Saoun Basalt, Kerkar-Ida Ougnidif Basalt, Azegour Basalt, Wirgane-Talat N’Yacoub Basalt, Boho/Bou Azzer Trachyandesite-Trachyte, Boho
Trachyte
Rhyolite Rhyodacite/Dacite
Rhyodacite/Dacite TrachyAnd
Trachyte
Andesite Bsn/Nph
Andesite/Basalt Alk-Bas 0.1
Bsn/Nph
Andesite/Basalt Alk-Bas
SubAlkaline Basalt 0.001 0.01
Phonolite
TrachyAnd
Andesite 0.01
Com/Pant
SubAlkaline Basalt 1
Nb/Y
10 0.01
0.1
1
Nb/Y
10
Fig. 4. Projection of Zr/TiO2 v. Nb/Y for the Late Neoproterozoic (a) and the Early Cambrian igneous rocks (b).
The spider diagrams of the Late Neoproterozoic rock analyses show subparallel patterns and a general enrichment of incompatible elements, and negative anomalies of Nb–Ta, Ti and P for the most fractionated rocks (Fig. 5a). These signatures are usually shown by the calc-alkaline series of the continental subduction zones: the low ionic potential elements are attributed to metamorphism in the mantle source by fluids rising from a subducted slab; the Ti and P negative anomalies suggest continental contamination. The Nb–Ta negative anomalies could suggest retention by accessory mineral phases at the source or from continental contamination. The most differentiated rocks show negative anomalies of Sr, suggesting plagioclase fractionation. For comparative purposes we have used a representative pattern of the most typical continental subduction zone, a calc-alkaline basalt from Central Chile (Fig. 5a). The good correlation of the profiles suggests that their signatures were produced in similar environments. The Ti/Alt ratios of the studied clinopyroxenes of the basalts also confirm the calc-alkaline character of these rocks (Ezzouhairi 2001). To determine the processes involved in the genesis and differentiation of these rocks the Zr/Y v. Ti/Y and Ce/Zr v. Zr diagrams are used. When these ratios are plotted on the Crowley et al. (2000) diagram they indicate that the series has evolved by fractional crystallization and crustal contamination processes (pelagic sediments and continental crust, PASC: Fig. 6a). As previously stated, the Early Cambrian volcanic rocks crop out either directly over the Basal Member conglomerates (e.g. Ouarzazate –Agdz, Kerkar –Ida Ougnidif and Azegour –Wirgane regions) or intercalated at the top of the Lower Limestone Unit (e.g. Jbel Boho–Bou Azzer
region), suggesting deposition over a long period of time. Massive basalts are the most frequent lithotype that crops out. Usually they occur as lava flows, and locally pillow-lava structures can appear. However, in the Jbel Boho region, where the basaltic type is also predominant, trachyandesites, trachytes and rhyolitic rocks also occur. All these rocks are usually metamorphosed at the low greenschist-facies conditions. Plagioclase, chlorite, actinolite, epidote, calcite and iron oxides are the most common phases in the mineralogical associations of basic and intermediate rocks; quartz and alkaline feldspar appear in the acidic rocks. Twenty-nine representative geochemical analyses carried out on the Early Cambrian volcanic rocks of Ouarzazate– Agdz, Kerkar –Ida– Ougnidif and Azegour –Wirgane (21) and Jbel Boho–Bou Azzer (8) are given in Table 2. Their compositions, based on the silica content, range from basic to acidic (44% , SiO2 , 72%). Like the Late Neoproterozoic igneous rocks, those of the Early Cambrian were affected by identical post-magmatic events and so the same conditions on the use of the low ionic potential elements are valid here. The Early Cambrian igneous rocks from Ouarzazate –Agdz, Jbel Kerkar –Ida Ougnidif, Azegour –Wirgane and Bou Azzer –Jbel Boho plotted on the Winchester & Floyd (1976) diagram show an enormous range of Nb/Y ratios, suggesting the existence of at least two groups: an alkaline group (Nb/Y . 1) for the Jbel Boho region, which plots in the alkaline basalt, trachyandesite and trachyte fields, and a sub-alkaline group (Nb/Y 1) from the other regions, which plots in the basalt and andesite –basaltic fields (Fig. 4b). Discrimination of the two Early Cambrian groups was achieved in two independent spider
MAGMATIC EVOLUTION OF NORTHERN WAC
-
100
-
-
10
-
-
-
Chilean basalt
5 TiO2
Sample/Chondrite (except K, Rb, P)
(a) 600
2
4
Sample/Chondrite (except K, Rb, P)
(b) 600
3 2 HPT -
100
1
-
-
-
LPT
P2O5
-
x
x
x -
x
0
-
0
1
-
-
x
x
x
x x
x
x
x x -
x
x x -
10 Parana Deccan N-MORB
2
Sample/Chondrite (except K, Rb, P)
(c) 800 100
10
1 Mt. Kenya basalt Mt. Kenya trachyte
.1 Ba
K Rb
Ta Nb
Ce La
Nd Sr
Sm P
Hf Zr
Tb Ti
Yb Y
Fig. 5. Normalized multi-elemental spectra for: (a) the Late Neoproterozoic rocks; (b) Early Cambrian sub-alkaline basalts; (c) Bou Azzer– Boho alkaline rocks. Inset in (b) shows TiO2 v. P2O5 projection of sub-alkaline basalts. Comparative reference profiles of Chilean calc-alkaline basalt and N-MORB (Pearce 1983), continental tholeiitic basalts of Parana´ and Deccan (Thompson et al. 1983) and Mt. Kenya basalt and trachyte (Price et al. 1985) are shown. Normalization values are from Thompson et al. (1984); Rb, K, P are from primitive mantle values of Sun (1980). Symbols as in Figure 4.
diagrams, which show their distinctive characteristics (Fig. 5b and c). The most primitive patterns of the sub-alkaline group show low fractionation of the elements from La to Y; there are pronounced negative anomalies in Nb–Ta and some
337
perturbation in the LILE is conspicuous. The most differentiated rock types display strong negative anomalies in Sr, suggesting plagioclase fractionation. The regular variation in the patterns (Nb– Ta behaviour) indicates they may correspond to the same igneous series (Fig. 5b). All the geochemical signatures suggest that these are the result of high-grade melting of a depleted mantle source with probable involvement of continental contamination. The two typical patterns of Parana´ (Brazil) and Deccan (India) continental tholeiites plotted in Figure 5b and the Early Cambrian sub-alkaline group have similar signatures. Despite the scatter of TiO2/P2O5 ratios of this group on the Bellieni et al. (1984) diagram, a high-P2O5 –TiO2 group (HPT) and a low-P2O5 –TiO2 group (LPT) can be distinguished (Fig. 5, inset). The rocks of the Azegour region, located on the northern side of the South Atlas Fault, belong to the LPT group. The rocks of all the other regions belong to the HPT group. According to Bellieni et al. (1984), this feature may indicate that the Azegour volcanic rocks were subjected to a stronger continental contamination than those in the other regions. The incompatible enrichment of the Early Cambrian alkaline group is observed on the spider diagram in Figure 5c. Their geochemical patterns and those of Mt. Kenya, usually cited as a typical example of alkaline intra-plate volcanism, are similar (Fig. 5c). The geochemical evolution of this alkaline series was controlled by plagioclase, clinopyroxene and clino-amphibole fractionation ´ lvaro et al. 2006). The magmatic source of (A this series was suggested to be the garnet´ lvaro et al. 2006). lherzolite-enriched mantle (A To constrain the processes involved in the genesis and differentiation of the sub-alkaline series of Early Cambrian volcanic rocks, the Zr/Y v. Ti/Y and Ce/Zr v. Zr ratios were plotted on the Crowley et al. (2000) diagram (Fig. 6b). The Early Cambrian series plot closely to the mantle array, unlike the Late Neoproterozoic magmatic series (Fig. 6a and b). The Early Cambrian rock projections suggest an evolution mostly controlled by fractional crystallization with minor continental upper crust contamination, unlike that of the Late Neoproterozoic series. In this series the role of the pelagic sediments and continental crust contamination seems to be more decisive in their differentiation (Fig. 6a and b).
Discussion and conclusion During the Late Neoproterozoic, a calc-alkaline magmatic series of potassic affinity evolved from basic to intermediate –acidic. This evolution was essentially controlled by fractional crystallization
338
H. EZZOUHAIRI ET AL.
(a) 10
+ + +
Pelagic Contaminant (r = 0.8)
PASC
8
4
AFC
+
0.3
0.2
+
AFC + UCC Contaminant (r = 0.2–0.8 range)
Average N-MORB (depleted)
2
0.1
+ Closed system F C
(b)
0.4
+ +
Zr / Y
6
0.5
Ce/Zr
Average OIB (enriched)
0 10
Pelagic Contaminant (r = 0.8)
PASC
8
0.0 0.5
Ce/Zr
Average OIB (enriched)
0.4
AFC
6 Zr / Y
0.3
4
0.2
AFC
UCC Contaminant (r = 0.2–0.8 range)
Average N-MORB (depleted)
2
0.1
Closed system FC
0
0
200
400
600
800 0
Ti /Y
100
200
300
0.0 400
Zr
Fig. 6. Zr/Y v. Ti/Y and Ce/Zr v. Zr projections of the Late Neoproterozoic rocks of Ouarzazate– Agdz region and of the Early Cambrian of the Anti-Atlas and High Atlas regions. Symbols as in Figure 4. PASC, pelagic assimilation and sediment contaminant; AFC, assimilation and fractional crystallization; UCC, upper continental crust.
and strong crustal contamination processes (assimilation –fractional crystallization AFC). However, the enrichment of the basic rocks in LILE, light REE (LREE) and K2O is related to the composition of its source having previously been modified by a subduction event (Youbi 1998; Ezzouhairi 2001). The partial melting of an ‘anomalous’ enriched mantle (by a previous subduction event) can generate such an enriched liquid (Kay & Kay 1989; Wilson 1994; MacInnes et al. 2001). The Early Cambrian magmatism, which follows the Late Neoproterozoic phase, began in the AntiAtlas on both sides of the AAMF (Ouarzazate and Kerdous regions) and extended later to other ´ lvaro et al. 2006). regions (El Archi et al. 2004; A Early Cambrian volcanism shows different geochemical signatures, suggesting different geodynamical environments. Except in the Jbel Boho area, the main geochemical features of the Early Cambrian volcanism correspond to continental
tholeiites and suggest emplacement in an extensional tectonic environment. This magmatism might be related to continental extension processes and rifting (Ait Ayad et al. 1998). Continental extension is also suggested by the polymictic and poorly sorted character and the thickness variations of the Basal Series conglomerate near the fault zones. The P2O5/TiO2 ratios of the continental tholeiites indicate the presence of two groups, the high-P2O5 –TiO2 group (HPT) and the low-P2O5 – TiO2 group (LPT). The HPT, which is very common in the Anti-Atlas and the High Atlas, corresponds to the main regional magmatism, whereas the LPT is located near the South Atlas Fault (SAF). Such features of these two similar groups (LPT and HPT), related to the Gondwanaland supercontinent fragmentation, were explained by different modifications of the mantle magmas sources by previous subduction events (Elliot 1975; Cox 1978).
Table 2. Selected chemical analyses of the Early Cambrian magmatic rocks AS-1
AS-2
AS-3
Jk1
JK3
JK5
JK8
IO6
IO7
IO8
IO9
BZ12
JB16
BZ13
SiO2 Al2O3 Fe2OT3 MnO CaO MgO Na2O K2 O TiO2 P2O5 LOI Total
46.08 15.02 15.84 0.11 8.15 2.93 3.67 0.8 3.53 0.34 3.08 99.55 0 17 56 38.0 143 4.0 101 0.2 0 4.3 0 13.0 30.4 4.6 24.4 6.20 2.50 3.40
45.89 15.12 15.6 0.17 5.69 4.53 3.73 1.73 3.36 0.34 3.24 99.4 0 42 61 38.0 144 5.0 73 0.3 0 4.3 0 10.7 26.4 4.2 21.3 5.80 2.30 3.30
46.63 14.23 17.25 0.15 6.06 3.39 4.4 1.27 3.26 0.33 2.6 99.57 0 21 96 35.0 137 5.0 161 0.3 0 4.1 0 9.4 24.2 3.9 19.3 5.20 1.90 3.20
45.6 13.85 14.55 0.2 6.87 7.68 4.16 0.62 2.14 0.41 3.28 99.36 0 12 211 47.0 172 6.7 254 0.4 0 5.2 305 13.0 33.0 0.0 26.5 7.20 3.00 3.18
48.4 12.5 15.2 0.15 3.39 7.35 5.7 0.65 2.48 0.31 3.11 99.24 0 11 143 34.5 192 6.0 345 0.4 0 5.8 290 5.2 14.0 0.0 13.7 4.18 1.20 2.90
46.6 14.95 14.25 0.15 5.35 6.76 4.42 0.96 2.78 0.38 2.69 99.29 0 10 197 44.0 178 7.5 442 0.5 0 5.3 390 13.2 34.0 0.0 26.0 7.05 2.75 3.53
47 12.75 14.85 0.2 4.75 8.42 4.85 0.67 2.3 0.35 3.2 99.34 0 11 336 41.0 164 7.8 255 0.5 0 4.9 306 12.6 31.0 0.0 23.0 6.30 1.98 3.05
47 11.7 14 0.36 5.5 10.5 3.97 0.47 1.92 0.26 3.6 99.2 0 10 208 27.0 130 4.5 186 0.3 0.6 3.9 241 7.8 22.0 0.0 18.0 4.90 1.32 2.02
47.5 13 16 0.26 6.02 6.3 4.7 0.47 2.15 0.27 3.32 99.99 0 13 400 42.0 135 5.3 220 0.3 0.7 4.1 330 10.1 27.0 0.0 22.0 5.85 2.45 3.28
49.4 11.4 15.4 0.23 4.78 7.25 5.5 0.9 2.54 0.38 2.11 99.8 0 18 192 43.0 252 11.5 547 0.7 0.9 7.6 274 13.6 38.5 0.0 28.0 7.00 2.43 3.00
48.7 14 14.3 0.21 6.55 4.91 5 0.82 2.25 0.32 2.77 99.8 0 35 445 43.0 280 10.0 292 0.6 1.1 8.4 288 13.2 38.0 0.0 28.0 7.50 1.92 3.24
45.15 15.91 11.97 0.07 4.67 9.65 1.6 3.42 2.86 0.47 3.72 99.49 0 46 118 20.0 167 26.0 916 11.0 33 38.0 14 0.0 0.0 0.0 0.0 0.00 0.00 13.00
45.64 15.11 14.18 0.18 6.19 4.23 4.72 2.72 2.9 0.68 3.1 99.66 0 40 88 20.0 160 25.0 484 6.0 16 14.0 311 0.0 0.0 0.0 0.0 0.00 0.00 7.00
46.21 15.46 9.65 0.07 6.04 9.87 2.22 2.99 2.99 0.6 3.47 99.58 0 29 90 22.0 176 31.0 387 8.0 20 20.0 4 0.0 0.0 0.0 0.0 0.00 0.00 7.00
Rb Sr Y Zr Nb Ba Ta Th Hf V La Ce Pr Nd Sm Eu Yb
MAGMATIC EVOLUTION OF NORTHERN WAC
Sample:
(Continued)
339
340
Table 2. Continued Sample:
Rb Sr Y Zr Nb Ba Ta Th Hf V La Ce Pr Nd Sm Eu Yb
JB8
JB1
JB3
JB6
Pz19
Pz20
Pz15
Pz18
TN-2
TN-3
TN-6
WT-1
WT-3
WT-6
55.27 17.52 10.62 0.05 1.32 2.25 4.62 3.93 1.75 0.71 1.53 99.57 0 70 39 41.0 415 85.0 806 8.0 18 20.0 1 0.0 0.0 0.0 0.0 0.00 0.00 8.00
62.93 15.79 7.86 0.04 0.56 0.64 2.78 7.91 0.38 0 0.82 99.72 0 152 15 66.0 784 149.0 321 9.0 20 24.0 2 0.0 0.0 0.0 0.0 0.00 0.00 8.00
64.15 12.64 11.1 0.06 0.17 3.63 0.24 5.37 0.56 0.07 1.55 99.54 0 90 13 61.0 960 135.0 559 10.0 28 24.0 4 0.0 0.0 0.0 0.0 0.00 0.00 5.00
66.48 14.94 5.7 0.04 0.56 1.91 2.61 5.98 0.4 0.05 1.17 99.84 0 94 19 90.0 1033 159.0 342 1.0 4 5.0 213 0.0 0.0 0.0 0.0 0.00 0.00 4.00
72.86 13.03 0.79 0.05 2.89 0.1 5.26 2.98 0.25 0.01 1.39 99.61 0 60 20 114.0 1821 196.0 131 2.0 10 8.0 215 0.0 0.0 0.0 0.0 0.00 0.00 2.00
48.65 14.51 11.58 0.18 10.74 7 2.83 0.85 2.15 0.21 0.83 99.53 0 22 367 30.5 144 4.0 418 0.2 3 2.4 0 4.5 11.8 1.8 10.0 3.60 1.49 3.12
44.59 13.17 12.72 0.25 15.94 6.61 1.15 1.46 2.08 0.2 1.37 99.54 0 30 339 38.6 136 3.0 220 0.2 3 2.6 0 5.8 16.2 2.7 14.4 5.10 1.96 4.01
58.8 18.27 3.85 0.12 2.84 2.74 8.06 0.29 0.79 0.14 3.73 99.63 0 7 569 10.4 98 4.0 160 0.2 0 1.8 0 6.4 16.9 2.7 10.7 2.50 0.77 1.05
51.33 19.88 9.07 0.1 3.26 4.94 6.57 0.42 1.15 0.22 2.57 99.51 0 12 705 12.6 139 13.0 177 0.8 0 2.1 0 12.9 29.2 3.8 15.5 3.40 0.94 1.39
45.89 16.47 15.21 0.2 4.56 3.26 5.48 0.75 4.18 0.74 2.95 99.69 0 18 339 57.0 343 15.0 428 0.9 5 9.0 292 17.0 59.0 0.0 47.0 13.00 0.00 0.00
43.69 15.89 14.64 0.24 8.47 3.31 5.31 0.34 3.65 0.66 3.38 99.58 0 7 244 44.0 291 13.0 212 0.8 0 6.0 249 8.0 51.0 0.0 35.0 7.00 0.00 0.00
46.81 15.57 13.6 0.23 5.91 4.35 5.05 0.43 3.68 0.73 3.27 99.63 0 10 289 48.0 331 14.0 287 0.8 0 7.0 232 9.0 47.0 0.0 37.0 16.00 0.00 0.00
51.86 16.43 15.23 0.12 1.93 2.1 5.87 0.77 2.72 1.14 1.65 99.82 0 18 209 39.0 282 19.0 241 1.1 0 6.0 138 42.0 103.0 0.0 64.0 14.00 0.00 0.00
52.66 17.96 9.1 0.15 1.82 4.57 5 0.51 4.6 0.54 2.95 99.86 0 11 249 26.0 267 16.0 173 1.0 0 9.0 231 16.0 55.0 0.0 27.0 3.00 0.00 0.00
47.83 19.32 13.06 0.11 1.26 3.17 5.54 1.43 4.65 0.75 2.75 99.87 0 32 115 36.0 321 24.0 410 1.4 0 6.0 190 34.0 86.0 0.0 51.0 10.00 0.00 0.00
AS, Ait Saoun basalts; JK, Jbel Kerkar basalts; IO, Ida Ougnidif basalts, BZ, Bou Azzer basalts; JB, Jbel Boho (JB16, basalt; JB2, JB8, JB1, JB3 and JB6, trachyandesites, trachytes and rhyolites); PZ, Azegour basalts; TN, Talat N’Yaacoub basalts (near de Wirgane village); WT, Wirgane basalts. The Lambert coordinates of these rocks are respectively: AS-1 (383.2; 413.8); AS-2 (383.3; 413.9); AS-3 (383.5; 414); BZ12 (348.8; 393.7); BZ13 (348.6; 393.8); JB16 (362.8; 384.4); JB2 (364; 380); JB8 (362.2; 382.4); JB1 (364; 378.8); JB3 (364.4; 380.6); JB6 (364.8; 380.4); PZ19 (221.9; 465.8); PZ20 (222; 466); PZ15 (222.2; 465.6); PZ18 (222.1; 465.7); TN-2 (238.6; 451); TN-3 (238.5; 450.6); TN-6 (238.6; 450.2); WT-1 (236; 458.5); WT-3 (236.2; 459); WT-6 (238.6; 450.2).
H. EZZOUHAIRI ET AL.
SiO2 Al2O3 Fe2OT3 MnO CaO MgO Na2O K2 O TiO2 P2O5 LOI Total
JB2
MAGMATIC EVOLUTION OF NORTHERN WAC
However, we think that more data are necessary to understand the LPT occurrence in such a region located on the north side of the South Atlas Fault. The Early Cambrian alkaline rocks of the top of Adoudou Formation and in the Lie-de-vin Formation, which crops out in a region near the SW side of the AAMF represent the latest magmatism. These rocks, related to partial melting of an enriched mantle, have patterns similar to those of Mt Kenya and the East African Rift rocks ´ lvaro et al. 2006). This source, deeper than that (A estimated for the previous Early Cambrian tholeiites, suggests a melting rate reduction over time. The evolution of tholeiitic basalts of the Early Cambrian sequences was interpreted as being related to a tectonic extensional event ending in an aborted continental rifting episode (Ait-Ayad et al. 1998; Ezzouhairi et al. 2005; ´ lvaro et al. 2006). The different geochemical A patterns observed near the major faults (SAF and AAMF) suggest that these structures played an important role in the Cambrian magmatism. We thank J. P. Lie´geois, B. Litvinovsky and K. Hefferan for highly constructive comments. Research on the Late Neoproterozoic and Early Palaeozoic magmatism has been supported by the GRICES/CNRST project ‘A comparative study of the Neoproterozoic– Cambrian transition between the Portuguese Ossa– Morena Zone and the Moroccan Anti-Atlas and Meseta: geologic and geochemical aspects, and geodynamical model’. Thanks are also due to the C. Geocieˆncias, Universidade de Coimbra for a short duration grant. This paper is a contribution to IGCP project 485 ‘Cratons, metacratons and mobile belts: keys from the West African craton boundaries, Eburnean versus Pan-African signature, magmatic, tectonic and metallogenic implications’.
References A IT A YAD , N., R IBEIRO , M. L., M ATA , J., F ERREIRA , P., E ZZOUHAIRI , H., C HARIF , A. & D IAS , R. 1998. Evolution du magmatisme cambrien en deux re´gions pe´rigondwanniennes: Azegour (Haut-Atlas) et Alter do Chao– Elvas (NE Alentejo). Comunicac¸oes do Instituto Geologico e Mineiro do Portugal, 84, 154– 157. A IT M ALEK , H., G ASQUET , D., B ERTRAND , J. M. & L ETERRIER , J. 1998. Ge´ochronologie U/Pb sur zircon de granitoides e´burne´ens et panafricains dans les botonnie`res prote´rozoiques d’Ighrem, du Kerdous et du Bas Draˆa (Anti Atlas occidental, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 327, 819–826. ´ LVARO , J. J., E ZZOUHAIRI , H., V ENNIN , E. ET AL . A 2006. The Early-Cambrian Boho Volcano of the El Graara massif, Morocco: Petrology, geodynamical setting and coeval sedimentation. Journal of African Earth Sciences, 44, 396–410. B ELLIENI , G., B ROTZU , P., C OMIN -C HIARAMONTI , P., E RNESTO , M., M ELFI , A., P ACCA , I. G. & P ICCIRILLO , E. M. 1984. Flood basalt to rhyolite
341
suite in southern Parana Plateau (Brazil): palaeomagnetism, petrogenesis and geodynamical implications. Journal of Petrology, 25, 579– 618. B ENSSAOU , M. 2005. Le Rift cambrien infe´rieur de L’Anti-Atlas (Maroc): Remplissage se´dimentaire, ge´odynamique et palae´o-ge´ographie. PhD thesis, University Ibn Zohr, Agadir. B ENSSAOU , M. & H AMOUMI , N. 2001. L’Anti-Atlas occidental du Maroc: e´tude se´dimentologique et reconstitutions pale´oge´ographiques au Cambrien infe´rieur. Journal of African Earth Sciences, 32, 351– 372. B UGGISCH , W. & F LU¨ GEL , E. 1988. The Precambrian/ Cambrian boundary in the Anti-Atlas. Discussion and new results (Morocco). Lecture Notes on Earth Sciences, 15, 81–90. C HARLOT , R. 1982. Caracte´risation des e´ve´nements e´burne´ens et panafricains dans l’Anti-Atlas marocain: apport de la me´thode ge´ochronologique Rb–Sr. Notes et Me´moires du Service Ge´ologique du Maroc, 313. C HEBBAA , B. 1996. Me´talloge´nie du cuivre associe´ aux roches volcaniques d’aˆge Pre´cambrien II supe´rieur dans l’Anti-Atlas marocain. PhD thesis, University of Lausanne. C HOUBERT , G. 1952. Le volcan ge´orgien de la re´gion d’Alougoum (Anti-Atlas). Comptes Rendus de l’Acade´mie des Sciences, 234, 350–352. C HOUBERT , G. 1963. Histoire ge´ologique du Pre´cambrien de l’Anti Atlas. Notes et Me´moires du Service Ge´ologique du Maroc, 162. C HOUBERT , G. & F AURE -M URET , A. 1970. Livret-guide de l’excursion ‘Anti-Atlas occidental et central’ du Colloque international sur les corre´lations du Pre´cambrien. Notes et Me´moires du Service Ge´ologique du Maroc, 229. C LAUER , N. 1974. Utilisation de la me´thode Rb–Sr pour la datation d’une schistosite´ de se´diments peu me´tamorphise´s: applications au Pre´cambrien II de la boutonnie`re de Bou Azzer–El Graara (Anti-Atlas, Maroc). Earth and Planetary Science Letters, 22, 404–412. C LAUER , N. 1976. Ge´ochimie isotopique du strontium des milieux se´dimentaires: application a` la ge´ochronologie de la couverture du craton Ouest-Africain. Sciences Ge´ologiques, 45, 1– 256. C LAUER , N., C ABY , R., J EANNETTE , D. & T ROMPETE , R. 1982. Geochronology of sedimentary and metasedimentary rocks of the West African craton. Precambrian Research, 18, 53– 71. C OMPSTON , W., W ILLIAMS , J. L., K IRSCHVINK , J. L., Z HANG , Z. & M A , G. 1992. Zircon U–Pb ages for the Early Cambrian time scale. Journal of the Geological Society, London, 127, 319–332. C OX , K. G. 1978. Flood basalts, subduction and the break-up of Gonwanaland. Nature, 274, 47–49. C ROWLEY , Q. G., F LOYD , P. A., W INCHESTER , J. A., F RANKE , W. & H OLLAND , J. G. 2000. Early Palaeozoic rift-related magmatism in Variscan Europe: fragmentation of the Armorican Terrane Assemblage. Terra Nova, 12, 171–180. D ESTOMBES , J., H OLLARD , H. & W ILLEFERT , S. 1985. Lower Palaeozoic rocks of Morocco. In: H OLLAND , C. H. (ed.) Lower Palaeozoic Rocks of the World, Vol. 4. Lower Palaeozoic of Northwestern and West– central Africa. Wiley, New York, 157–184.
342
H. EZZOUHAIRI ET AL.
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, 65–78. D UCROT , J. & L ANCELOT , J. R. 1977. Proble`me de la limite Pre´cambrien –Cambrien: e´tude radiochronologique par la me´thode U/Pb sur zircon du volcan du Jbel Boho. Canadian Journal of Earth Sciences, 14, 1771–1777. E L A RCHI , A., E L H OUICHA , M., J OUHARI , A. & B OUABDELLI , M. 2004. Is the Cambrian basin of the western High-Atlas (Morocco) related either to a subduction zone of a major shear zone? Journal of African Earth Sciences, 39, 311 –318. E LLIOT , D. H. 1975. Tectonics in Antarctica: a review. American Journal of Science, 275A, 45– 106. E NNIH , N. & L IE´ GEOIS , J. P. 2001. The Moroccan Anti-Atlas: 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 RRAMI , E., L IE´ GEOIS , J. P., L ADURON , D. & E NNIH , N. 2002. The Pan-African post-collisional high-K calc-alkaline granitoids (Eastern Saghro, Anti-Atlas, Morocco), witness for a metacratonic margin? In: ENNIH , N. (ed.) Abstracts of the 19th Colloquium of African Geology, El Jadida, Morocco, 85. E ZZOUHAIRI , H. 2001. Le magmatisme post-collisionnel panafricain (tardi a` post-oroge´nique) des re´gions d’Aghbalou, Sidi Flah Bouskour et Oued Imini (Ouarzazate, Anti Atlas central, Maroc). Lithostratigraphie, ge´ochimie, pe´trogene`se et contexte ge´odynamique. PhD thesis, University Chouaı¨b Doukkali, El Jadida. E ZZOUHAIRI , H., R IBEIRO , M. L., F ERREIRA , P., A IT A YAD , N., C HARIF , A. & R AMOS , F. 1998. Magmatisme pre´cambrien de la re´gion d’Aghbalou–Oued Imini (Anti-Atlas central du Maroc): nature ge´ochimique et quelques aspects significatifs. Comunicac¸oes do Instituto Geologico e Mineiro do Portugal, 84, 178– 181. E ZZOUHAIRI , H., R IBEIRO , M. L., C HARIF , A., R AMOS , F., A IT A YAD , N., M OREIRA , M. E. & C OKE , C. 2005. Le magmatisme Cambrien infe´rieur de Jbel Boho– Bou Azzer: quelques aspects pe´trographique et ge´ochimiques (Anti Atlas, Maroc). In: I KENNE , M. (ed.) 4th International Colloquium of Magmatism, Metamorphism and Associated Mineralization, University Ibn Zohr, Agadir, Maroc, 103–104. F EKKAK , A., B OUALOUL , M., B ADRA , L., A MENZOU , M., S AQUAQUE , A. & E L A MRANI , I. E. 2000. Origine et contexte ge´otectonique des de´poˆts de´tritiques du Groupe ne´oprote´rozoique infe´rieur de Kelaaˆt M’gouna (Anti-Atlas oriental). Maroc. Journal of African Earth Sciences, 30, 295 –311. G ASQUET , D., L EVRESSE , G., C HEILLETZ , A., A ZIZI -S AMIR , R. & M OUTTAQI , A. 2005. Contribution to a geodynamical 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 EYER , G. 1990. Proposal of formal lithostratigraphical units for the Terminal Proterozoic to Early Middle
Cambrian of southern Morocco. Newsletter on Stratigraphy, 22, 87– 109. H EFFERAN , K., K ARSON , J. & S AQUAQUE , A. 1992. Proterozoic collisional basins in a Pan-African suture zone, Anti-Atlas Mountains, Morocco. Precambrian Research, 54, 295– 319. H EFFERAN , K. P., A DMOU , H., K ARSON , J. A. & S AQUAQUE , A. 2000. Anti-Atlas (Morocco) role in Neoproterozoic Western Gondwana reconstitution. Precambrian Research, 103, 89– 96. 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. 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 Bleı¨da granodiorite, Anti-Atlas, Morocco: implication 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 the Late Precambrian intrusion, metamorphism, and tectonism in the Anti-Atlas Mountains. Journal of Geology, 113, 439 –450. K AY , R. W. & K AY , S. M. 1989. Recycled continental crustal components in Aleutian arc magmas, implications for crustal growth and mantle heterogeneity. In: H ART , S. R. & G ULEN , L. (eds) Crust/Mantle Recycling at Convergence Zones. Kluwer, Dordrecht, 258, 145– 161. L ANDING , E., B OWRING , S. A., D AVIDEK , K. L., W ETROP , S. R., G EYER , G. & H ELDMAIER , W. 1998. Duration of the Early Cambrian: U–Pb ages of volcanic ashes from Avalon and Gondwana. Canadian Journal of Earth Sciences, 35, 329–338. L EBLANC , M. & L ANCELOT , J. 1980. Interpre´tation ge´odynamique du domaine panafricain de l’Anti Atlas (Maroc) a` partir de donne´es ge´ologiques et ge´ochronologiques. Canadian Journal of Earth Sciences, 17, 142–155. L EVRESSE , G. 2001. Contribution a` l’e´tablissement d’un mode`le ge´ne´tique des gisements d’Imiter (Ag–Hg) Bou-Madine (Pb–Zn–Cu–Au) et Bou Azzer (Co–Ni– As–Au–Ag) dans l’Anti-Atlas marocain. PhD thesis, Institut National Polytechnique de Lorraine, Nancy. M AALOUF , A. C., S CHRAG , D. P., C ROWLEY , J. L. & B OWRING , S. A. 2005. An expanded record of Early Cambrian carbon cycling for the Anti-Atlas margin, Morocco. Canadian Journal of Earth Sciences, 42, 2195– 2216. M C I NNES , B. I. A., G RE´ GOIRE , M., B INNS , R. A., H ERZIG , P. & H ANNINGTON , M. 2001. Hydrous metasomatism of oceanic sub-arc mantle, Lihir, Papua New Guinea: petrology and geochemistry of fluid-metasomatised mantle wedge xenoliths. Earth and Planetary Science Letters, 188, 169– 183. M IFDAL , A. & P EUCAT , J. 1985. Datation U– Pb et Rb–Sr du volcanisme acide de l’Anti-Atlas marocain et du socle sous-jacent dans la re´gion de Ouarzazate. Apport au proble`me de la limite Pre´cambrien– Cambrien. Bulletin de Sciences Ge´ologiques, 38, 185–200.
MAGMATIC EVOLUTION OF NORTHERN WAC M RINI , Z. 1993. Chronologie (Rb– Sr; U– Pb); trac¸age isotopique (Sr–Nd –Pb) des sources des roches magmatiques e´burne´ennes, panafricaines et hercyniennes du Maroc. PhD thesis, University Cadi Ayyad, Marrakech. P EARCE , J. A. 1983. Role of the sub-continental lithosphere in magma genesis at active continental margins. In: H AWKESWORTH , C. J. & N ORRY , M. J. (eds) Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, 230 –249. P RICE , R. C., J OHNSON , R. W., G RAY , C. M. & F REY , F. A. 1985. Geochemistry of phonolites and trachytes from the summit region of Mt. Kenya. Contributions to Mineralogy and Petrology, 89, 394–409. S AQUAQUE , A., A DMOU , H., K ARSON , J. A., 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 OULAIMANI , A., B OUABDELLI , M. & P IQUE´ , A. 2003. L’extension continentale au Ne´o-Prote´rozoique supe´rieur– Cambrien infe´rieur dans l’Anti-Atlas (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, 174, 83–92. 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 (Anti-Atlas occidental, Maroc). Comptes Rendus Ge´oscience, 336, 1433–1441. S OULAIMANI , A., J AFFAL , M., M AACHA , L., K CHIKACH , A., N AJINE , A. & S AIDI , A. 2006. Mode´lisation magne´tique de la suture ophiolitique de Bou Azzer–El Graara (Anti-Atlas central, Maroc). Implications sur la reconstitution ge´odynamique panafricaine. Comptes Rendus Ge´oscience, 338, 153–160. S UN , S. S. 1980. Lead isotopic study of young volcanic rocks from mid-ocean ridges, ocean islands and island arcs. Philosophical Transactions of the Royal Society of London, Series A, 297, 409–445.
343
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 HOMPSON , R. N., M ORRISON , M. A., D ICKIN , A. P. & H ENDRY , G. L. 1983. Continental flood basalts . . . arachnids rule OK? In: H AWKESWORTH , C. J. & N ORRY , M. J. (eds) Continental Basalts and Mantle Xenoliths, Shiva, Nantwich, 158–185. T HOMPSON , R. N., M ORRISON , M. A., H ENDRY , G. L. & P ARRY , S. J. 1984. An assessment of the relative roles of crust and mantle in magma genesis. Philosophical Transactions of the Royal Society of London, Series A, 310, 549 –590. W ALSH , G. J., A LEINIKOFF , J. N., B ENZIANE , F., Y AZIDI , A. & A MSTRONG , T. R. 2002. Pb–zircon geochronology of the Palaeoproterozoic Tagragra de Tata inlier and its Neoproterzoic cover, western Anti-Atlas, Morocco. Precambrian Research, 117, 1– 20. W ILSON , M. 1994. Igneous Petrogenesis. A Global Tectonic Approach. Chapman & Hall, London. W INCHESTER , J. A. & F LOYD , P. A. 1976. Geochemical magma type discrimination; application to altered and metamorphosed basic igneous rocks. Earth and Planetary Science Letters, 28, 459–469. Y OUBI , N. 1998. Le volcanisme post ‘collisionnel’: un magmatisme intraplaque relie´ a` des panaches mantelliques. Etudes volcanologique et ge´ochimique. Exemples d’application dans le Ne´oprote´rozoique terminal (PIII) de l’Anti Atlas et le permien du Maroc. PhD thesis, University Caddi Ayyad, Marrakech. Z AHOUR , G. 2001. Le Ne´oprote´rozoı¨que terminal de la boutonnie`re de Toubkal (Haut-Atlas Occidental) et de Siroua (Anti-Atlas Central): un exemple de volcanisme intraplaque continental associe´ a` un volcanisme calco-alcalin post-collisionnel. PhD thesis, University Hassan II-Mohammedia, Casablanca.
Neoproterozoic –early Palaeozoic tectonostratigraphy and palaeogeography of the peri-Gondwanan terranes: Amazonian v. West African connections R. DAMIAN NANCE1, J. BRENDAN MURPHY2, ROB A. STRACHAN3, J. DUNCAN KEPPIE4, GABRIEL GUTIE´RREZ-ALONSO5, ´ NDEZ-SUA ´ REZ6, CECILIO QUESADA7, ULF LINNEMANN8, JAVIER FERNA RICHARD D’LEMOS9 & SERGEI A. PISAREVSKY10 1
Department of Geological Sciences, 316 Clippinger Laboratories, Ohio University, Athens, OH 45701, USA (e-mail:
[email protected])
2
Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, B2G 2W5, Canada 3
School of Earth & Planetary Sciences, University of Portsmouth, Burnaby Road, Portsmouth PO1 3QL, UK 4
Instituto de Geologı´a, Universidad Nacional Autonoma de Me´xico, 04510, Me´xico D.F., Mexico
5
Departmento de Geologı´a, Universidad de Salamanca, 33708 Salamanca, Spain
6
Departmento de Petrologı´a y Geoquı´mica, Universidad Complutense, 28040 Madrid, Spain 7
Instituto Geolo´gico y Minero de Espan˜a, Direccion de Geologı´a y Geofisica, Rio Rosas 23, 28003 Madrid, Spain
8
Staatliche Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Ko¨nigsbru¨cker Landsstrasse 159, D-01109 Dresden, Germany 9
Deers Cottage, Aston View, Somerton OX25 6NP, UK
10
Tectonics Special Research Centre, University of Western Australia, Crawley, WA, Australia Abstract: Within the Appalachian –Variscan orogen of North America and southern Europe lie a collection of terranes that were distributed along the northern margin of West Gondwana in the late Neoproterozoic and early Palaeozoic. These peri-Gondwanan terranes are characterized by voluminous late Neoproterozoic (c. 640–570 Ma) arc magmatism and cogenetic basins, and their tectonothermal histories provide fundamental constraints on the palaeogeography of this margin and on palaeocontinental reconstructions for this important period in Earth history. Field and geochemical studies indicate that arc magmatism generally terminated diachronously with the formation of a transform margin, leading by the Early–Middle Cambrian to the development of a shallow-marine platform– passive margin characterized by Gondwanan fauna. However, important differences exist between these terranes that constrain their relative palaeogeography in the late Neoproterozoic and permit changes in the geometry of the margin from the late Neoproterozoic to the Early Cambrian to be reconstructed. On the basis of basement isotopic composition, the terranes can be subdivided into: (1) Avalonian-type (e.g. West Avalonia, East Avalonia, Meguma, Carolinia, Moravia–Silesia), which developed on juvenile, c. 1.3– 1.0 Ga crust originating within the Panthalassa-like Mirovoi Ocean surrounding Rodinia, and which were accreted to the northern Gondwanan margin by c. 650 Ma; (2) Cadomian-type (e.g. North Armorican Massif, Ossa– Morena, Saxo-Thuringia, Moldanubia), which formed along the West African margin by recycling ancient (c. 2.0–2.2 Ga) West African crust; (3) Ganderian-type (e.g. Ganderia, Florida, the Maya terrane and possible the NW Iberian domain and South Armorican Massif), which formed along the Amazonian margin of Gondwana by recycling Avalonian and older Amazonian basement; and (4) cratonic terranes (e.g. Oaxaquia and the Chortis block), which represent displaced Amazonian portions of cratonic Gondwana. These contrasts imply the existence of fundamental sutures between these terranes prior to c. 650 Ma. Derivation of the Cadomian-type terranes from the West African craton is further supported by detrital zircon
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 345–383. DOI: 10.1144/SP297.17 0305-8719/08/$15.00 # The Geological Society of London 2008.
346
R. D. NANCE ET AL. data from their Neoproterozoic–Ediacaran clastic rocks, which contrast with such data from the Avalonian- and Ganderian-type terranes that suggest derivation from the Amazonian craton. Differences in Neoproterozoic and Ediacaran palaeogeography are also matched in some terranes by contrasts in Cambrian faunal and sedimentary provenance data. Platformal assemblages in certain Avalonian-type terranes (e.g. West Avalonia and East Avalonia) have cool-water, highlatitude fauna and detrital zircon signatures consistent with proximity to the Amazonian craton. Conversely, platformal assemblages in certain Cadomian-type terranes (e.g. North Armorican Massif, Ossa–Morena) show a transition from tropical to temperate waters and detrital zircon signatures that suggest continuing proximity to the West African craton. Other terranes (e.g. NW Iberian domain, Meguma) show Avalonian-type basement and/or detrital zircon signatures in the Neoproterozoic, but develop Cadomian-type signatures in the Cambrian. This change suggests tectonic slivering and lateral transport of terranes along the northern margin of West Gondwana consistent with the transform termination of arc magmatism. In the early Palaeozoic, several peri-Gondwanan terranes (e.g. Avalonia, Carolinia, Ganderia, Meguma) separated from West Gondwana, either separately or together, and had accreted to Laurentia by the Silurian– Devonian. Others (e.g. Cadomian-type terranes, Florida, Maya terrane, Oaxaquia, Chortis block) remained attached to Gondwana and were transferred to Laurussia only with the closure of the Rheic Ocean in the late Palaeozoic.
Scattered throughout the Palaeozoic orogens of the circum-North Atlantic (Fig. 1) lie a collection of arc-related terranes that are generally considered to have occupied positions along the northern margin of West Gondwana during the late Neoproterozoic and early Palaeozoic (e.g. Strachan & Taylor 1990; Nance & Thompson 1996; Murphy et al. 2002a; Do¨rr et al. 2004; Linnemann et al. 2007a). Although the precise location that each of these so-called peri-Gondwanan terranes occupied along this active margin is uncertain, a variety of field and analytical data suggest that, whereas some lay proximal to Amazonia, others were built upon West African basement (e.g. Nance & Murphy 1994, 1996). The palaeogeography of these terranes in the late Neoproterozoic, therefore, plays a central role in defining the margins of the West African craton during this time interval. Furthermore, because these data require the periGondwanan terranes to have faced an open ocean during the late Neoproterozoic and early Palaeozoic (e.g. Murphy et al. 2000; Nance et al. 2002; Keppie et al. 2003a), their tectonothermal records provide fundamental constraints on palaeocontinental reconstructions for this time interval, which was a critical period in Earth history marked by widespread orogeny and profound changes in ocean biology and chemistry (e.g. Knoll 1992; Hoffman et al. 1998; James et al. 2005). In this paper, we summarize the tectonostratigraphic records of these subduction-related complexes and review their relative palaeogeography in the late Neoproterozoic–early Palaeozoic as revealed primarily by: (1) the isotopic character of their basement (e.g. Nance & Murphy 1994, 1996), supplemented by: (2) the provenance of their sedimentary successions as determined from detrital mineral studies (e.g. Gutie´rrez-Alonso
et al. 2005), (3) the affinities of their fauna (e.g. ´ lvaro et al. 2003; Landing 2005), and (4) the A results of palaeomagnetic studies (e.g. McNamara et al. 2001; Mac Niocaill et al. 2002). We then examine the implications of the inferred palaeogeography for late Neoproterozoic –early Palaeozoic continental reconstructions. The time scale used is that of Gradstein et al. (2004).
Geological setting Characteristics and location The peri-Gondwanan terranes occur as exotic blocks within the Appalachian belt of North America, the Cordillera of Middle America, and the Variscan belt of southern and central Europe (Fig. 1), having been either detached or transferred from Gondwana in the Palaeozoic. For the late Neoproterozoic –early Palaeozoic, however, a wealth of tectonostratigraphic, palaeomagnetic, faunal and isotopic data place these terranes along the northern margin of West Gondwana at considerable latitudinal distance from contemporary Laurentia and Baltica (e.g. Cowie 1974; Theokritoff 1979; Johnson & Van der Voo 1986; Cocks & Fortey, 1988, 1990; Van der Voo, 1988; Nance et al. 1991; McKerrow et al. 1992; Nance & Murphy 1994, 1996; Cocks 2000; McNamara et al. 2001; Murphy et al. 2001; Mac Niocaill et al. 2002). Although important differences exist between them, the peri-Gondwanan terranes collectively record a protracted history of subduction beneath this margin that begins at about 760 Ma and terminates diachronously with the progressive development of a transform system from c. 610 Ma to c. 540 Ma, or continues into the
THE PERI-GONDWANAN TERRANES
347
Fig. 1. Location of peri-Gondwanan terranes on Early Mesozoic (Pangaea A) reconstruction of the circum-North Atlantic region (modified from Nance & Murphy 1994; Keppie & Ortega-Gutie´rrez 1999; Weil et al. 2001). CBI, Cape Breton Island; CH, Chortis block; Cp, Chiapas; F, Floresta; G, Garzo´n; Gu, Guajira; H, Huiznopala; M, Mixtequita; MA, Me´rida Andes; MSZ, Moravo-Silesian zone; MZ, Moldanubian zone; NAM, North Armorican Massif; No, Novillo; NWI, NW Iberian domain; OMZ, Ossa-Morena zone; Ox, Oaxacan Complex; Q, Quetame; S, Santander; SAM, South Armorican Massif; SM, Santa Marta; STZ, Saxo-Thuringian zone. OAXAQUIA comprises Huiznopala, Mixtequita, Novillo and Oaxacan Complex.
Cambrian. Neoproterozoic tectonothermal activity resembles that of modern Andean or island arc-back arc complexes and is characterized by abundant calc-alkaline volcanic rocks, cogenetic plutons, and synorogenic sedimentation and deformation associated with the opening and closing of arc-related basins (e.g. Murphy & Nance 1989; Keppie et al. 1991, 2003a; Nance et al. 1991, 2002; Murphy et al. 1999a, 2004a). The Neoproterozoic history of Andean-type arc magmatism generally concludes with the deposition of Cambrian platformal or passive margin sequences containing Gondwanan fauna. In North America, peri-Gondwanan terranes now occupy much of the eastern flank of the Appalachian orogen (Fig. 2), where they include West
Avalonia, Ganderia, Meguma and Carolinia, and the Suwanee terrane of the Florida subsurface (e.g. O’Brien et al. 1983; Heatherington et al. 1996; van Staal et al. 1998; Hibbard 2000; Hibbard et al. 2002; Nance et al. 2002; Murphy et al. 2004a). In Middle America (Fig. 3), periGondwanan terranes also make up Oaxaquia and the Maya terrane (Yucatan block) in Mexico (e.g. Keppie & Ortega-Gutie´rrez 1999; Keppie 2004), and include the Chortis block of Honduras and Guatamala (e.g. Keppie & Ramos 1999). In southern and central Europe (Fig. 4), periGondwanan terranes make up: (1) East Avalonia, which underlies southern Britain (Midland craton), the Brabant Massif of Belgium and, most probably, the Moravo-Silesian zone
Fig. 2. Peri-Gondwanan terranes of within the Appalachian orogen of eastern North America (modified from Hibbard et al. 2007). CBI, Cape Breton Island; G, Goochland terrane.
348
R. D. NANCE ET AL.
Fig. 3. Cratonic peri-Gondwanan terranes of Middle America (modified from Keppie 2004).
Fig. 4. Peri-Gondwanan terranes of southern and central Europe (modified from Franke 1989; Quesada 2006; Linnemann et al. 2007b). An, Anglesey; BCSZ, Badajos–Co´rdoba shear zone; CIZ, Central Iberian zone; CZ, Cantabrian zone; MC, Midland craton; MSL, Menai Strait Line; MZ, Moldanubian zone; NAM, North Armorican Massif; NASZ, North Armorican shear zone; NWI, NW Iberian domain; OMZ, Ossa–Morena zone; SAM, South Armorican Massif; SASZ, South Armorican shear zone; SMZ, Moravo-Silesian zone; STZ, Saxo-Thuringian zone; WALZ, West Asturian–Leonese zone.
THE PERI-GONDWANAN TERRANES
(Brunovistulian) of the Bohemian Massif (e.g. Finger et al. 2000; Pharaoh & Carney 2000; Sintubin et al. 2002), and (2) the Neoproterozoic– early Palaeozoic rocks (collectively referred to as Cadomian basement) that is preserved in the North and South Armorican Massifs (Cadomia) of northwestern France (e.g. Strachan et al. 1990, 1996a; Egal et al. 1996), the Saxo-Thuringian and Moldanubian (Tepla´ –Barrandian) zones of the Bohemian Massif (Bohemia) in Germany and the Czech Republic (e.g. Do¨rr et al. 2002; Linnemann & Romer 2002; Linnemann et al. 2004), and the NW Iberian domain and Ossa–Morena zone of the Iberian Massif (Iberia) in Spain and Portugal (e.g. Quesad a 1990, 1991; Ferna´ndez-Sua´rez et al. 2000, 2002a, b). Remnants of Cadomian basement also occur as inliers in the Alps, the Carpathians and Turkey (e.g. Neubauer 2002; von Raumer et al. 2002), although the palaeogeography of these remnants relative to other peri-Gondwanan terranes is poorly constrained (e.g. Neubauer 2002).
Palaeogeographical constraints Reconstruction of the late Neoproterozoic palaeogeography of the peri-Gondwanan terranes makes use of various lines of evidence, which also testify to the exotic nature of these terranes with respect to their present locations. Chief among these constraints are indicators of basement composition and sources, supplemented by the faunal affinities of sedimentary strata and palaeomagnetic data. Basement composition and sources. Exposure of undisputed basement in the peri-Gondwanan terranes occurs only in the North Armorican Massif of Cadomia, where it takes the form of the c. 2.1 Ga Icart Gneiss (e.g. Samson & D’Lemos 1998), the Moldanubian zone of Bohemia, where it takes the form of the c. 2.1 Ga Svetlik gneiss (Kro¨ner et al. 1988; Wendt et al. 1993), and the Moravo-Silesian zone of Bohemia, where it takes the form of the c. 1.38 Ga Dobra orthogneiss (Gebauer & Friedl 1994; Friedl et al. 2004). Basement of c. 1 Ga age is exposed in Oaxaquia and the Chortis block of Middle America (Keppie & Ortega-Gutie´rrez 1995, 1999), and, in Oaxaquia, is locally overlain unconformably by uppermost Cambrian –Lower Ordovician or Silurian rocks containing a Gondwanan fauna (Robison & Pantoja-Alor 1968, revised by Shergold 1975; Boucot et al. 1997; Stewart et al. 1999; Landing et al. 2006). However, these terranes lack Neoproterozoic arc magmatism and are thought to represent cratonic fragments of Amazonia that lay inboard of the arc (e.g. Keppie et al. 2001, 2003b, 2006).
349
Despite the lack of exposed basement, clues to basement tectonothermal evolution in a given periGondwanan terrane are provided by (1) the Sm –Nd isotopic composition and U –Pb xenocrystic zircon ages of younger, crustally derived felsic igneous rocks that represent melt fractions extracted from the basement, and (2) U –Pb detrital zircon and 40 Ar– 39Ar detrital muscovite data that give an indication of basement sources (e.g. Nance & Murphy 1994, 1996; Keppie et al. 1998; Gutie´rrez-Alonso et al. 2005). Because of the mineral’s resistance, U –Pb detrital zircon data tend to provide a broad fingerprint of sedimentary provenance that includes both proximal and distal source areas, especially if sediment transport was extensive. However, muscovite is considerably less robust and so is unlikely to survive more than one sedimentary cycle or even long waterborne transport. It is therefore considered to constrain proximal source areas containing ‘primary’ muscovite-bearing rocks (Haines et al. 2004). Based on the isotopic character of their basement, the peri-Gondwanan terranes can be grouped into four broad ‘palaeogeographical’ categories (Fig. 1): (1) those, such as Avalonia, that evolved upon juvenile crust of c. 1 Ga mantle extraction age; (2) those, such as Cadomia, that evolved upon ancient continental crust like that of the West African craton; (3) those, such as Ganderia, that evolved upon ancient continental crust like that of the Amazon craton; and (4) cratonic terranes, such as Oaxaquia, which lack a Neoproterozoic arc component and represent displaced portions of cratonic Gondwana (e.g. Murphy et al. 2004a). Faunal affinities. The Gondwanan affinity of those peri-Gondwanan terranes that preserve an early Palaeozoic sedimentary record is evident from the faunal data of Cocks & Fortey (1988, 1990) and Fortey & Cocks (2003). However, Landing (1996, 2005) and Geyer & Landing (2001) pointed out that important faunal differences exist during the Cambrian between the Avalonian terranes of North America and Britain, which record cool, high-latitude fauna, and Cadomia and Iberia, which record an evolution of tropical to temperate fauna matching that of West Africa. Similarly, Samson et al. (1990) identified differences between the Cambrian fauna of Avalonia and that of Carolinia. These distinctions are important because they provide critical information on the distribution of the peri-Gondwanan terranes in the Cambrian and, by comparing this distribution with their inferred Neoproterozoic locations, can be used to document relative movement between the terranes during Neoproterozoic – Cambrian time.
350
R. D. NANCE ET AL.
Palaeomagnetic data. Unequivocal evidence of the proximity of the peri-Gondwanan terranes to Gondwana from the Neoproterozoic to the Early Ordovician, and of their considerable latitudinal separation from contemporary Laurentia and Baltica, is provided by palaeomagnetic data (e.g. Johnson & Van der Voo 1986; Van der Voo 1988; McNamara et al. 2001; Thompson et al. 2006). At present, however, these data are not sufficient to resolve positional differences between these terranes (e.g. Mac Niocaill et al. 2002; Murphy et al. 2002b). The principal reason for this uncertainty is the lack of reliable palaeomagnetic data for Amazonia or West Africa for this time interval. Consequently, because most reconstructions (e.g. Cawood & Pisarevsky 2006) juxtapose the southwestern margin of Amazonia and the southeastern margin of Laurentia during this period (a configuration supported by palaeomagnetic data at c. 1.2 Ga for Amazonia in an inverted position; Tohver et al. 2002), the positions of Amazonia and West Africa are inferred from Laurentian palaeomagnetic poles (e.g. Dalziel 1991, 1992, 1994, 1997; Hoffman 1991; Weil et al. 1998). However, the palaeolatitude of Laurentia (high v. low) from c. 600 to 540 Ma is controversial (e.g. Tanczyk et al. 1987; Symons & Chiasson 1991), although a low-latitude position (at least at c. 550 Ma) is favoured by recent data (McCausland & Hodych 1998; Hodych et al. 2004). A connection between Laurentia and Gondwana at this time has also been called into question by data that suggest they were separate. These data include the suggestion that the c. 1 Ga Arequipa–Antofalla terrane of Peru and Chile, which borders the southwestern margin of Amazonia, may have been a microcontinent flanked by oceans (Loewy et al. 2004), and the possible existence of a Neoproterozoic orogen (the Maran˜on belt of Peru and its continuation into the Tucavaca belt of Bolivia: Ramos & Aleman 2000; V. Ramos, pers. commun.) between the Arequipa Massif and the Amazon craton.
Magmatic and tectonostratigraphic history The geology of the peri-Gondwanan terranes has been described in several compilations (e.g. O’Brien et al. 1983; D’Lemos et al. 1990; Strachan & Taylor 1990; Nance et al. 1991, 2002; Nance & Thompson 1996; Eguı´luz et al. 2000; Gutie´rrez-Alonso et al. 2003; Keppie et al. 2003a; Drost et al. 2004; Linnemann et al. 2004; Murphy et al. 2004a; Hibbard et al. 2007) and only brief summaries of the magmatic and tectonostratigraphic records of individual terranes are provided here. A review of the timeequivalent (Pan-African) tectonic evolution of West Africa has been provided by Deynoux et al. (2006).
Avalonia The peri-Gondwanan province of Avalonia, which we define in the more restrictive sense (Avalon Zone sensu stricto) of O’Brien et al. (1996) rather than the more broadly defined Avalon Zone of van Staal et al. (1996), is divided by the Atlantic Ocean into West Avalonia in New England and Atlantic Canada (Fig. 2), and East Avalonia in southern Britain, the subsurface Brabant Massif of Belgium and Rhenish Massif of NW Germany and, probably, the Moravo-Silesian zone–Brunovistulian of the Bohemian Massif (Fig. 4). These regions are characterized by a Cambrian– Ordovician overstep sequence containing a distinctive Acado-Baltic (Avalonian) fauna (e.g. Theokritoff 1979; Keppie 1993; Landing 1996, 2005) and formed a continuous tectonostratigraphic belt prior to the opening of the North Atlantic (e.g. Keppie et al. 1991). The main events in the evolution of Avalonia (e.g. Nance et al. 2002; Keppie et al. 2003a) are: (1) the development of juvenile crust at c. 1.3– 1.0 Ga; (2) an early phase(s) of arc magmatism prior to c. 650 Ma; (3) the accretion of this arc to Gondwana in the interval c. 665– 650 Ma; (4) main phase arc magmatism from c. 640 Ma to c. 570 Ma; (5) the diachronous transition from an arc to a platform between c. 600 Ma and c. 540 Ma; (6) rifting of Avalonia from Gondwana at c. 540–515 Ma and development of an endemic Avalonian fauna; (7) the gradual reintroduction of Gondwanan fauna in the Late Cambrian– Early Ordovician; (8) mid-Ordovician separation of Avalonia from Gondwana and development of an endemic fauna; and (9) the accretion of Avalonia to Laurussia sometime in the interval c. 440–415 Ma (Fig. 5). The basement upon which the main Avalonian arc was built is nowhere unequivocally exposed. However, constraints on its character are provided by the Sm –Nd isotopic composition of crustally derived Neoproterozoic and Early Cambrian felsic igneous rocks that represent melt fractions extracted from the basement. These rocks show similar initial 1Nd values that range between –2.5 and þ5.0 (Thorogood 1990; Barr & Hegner 1992; Whalen et al. 1994; Kerr et al. 1995; Murphy et al. 1996a, 2000; Keppie et al. 1997; Samson et al. 2000) and yield depleted mantle model ages (TDM) of 0.8 – 1.1 Ga in Atlantic Canada and 1.0–1.3 Ga in southern Britain (Thorogood 1990; Murphy et al. 2000). Detailed arguments for the interpretation of the Avalonian model ages have been presented elsewhere (e.g. Nance & Murphy 1994, 1996; Murphy et al. 2000; Murphy & Nance 2002), but collectively they suggest that the Neoproterozoic – Early Cambrian felsic magmas were produced
THE PERI-GONDWANAN TERRANES
Fig. 5. Interpretive late Neoproterozoic-Palaeozoic tectonostratigraphic column for Avalonia (modified from Murphy et al. 1999a; Nance et al. 2002).
largely as a result of recycling c. 1.0– 1.3 Ga crust. The depleted mantle model ages are therefore thought to record a genuine tectonothermal event during which the bulk of Avalonian basement was itself extracted from the mantle; the primitive isotopic signature suggests that it formed in one or more largely juvenile oceanic island arcs or oceanic plateau (Murphy et al. 2000). These model ages are broadly coeval with the c. 1 Ga (‘Grenville’) age of several collisional orogenic belts thought to record the amalgamation of the supercontinent Rodinia at c. 1.0–1.1 Ga. Hence, the basement to the main Avalonian arc is thought to have developed within the Panthalassa-like ocean (Mirovoi Ocean; McMenamin & McMenamin 1990; Meert & Powell 2001) that would have surrounded Rodinia following its amalgamation. Fragmentary evidence for subduction exists within Avalonia from at least 730 Ma, and for the interval 730 –650 Ma this is termed the early arc phase. In Atlantic Canada, examples of this activity include the c. 734 Ma calc-alkaline Economy River Gneiss in mainland Nova Scotia (Doig et al. 1993),
351
c. 729 Ma portions of the Hawkes Hill Tuff in the central Avalon Peninsula, Newfoundland (O’Brien et al. 2001), the c. 681 Ma arc-related Stirling Belt in Cape Breton Island (Bevier et al. 1993), the 680– 630 Ma back-arc volcanic rocks in Central Cape Breton Island (Keppie & Dostal 1998), and the calc-alkaline, c. 683 Ma Tickle Point Formation and c. 673 Ma Furby’s Cove Intrusive Suite in southern Newfoundland (Swinden & Hunt 1991; O’Brien et al. 1996). The rift ophiolite volcanic rocks of the Burin Group in Newfoundland (Strong et al. 1978) may extend this early Avalonian magmatic activity to c. 763 Ma (Krogh et al. 1988). In Nova Scotia, geochemical and isotopic data from Neoproterozoic pelitic schists and quartzites (Gamble Brook Formation) that may date from this early arc phase are consistent with deposition in a rifted arc setting. Data for the pelites suggest their derivation from a moderately differentiated ocean island arc source composed of juvenile Avalonian basement crust, whereas those for the quartzites suggest that they were derived from ancient cratonic basement (Murphy 2002). An Amazonian source for the quartzite is suggested by detrital zircon ages (Fig. 6) that generally correspond to the age provinces of the Amazon craton (e.g. Sadowski & Bettencourt 1996; Tassinari & Macambira 1999; Santos et al. 2000) and cluster in the ranges c. 1 Ga, c. 1.1– 1.2 Ga, c. 1.45 –1.55 Ga, c. 1.6 Ga, c. 1.9 Ga, c. 2.1 Ga and c. 2.7–2.8 Ga (Keppie et al. 1998; Barr et al. 2003a). In Britain, evidence for early arc-related activity is represented by the c. 700 Ma calcalkaline Stanner– Hanter Complex of central Wales (Patchett et al. 1980) and the c. 677 Ma calc-alkaline Malverns Plutonic Complex of the British Midlands (Tucker & Pharoah 1991). Early arc activity may also be represented in the undated gneisses of the Rosslare Complex in southeastern Ireland and the Coedana Complex in Anglesey, North Wales (Gibbons & Hora´k 1996). Detrital zircons from paragneisses from both the Malverns Plutonic Complex and the Coedana Complex are dominated by Mesoproterozoic ages, with major age concentrations at c. 1150–1250 Ma and c. 1500– 1650 Ma (Strachan et al. 2007). A period of high-grade metamorphism is recorded in the Coedana Complex at c. 666 Ma (Strachan et al. 2007) and in the Malvern Plutonic Complex at c. 650 Ma (Strachan et al. 1996b). Amphibolite-facies metamorphism at c. 650 Ma is also recorded in coastal Maine (Stewart & Tucker 1998; Stewart et al. 2001) and high-grade metamorphism of pre-630 Ma age may be present in central Cape Breton Island (Keppie et al. 1998) and southern Newfoundland (O’Brien et al. 1996). However, the Avalonian (v. Ganderian) status of
352 R. D. NANCE ET AL. Fig. 6. Detrital zircon data for late Neoproterozoic–early Palaeozoic successions in Avalonia (Bevier et al. 1990; Karabinos & Gromet 1993; Keppie et al. 1998; Thompson & Bowring 2000; Barr et al. 2003a), Ganderia (van Staal et al. 1996; Barr et al. 2003a), Meguma (Krogh & Keppie 1990; Murphy et al. 2004c), Carolinia (inherited zircons: Mueller et al. 1996; Ingle et al. 2003), Florida (Mueller et al. 1994), Cadomia (Miller et al. 2001; Ferna´ndez-Sua´rez et al. 2002a; Samson et al. 2005), Iberia (Ferna´ndez-Sua´rez et al. 2000, 2002a; Gutie´rrez-Alonso et al. 2003) and Bohemia (Linnemann et al. 2004) compared with the cratonic age provinces of Eastern Laurentia (Cawood et al. 2001), Baltica (Gower et al. 1990; Roberts 2003), Amazonia (Sadowski & Bettencourt 1996; Tassinari & Macambira 1999; Santos et al. 2000) and West Africa (Rocci et al. 1991; Boher et al. 1992; Potrel et al. 1998; Hirdes & Davis 2002). C-O, Cambrian– Lower Ordovician; S, Lower Silurian; NS, Nova Scotia; NB, New Brunswick; NE, New England; NWI, NW Iberian domain; O-M, Ossa –Morena zone; UK, United Kingdom.
THE PERI-GONDWANAN TERRANES
these latter regions is debated. Some form of accretion is also implicated by the ophiolitic rocks of the c. 760 Ma Burin Group (Keppie et al. 1991). This metamorphism is interpreted to reflect the accretion of Avalonia to the Gondwanan continental margin prior to the beginning of the main phase of Avalonian magmatism at c. 640 Ma and coincides with a temporary cessation (c. 650 –640 Ma) in subductionrelated magmatism (e.g. Murphy et al. 2000). The main phase of Avalonian magmatism is recorded in voluminous late Neoproterozoic magmatic arc-related volcanic and cogenetic plutonic rocks with crystallization ages of 640–570 Ma (e.g. Hermes & Zartman 1985, 1992; Krogh et al. 1988; Barr et al. 1990, 1994; Bevier & Barr 1990; Tucker & Pharoah 1991; Bevier et al. 1993; Doig et al. 1993; Thompson et al. 1996; Murphy et al. 1997; O’Brien et al. 2001; Compston et al. 2002). Coeval sedimentary successions that are dominated by volcanogenic turbidites are locally associated with these arc-related magmatic rocks, and have been attributed to deposition in a variety of intra-arc, interarc and back-arc basins (e.g. Pe-Piper & Murphy 1989; Pe-Piper & Piper 1989; Murphy et al. 1990; Pauley 1990; Smith & Socci 1990; O’Brien et al. 1996). This magmatic activity and the generation of arc-related basins are interpreted to reflect oblique (sinistral) subduction beneath the northern margin of West Gondwana (e.g. Murphy & Nance 1989). Detrital zircon data (Fig. 6) from late Neoproterozoic sedimentary successions (c. 0.6–0.7 Ga, c. 1.1–1.2 Ga, c. 1.3 Ga, c. 1.5 Ga, c. 1.7 Ga, c. 1.9–2.1 Ga, c. 2.2–2.3 Ga and c. 2.6– 2.7 Ga; Keppie et al. 1998; Barr et al. 2003a) contain important Mesoproterozoic populations and broadly match the age provinces of the Amazon craton (e.g. Sadowski & Bettencourt 1996; Tassinari & Macambira 1999; Santos et al. 2000). Although the onset of this main phase of activity was broadly synchronous throughout much of Avalonia, its cessation was diachronous, terminating at c. 590 Ma in New England (Kaye & Zartman 1980; Hermes & Zartman 1985, 1992; Thompson et al. 1996; Thompson & Bowring 2000), 600 Ma in southern New Brunswick (Bevier & Barr 1990; Barr et al. 1994; Currie & McNicoll 1999), 605 Ma in mainland Nova Scotia (Doig et al. 1991; Murphy et al. 1997; Keppie et al. 1998), 575 Ma in southern Cape Breton (Barr et al. 1990; Bevier et al. 1993), 585 Ma in Newfoundland (Krogh et al. 1988; O’Brien et al. 1996), and 600 Ma in the British Isles (e.g. Tucker & Pharaoh 1991; Hora´k 1993; Noble et al. 1993). Cessation of main-phase subduction was accompanied by a transition to intracontinental extension, marked by the onset of bimodal magmatism. This onset was similarly diachronous, occurring at c. 595 Ma in New England (Mancusco
353
et al. 1996), at c. 560 Ma in southern New Brunswick (Bevier & Barr 1990; Barr et al. 1994; Currie & McNicoll 1999), at c. 605 Ma in mainland Nova Scotia (Murphy et al. 1997), between 575 and 560 Ma in southern Cape Breton Island (Bevier et al. 1993), at c. 570 Ma in Newfoundland (O’Brien et al. 1996), and in the interval 570–560 Ma in Britain (Tucker & Pharoah 1991; Compston et al. 2002). Although the transition was locally accompanied by deformation and metamorphism, no evidence exists for the regional orogenesis, crustal shortening, and crustal thickening and uplift characteristic of continental collision zones. Instead, deformation is usually localized and resulted only in the inversion of some of the earlier volcanic arc basin successions. To account for such a tectonic transition in the apparent absence of a major collisional event, Murphy & Nance (1989) proposed that Avalonian subduction was terminated as a result of transform activity (Fig. 7). In their model, the main phase of
Fig. 7. General model for the late Neoproterozoic evolution of Avalonia modified after Murphy & Nance (1989). (a) Oblique subduction during the interval c. 635–590 Ma produces the main arc phase of Avalonian magmatism and opens a variety of volcanic arc basins in response to sinistral motion on basin-bounding faults. (b) Ridge–trench collision results in the structural inversion of the volcanic arc basins, the opening of new wrench basins, and the diachronous termination of subduction at c. 590 –540 Ma, in response to the progressive development of a dextral continental transform fault (after Murphy et al. 1999a; Nance et al. 2002). , Cambrian; Pre , Precambrian; A, away; T, towards.
354
R. D. NANCE ET AL.
Avalonian magmatism at c. 640 –590 Ma occurred as the result of oblique subduction, leading to the development of an extensional magmatic arc and a variety of volcanic arc basins. Subsequently, the interaction of a continental margin transform system with the subduction zone resulted in the termination of subduction, the structural inversion of some volcanic arc basins, and the formation of new rift- and wrench-related basins in the interval c. 590 –540 Ma. Murphy et al. (1999a), Keppie et al. (2000, 2003a) and Nance et al. (2002) proposed ridge – trench collision as a mechanism for the transition, to account for the diachronous cessation of arc volcanism and the apparent reversal of kinematics on major basin-bounding faults (e.g. Nance & Murphy 1990). Palaeomagnetic data from the c. 580 –570 Ma volcanic and interbedded clastic rocks of the Marystown Group in southern Newfoundland indicate that the strata were deposited at a palaeolatitude of 348 þ88/278 (McNamara et al. 2001), whereas data from the c. 595 Ma Lynn –Mattapan volcanic complex in southeastern New England yield a palaeolatitude of c. 408 (Thompson et al. 2006). However, whether these data imply proximity to West Africa (McNamara et al. 2001; Mac Niocaill et al. 2002), North Africa (Thompson et al. 2006), or Amazonia (Murphy et al. 2002b) is unclear. Cambrian platformal sequences in Avalonia contain shallow-water and intertidal limestones and siliciclastic rocks, as well as bimodal volcanic rocks (e.g. Murphy et al. 1985). The Avalonian fauna these rocks contain are indicative of coolwater, high-latitude conditions and thereby provide first-order constraints on palaeogeographical reconstructions (Landing 2004, 2005). The distinct provinciality of this fauna suggests that Avalonia was separated from Gondwana in the Early Cambrian (Theokritoff 1979; Landing 1996, 2005), which is consistent with the rapid subsidence recorded by its sedimentary rocks (e.g. Prigmore et al. 1997). Because this separation is not detectable palaeomagnetically (e.g. Van der Voo 1988), the seaway between Avalonia and Gondwana was probably narrow. That Avalonia remained broadly peri-Gondwanan in the Cambrian is also indicated by the breakdown of faunal barriers between it and Gondwana in the Middle Cambrian (Landing 2005) and by its Early Ordovician Gondwanan fauna. The gradual replacement of this fauna by endemic forms in the Arenig – Llanvirn, and then by fauna of Baltic and Laurentian affinities in the Llandeilo–Ashgill, suggests that large separation between Avalonia and Gondwana began in the Early Ordovician (Fortey & Cocks 2003). The accretion of Avalonia to Laurussia is of debated timing and is likely have been diachronous. On the basis of a Laurentian neodymium isotopic
signature in Early Silurian clastic sedimentary rocks in mainland Nova Scotia, Murphy et al. (1996b) have argued that the accretion of West Avalonia to Laurentia had occurred by the Early Silurian. Closure of the Iapetus suture in Britain, however, is not thought to have occurred until c. 420 Ma (e.g. Soper & Woodcock 1990), consistent with the available palaeomagnetic data, which suggest that any palaeolatitudinal separation between Avalonia and Laurussia had disappeared by the mid-Silurian (e.g. Miller & Kent 1988; Trench & Torsvik 1992; Potts et al. 1993; Hodych & Buchan 1998). However, following the cessation of subduction in the late Neoproterozoic, Avalonia remained unaffected by orogenesis until the Late Silurian –Early Devonian (e.g. Dallmeyer & Nance 1994; Waldron et al. 1996). Van Staal et al. (1998) consequently held the accretion of Avalonia to be responsible for the Late Silurian – Early Devonian Acadian orogeny.
Ganderia Inboard of Avalonia in the northern Appalachians are several terranes that have similar histories to that of Avalonia prior to c. 570 Ma, but record a distinct tectonic evolution during the latest Neoproterozoic and early Palaeozoic. They are thought either to represent portions of the leading edge of Avalonia following its separation from Gondwana (e.g. Keppie et al. 2003a) or to be parts of a separate periGondwanan basement block termed Ganderia (e.g. van Staal et al. 1998; Barr et al. 2002). The terranes include not only the early Palaeozoic siliciclastic passive margin succession of the Gander Zone of Newfoundland, New Brunswick and northern New England, but also the latest Neoproterozoic and younger rocks interpreted to represent the basement to the Gander Zone (e.g. van Staal et al. 1996; Van der Velden et al. 2004) in the Hermitage Flexure and Exploits subzone of Newfoundland, the Bras d’Or terrane of Cape Breton Island, the c. 550 Ma Upsalquitch gabbro of northwestern New Brunswick, the Brookville and New River terranes of southern New Brunswick, and the Seven Hundred Acre Island sequence in coastal Maine (Fig. 2). The Gander Zone has also been extended to the British Isles (Fig. 4) to include Cambrian–Early Ordovician siliciclastic successions in southeastern Ireland, Anglesey, the Isle of Man and the English Lake District (e.g. van Staal et al. 1996). In Newfoundland, the boundary between Ganderia and Avalonia is the Dover Fault, a vertical transcurrent structure that offsets the Moho (Keen et al. 1986). The inboard margin of Ganderia coincides with the Red Indian Line, which separates periGondwanan and peri-Laurentian elements in the northern Appalachians (van Staal et al. 1998).
THE PERI-GONDWANAN TERRANES
The main events in the evolution of Ganderia (e.g. Hibbard et al. 2007) are: (1) the development on unexposed basement of a carbonate –siliciclastic platform prior to c. 750 Ma; (2) an early phase of arc magmatism from c. 625 Ma to c. 605 Ma; (3) a younger arc magmatic phase between c. 570 and c. 525 Ma that accompanied metamorphism and deformation; (4) the development, following Cambrian cessation of subduction and possibly the separation of Ganderia from Gondwana, of a Middle Cambrian –Early Ordovician arc and a clastic back-arc passive margin; and (5) the accretion of Ganderia to Laurentia in the Late Ordovician–Silurian (Fig. 8). As is the case in Avalonia, the basement upon which the Neoproterozoic Ganderian arc was built is nowhere exposed, the oldest known rocks being stromatolite-bearing platform carbonates and quartzites in southern New Brunswick (Green Head Group) and coastal Maine (Seven Hundred Acre Island Formation), the depositional ages of which are very loosely constrained to the interval 670– 1230 Ma by 40Ar/39Ar metamorphic (Stewart et al. 2001) and U –Pb detrital zircon (Samson et al. 2000; Barr et al. 2003a) ages. However,
Fig. 8. Interpretive late Neoproterozoic –Palaeozoic tectonostratigraphic column for Ganderia.
355
initial 1Nd values of 24.8 to þ2.8 and TDM model ages of 0.85–1.7 Ga from felsic volcanic and plutonic rocks (e.g. Barr & Hegner 1992; Whalen et al. 1994, 1996; Kerr et al. 1995; Barr et al. 1998, 2003b; Samson et al. 2000) suggest that Ganderia was built upon basement that isotopically overlaps but is somewhat more evolved than that of Avalonia. Detrital zircon ages (Fig. 6) from metasedimentary rocks of the Bras d’Or terrane (George River Group: c. 1.1–1.2 Ga, c. 1.8 Ga, c. 1.9–2.0 Ga and c. 2.6 Ga; paragneiss: c. 0.5 Ga, c. 0.6–0.7 Ga, c. 1.0–1.2 Ga, c. 1.3 Ga, c. 1.5– 1.6 Ga; Keppie et al. 1998), Brookville terrane (Brookville paragneiss: c. 0.6 Ga, 0.9–1.1 Ga, 1.2–1.5 Ga, 1.6 Ga, 1.9 Ga and 2.6–2.7 Ga; Bevier et al. 1990; Green Head Group: c. 1.2–1.3 Ga, 1.5 –1.6 Ga, 1.7– 1.8 Ga, 1.9–2.0 Ga and 2.7 Ga; Barr et al. 2003a) and Gander Zone (c. 0.55 –0.8 Ga, 1.0–1.55 Ga and 2.5–2.7 Ga; van Staal et al. 1996) indicate an important Mesoproterozoic source and broadly match the age provinces of the Amazon craton (e.g. Sadowski & Bettencourt 1996; Tassinari & Macambira 1999; Santos et al. 2000). Evidence of an early phase of Neoproterozoic arc activity in Ganderia is restricted to the Brookville and New River terranes of southern New Brunswick, where it takes the form of the c. 625 Ma Lingley Suite (Currie & McNicoll 1999), the c. 622 Ma Blacks Harbour Granite (Barr et al. 2003b), and the c. 605 Ma protolith age for the Brookville orthogneiss (Bevier et al. 1990). In the British Isles, this early phase would also include the c. 612 Ma Coedana granite (Tucker & Pharoah 1991) in Anglesey, according to the correlation of van Staal et al (1998). However, based on detrital zircon data from associated paragneisses, Strachan et al. (2007) questioned the separation of Anglesey from neighbouring Avalonia. The younger phase of arc magmatism is more widespread and includes the c. 578–564 Ma Roti intrusive suite in southern Newfoundland (Dunning & O’Brien 1989; O’Brien et al. 1991, 1993), the c. 565 Ma Crippleback Intrusive Suite and associated volcanic rocks in central Newfoundland (Evans et al. 1990; Rogers et al. 2006), the c. 565–555 Ma plutons of the Bras d’Or terrane in Cape Breton Island (Barr et al. 1990; Dunning et al. 1990; Dostal et al. 1996) and the c. 553– 527 Ma Golden Grove Plutonic Suite (Dallmeyer & Nance 1992; White et al. 2002) and c. 539 Ma Simpsons Island Formation (Barr et al. 2003b) in souhern New Brunswick. Emplacement of these younger plutons commonly accompanied metamorphism and deformation that is bracketed between 571 and 564 Ma in southern Newfoundland (O’Brien et al. 1991) and has been dated at c. 550–540 Ma in Cape Breton Island (Dunning et al. 1990) and at c. 564 Ma (Bevier et al. 1990)
356
R. D. NANCE ET AL.
and c. 550 –540 Ma (Nance & Dallmeyer 1994) in southern New Brunswick. From the Middle Cambrian to Early Ordovician, Ganderia experienced renewed arc magmatism concomitant with the deposition of a clastic passive margin succession that is interpreted to record the development of an arc –back-arc system along its leading edge (Penobscot arc– Exploits–Tetagouche back arc; van Staal et al. 1998). Examples of arc and back-arc magmatism include the c. 513 Ma Lake Ambrose volcanic belt (Dunning et al. 1991) and c. 494 Ma Pipestone Pond Complex (Dunning & Krogh 1985) in Newfoundland, the c. 505 Ma North Boisdale Hills volcanic rocks in Cape Breton (White et al. 1994), the c. 515 Ma Mosquito Lake Road Formation (Johnson 2001) and c. 493– 497 Ma Annidale volcanic series in New Brunswick (McLeod et al. 1992), and the c. 509 Ma Ellsworth Formation in Maine (Stewart et al. 1995). The coeval siliciclastic passive margin succession (Gander margin; van Staal 1994) is represented by monotonous sequences of Middle Cambrian to Tremadoc sandstones, siltstones and shales in the Gander Group in Newfoundland, the Miramichi, Woodstock and Cookson groups in New Brunswick, and the Penobscot Formation in Maine. Potential correlatives of the Gander Group in the British Isles include the Bray, Cullenstown, Cahore and Ribband groups in SE Ireland, the Monian Supergroup in Anglesey, North Wales, the Manx Group in the Isle of Man and the Skiddaw Group in the English Lake District (e.g. van Staal et al. 1996, 1998). The accretion of Ganderia to Laurentia is thought to be recorded in widespread deformation and metamorphism of Late Ordovician–Silurian age that marks the accretion of the Penobscot arc and telescoping of the Exploits– Tetagouche back arc (e.g. Brunswick subduction complex; van Staal 1994), and in voluminous Early Silurian magmatism taken to reflect subduction of Ganderia beneath Laurentia (van Staal et al. 1998) and subsequent slab break-off (Whalen et al. 2006).
Meguma Outboard of Avalonia in the northern Appalachians lies the Meguma terrane (Fig. 2), which is exposed only in Nova Scotia but extends oceanward to the edge of the continental shelf from the Grand Banks to Cape Cod (e.g. Pe-Piper & Jansa 1999). It is separated from Avalonia by the Minas fault zone, a major late Palaeozoic strike-slip boundary (e.g. Webster et al. 1998). The main events in the evolution of the Meguma terrane (e.g. Murphy & Keppie 2005) are: (1) the development of a (?) Neoproterozoic –Early Ordovician siliciclastic passive margin; (2) the
development of a Late Ordovician– Early Devonian marine platform; (3) Devonian deformation, metamorphism and voluminous syn- to post-tectonic granitoid plutonism; (4) rapid uplift between c. 370 and c. 360 Ma; and (5) latest Devonian – Early Carboniferous basin formation. The basement upon which the sedimentary succession of the Meguma terrane was deposited is not exposed and the oldest rocks are those of the Meguma Group, a thick (.10 km), (?) Neoproterozoic –Early Ordovician succession (White et al. 2006) of turbiditic sandstones (Goldenville Formation) and overlying shales (Halifax Formation) that are widely attributed to deposition in a continental slope environment along the West African margin of Gondwana (e.g. Waldron 1992; Schenk 1997). Such a provenance is supported by the peri-Gondwanan (Acado-Baltic) affinity of its sparse trilobite fauna (Pratt & Waldron 1991; White et al. 2006) and by detrital zircon data (c. 0.6 Ga, 2.0–2.1 and 3.0 Ga; Krogh & Keppie 1990) that match the age provinces of the West African craton (e.g. Rocci et al. 1991; Boher et al. 1992; Potrel et al. 1998; Hirdes & Davis 2002) (Fig. 6). However, initial 1Nd values of þ0.2 to þ6.8 and TDM model ages of 0.8–1.3 Ga from crustally derived Early Silurian volcanic rocks of the White Rock Formation (Keppie et al. 1997) suggest that Meguma was floored in the Early Silurian by basement that is isotopically indistinguishable from that of Avalonia. Because there is no record of any deformation in Meguma prior to the Early Silurian, this basement is likely to have been present from the onset of deposition of the Goldenville Formation; that is, since the late Neoproterozoic. The Meguma Group is disconformably overlain by sandstones and shales of the Early Ordovician– Early Devonian Annapolis Group, which records the development of a shallow-marine siliciclastic shelf (e.g. Schenk 1995). The White Rock Formation at the base of the succession includes c. 440 Ma bimodal, rift-related volcanic and clastic rocks (Keppie & Krogh 2000) and contains a detrital zircon population that differs from that of the Meguma Group, but resembles those of coeval strata in Avalonia (Fig. 6). As these zircons do not occur in the underlying strata, Murphy et al. (2004b) argued that Meguma could not have been an isolated terrane at that time. The data include an important Mesoproterozoic (1.0– 1.4 Ga) population and have been taken to imply proximity of Meguma to Avalonia by the Early Silurian (Murphy et al. 2004b). Deformation and metamorphism of the Meguma terrane continued intermittently from c. 395– 388 Ma (Hicks et al. 1999) until c. 320 Ma (Culshaw & Reynolds 1997), and may record its
THE PERI-GONDWANAN TERRANES
357
accretion to Laurentia in the Middle Devonian– Early Carboniferous; regional fold trends and the kinematics of the Minas fault zone are consistent with a dextral transpressive regime (e.g. Mawer & White 1987). This is age equivalent to the Neo-Acadian orogeny of New England (Robinson et al. 1998). However, on the basis of detrital zircon data that suggest their proximity, Murphy et al. (2004b) argued that Meguma arrived with Avalonia in the Early Silurian. Regional deformation was followed by the intrusion of the peraluminous South Mountain Batholith and related plutons at c. 372 Ma (Clarke et al. 1997). Coeval with continued dextral movement on the Minas fault zone (Dallmeyer & Keppie 1987), this shortlived episode of voluminous magmatism was followed by rapid uplift at c. 370 –360 Ma (Keppie & Dallmeyer 1995) and, in the Late Devonian– Early Carboniferous, by the development of the St. Marys (Murphy 2000, 2003) and other wrench-related basins.
Carolinia Carolinia (Carolina Zone of Hibbard et al. 2002) comprises several peri-Gondwanan terranes located along the eastern margin of the southern Appalachians (Fig. 2) that were brought together in the late Neoproterozoic by the collision of the Carolina and Charlotte arcs (e.g. Hibbard et al. 2007). To the west, Carolinia is juxtaposed against a variety of Piedmont terranes and, to the east, it surrounds on three sides and is in tectonic contact with the c. 1.0 Ga Goochland terrane, which is thought to be a displaced fragment of Laurentia (e.g. Farrar 1984; Bartholomew & Tollo 2004). The main events in the evolution of Carolinia (e.g. Hibbard et al. 2002, 2007) are: (1) an early phase of arc magmatism at c. 670 Ma; (2) a phase of juvenile arc magmatism like that of Avalonia between c. 633 Ma and c. 605 Ma; (3) a phase of mature arc magmatism like that of Ganderia, which accompanied metamorphism and deformation between c. 580 and c. 540 Ma; (4) the development of a Middle Cambrian passive margin; and (5) the accretion of Carolinia to Laurentia in the Late Ordovician–Silurian (Fig. 9). The oldest rocks in Carolinia are c. 672 Ma granitoid bodies of the Roanoke Rapids terrane, which are interpreted as evidence of early arc magmatism broadly coeval with that in Avalonia (Hibbard et al. 2002). The basement of Carolinia is not exposed. However, initial 1Nd values of þ0.1 to þ5.9 and TDM model ages of 0.7– 1.1 Ga from volcanic and plutonic rocks of the Virgilina sequence (Samson et al. 1995; Mueller et al. 1996; Ingle et al. 2003) suggest that this part of Carolinia, like Avalonia, was formed from juvenile
Fig. 9. Interpretive late Neoproterozoic– Palaeozoic tectonostratigraphic column for Carolinia.
peri-Rodinian (Mirovoi) crust and was located outboard from the northern margin of West Gondwana until at least 700 Ma (Murphy et al. 2004a). The Virgilina sequence is dominated by c. 633– 605 Ma subaqueous felsic– intermediate pyroclastic rocks and accompanying granitoids (Wortman et al. 2000; Ingle et al. 2003) but includes turbidites and metabasalts, and records the development of a juvenile arc similar to that recorded in the main phase of arc magmatism in Avalonia. Metavolcanic and plutonic rocks of similar age (c. 626–619 Ma) and 1Nd values (þ2 to þ3.5) to those of the Carolina arc have also been reported from the subsurface of the Atlantic Coastal Plain in South Carolina (Dennis et al. 2004). Unlike Avalonia, however, this juvenile arc assemblage is overlain unconformably by the Uwharrie –Albemarle mature arc sequence that includes volcanic and plutonic rocks with ages in the range 586– 539 Ma (Wright & Seiders 1980; Ingle et al. 2003). The bulk of these rocks show initial 1Nd values of 20.7 to þ4.1 and TDM model ages of 0.9–1.5 Ga (Mueller et al. 1996; Ingle et al. 2003). Possible episodes of arc rifting have been documented during this younger mature arc
358
R. D. NANCE ET AL.
phase (e.g. c. 538 –535 Ma Means Crossroads complex; Dennis & Shervais 1991, 1996), the magmatism of which was accompanied by metamorphism and deformation (e.g. Shervais et al. 1996). This tectonothermal activity, which occurred prior to c. 547 Ma in North Carolina (Hibbard & Samson 1995) and has been constrained to the intervals 570– 535 Ma (Dennis & Wright 1997) and 557– 548 Ma (Barker et al. 1998) in South Carolina, has been attributed to the collision of the Uwharrie–Albemarle (Carolina) arc with the broadly contemporaneous Charlotte arc (e.g. Shervais et al. 2003). Arc magmatism in Carolinia continued into at least the Early Cambrian before terminating, and the arc-related rocks are succeeded unconformably by Middle Cambrian quartz-rich platformal strata (Asbill Pond Formation), which contain mixed Tethyan –Avalonian, cool-water trilobites that share faunal affinities with Cadomia, Gondwana and Avalonia (Theokritoff 1979; Secor et al. 1983; Samson et al. 1990). However, the presence of 1.0–1.6 Ga inherited zircons in the Uwharrie – Albemarle sequence (Mueller et al. 1996; Ingle et al. 2003) and 1.1–1.8 Ga detrital zircons in latest Neoproterozoic –Cambrian sandstones and conglomerates (Samson et al. 1999, 2001) suggest a provenance in Amazonia and/or Baltica (Fig. 6). As in Ganderia, the accretion of Carolinia to Laurentia is thought to be recorded in widespread deformation and low-grade metamorphism of Late Ordovician–Silurian age. 40Ar/39Ar muscovite and biotite ages from the Albemarle Group (Carolina arc) suggest peak metamorphism and cleavage formation at c. 455 –443 Ma (Offield et al. 1995), whereas amphibolite-facies rocks in the Charlotte arc record amphibole ages of c. 430 –425 Ma (Sutter et al. 1983). Subduction of Carolinia beneath Laurentia is interpreted to have been oblique with a sinistral component (Hibbard 2000). On the basis of its Cambrian tectonomagmatic record, and the age and kinematics of its accretion, Hibbard & van Staal (2005) and Hibbard et al. (2007) have suggested that Carolinia was more closely affiliated with Ganderia than with Avalonia, at least after c. 570 Ma. However, the Virgilina sequence records crystallization ages and isotopic signatures (Samson et al. 1995; Wortman et al. 2000) that closely match those of Avalonia (Nance & Murphy 1996), suggesting that the juvenile arc, which evolved prior to c. 570 Ma, was built upon Avalonian basement.
Florida Peri-Gondwanan rocks occur in the Florida subsurface (Suwannee terrane) and consist of c. 550 Ma arc-related volcanic rocks with initital 1Nd values
(þ1.1 to 24.1) and TDM model ages (1.0–1.6 Ga) that resemble those of Ganderia, and c. 625 and c. 550 Ma granitoid rocks with 1.1– 1.2 Ga inherited zircons that suggest a crustal basement of ‘Grenville’ (c. 1 Ga) age (Heatherington et al. 1996). The Suwannee terrane has been tectonically linked to the Bassaride –Rokelide orogen of West Africa (e.g. Dallmeyer 1989). However, the common lead isotopic signature of the igneous rocks of the Suwannee terrane resembles that of Carolinia and plots in the same field as Amazonian basement rocks, suggesting a common Amazonian origin (Heatherington & Mueller 2005). Likewise, Sm– Nd isotope data from Mesozoic tholeiitic basalts in north Florida suggest their derivation from a c. 1 Ga source (Heatherington & Mueller 1999), which is similar to the age suggested for the source of Mesozoic tholeiites in Carolinia (Pegram 1990). The Neoproterozoic rocks of the Suwannee terrane are overlain unconformably by undeformed Ordovician– Devonian strata with high-latitude Gondwanan fauna (Keppie et al. 2003a). Detrital zircons from a sandstone in these strata (Fig. 6) yielded minimum ages of 515–637 Ma and 1.7 – 1.8, 2.0–2.3, 2.4 and 2.7 Ga (Mueller et al. 1994).
Cadomia Cadomia in the Armorican Massif of northwestern France (Fig. 4) is the type area of the late Neoproterozoic ‘Cadomian’ orogeny, evidence for which is also present in portions of Iberia and Bohemia. The Armorican Massif is divided into three contrasting crustal blocks by the Variscan North and South Armorican shear zones, both of which may have been localized on pre-existing Cadomian lineaments (e.g. Watts & Williams 1979). In all three blocks, Neoproterozoic volcanic– sedimentary successions are overlain unconformably by marine sedimentary sequences of Cambrian age. Unlike in central and southern Armorica (South Armorican Massif), Variscan orogenic activity in north Armorica was weak, with the result that here a protracted late Neoproterozoic tectonothermal evolution is largely preserved intact. The North and South Armorican Massifs have been correlated around the oroclinal curvature of the late Palaeozoic Iberian –Armorican arc (Fig. 4) with the NW Iberian domain and Ossa–Morena zone of Iberia (e.g. Quesada 1990, 1991). In the North Armorican Massif (Fig. 10), the Icartian gneiss exposed in northwestern segments of the Cadomian belt represents one of the few examples of undisputed cratonic basement exposed in any of the peri-Gondwanan arc terranes. The age (c. 2.1 Ga) and isotopic signature (initial 1Nd ¼ þ0.2 to þ1.5, TDM ¼ 2.2 –2.4 Ga) of this
THE PERI-GONDWANAN TERRANES
Fig. 10. Interpretive late Neoproterozoic– Palaeozoic tectonostratigraphic column for Cadomia.
basement (Samson & D’Lemos 1998; Inglis et al. 2004) resembles that of the 2.1 Ga Eburnean basement of the West African craton (Alle`gre & Ben Othman 1980; Boher et al. 1992). Similarly, initial 1Nd values of þ1.6 to 29.9 and TDM model ages of 1.0–1.9 Ga from late Cadomian granitoids (D’Lemos & Brown 1993) suggest mixing of mantle-derived material at c. 600 Ma with Icartianlike basement derived from the mantle at c. 2.1 Ga. Further to the SE, early arc-related magmatism is recorded by c. 755 –745 Ma orthogneisses within the Penthie`vre Complex (Egal et al. 1996; Nagy et al. 2002), a c. 625 Ma intrusive granodiorite (Samson et al. 2003), and detrital zircon ages of c. 670 –625 Ma obtained from unconformably overlying conglomerates (Guerrot & Peucat 1990; Samson et al. 2003). Continued calc-alkaline magmatism accompanied deformation and metamorphism of the early Cadomian arc at c. 620– 610 Ma (Miller et al. 1999; Samson & D’Lemos 1999; D’Lemos et al. 2001; Inglis et al. 2005).
359
Exhumation and erosion of the early Cadomian arc was followed by crustal extension and accumulation of Brioverian volcanic– sedimentary successions in the interval c. 610–580 Ma (e.g. Dennis & Dabard 1988; Chantraine et al. 1994; Egal et al. 1996; Strachan et al. 1996a; Miller et al. 1999; Cocherie et al. 2001). A second major period of intermediate to granitic magmatism occurred at c. 580– 570 Ma (Brown et al. 1990; Nagy et al. 2002; Inglis et al. 2005) and partly overlapped regional deformation of parts of the Brioverian (Dallmeyer et al. 1991). Regionally significant sinistrally transpressive deformation at c. 540 Ma (e.g. Strachan et al. 1989) was accompanied by intracrustal melting, migmatization and granitoid emplacement at c. 550– 540 Ma (e.g. Gapais & Bale´ 1990; D’Lemos et al. 1992; Egal et al. 1996). The basement isotopic signatures of Cadomia, together with U –Pb detrital zircon data from Brioverian sedimentary succession(s) that cluster in the intervals c. 600–650 Ma, c. 1.9–2.2 Ga, c. 2.3– 2.5 Ga, c. 2.6–2.7 Ga and c. 3.1 –3.2 Ga (Miller et al. 2001; Ferna´ndez-Sua´rez et al. 2002a; Samson et al. 2005) (Fig. 6), suggest a palaeogeographical position near the West African craton, a linkage supported by the recent identification of tectonothermal events at c. 755–700 Ma (D’Lemos et al. 2006) and c. 640– 650 Ma (Inglis et al. 2005) in the Anti-Atlas Mountains of Morocco. Hence, Cadomia, in contrast to Avalonia and Carolinia, appears to have originated above Palaeoproterozoic crust along the continental margin of West Africa, rather than within the peri-Rodinian Mirovoi Ocean. These data imply that Avalonia and Cadomia (i.e. the ‘Avalonian–Cadomian belt’ of Murphy & Nance 1989) did not form a coherent continental margin orogenic belt until the accretion of Avalonia to northern West Gondwana at c. 665– 650 Ma. The Cambrian strata of Cadomia comprise thick platformal assemblages with a stratigraphy and fauna that are virtually indistinguishable from those of Morocco (Landing 1996). The lower part of the succession consists of basal conglomerates that are overlain by c. 370 m of Lower Cambrian thin- to thick-bedded dolostone and c. 200 m of trilobite-bearing limestone with a fauna typical of tropical, low-latitude waters. Landing (2005) stressed that the presence of evaporites, oncoids, ooids and thrombolite build-ups are features that do not occur in Avalonian successions.
Iberia Neoproterozoic rocks in the Iberian Massif occur in two distinct tectonostratigraphic belts (Quesada 1990, 1991), the NW Iberian domain and the Ossa –Morena zone (Fig. 4). The NW Iberian
360
R. D. NANCE ET AL.
domain, which includes the Central Iberian, West Asturian– Leonese and Cantabrian zones, is dominated by thick c. 600 –540 Ma pelite –greywacke sequences and interbedded calc-alkaline volcanic rocks of Ediacaran age (c. 559 Ma, Gutie´rrezAlonso et al. 2003; Rodrı´guez-Alonso et al. 2004), overlain, in some places uncomformably, by late Vendian –Early Cambrian synorogenic clastic and shallow-marine coastal sedimentary rocks (Fig. 11). Occasional intrusive, c. 600 Ma, calc-alkaline, volcanic arc-related rocks are also present (Ferna´ndez-Sua´rez et al. 1998). Although no penetrative fabric of late Neoproterozoic age has been described, the angular unconformity separating the Neoproterozoic from the Palaeozoic succession has been attributed to deformation of this age (Dı´az Garcı´a 2006). The Ossa –Morena zone to the south exposes abundant c. 600 – 575 Ma igneous rocks that are interpreted to represent a volcanic arc (Quesada 1990, 1991). These rocks were repeatedly deformed and metamorphosed during the latest Neoproterozoic to Early
Fig. 11. Interpretive late Neoproterozoic –Palaeozoic tectonostratigraphic column for Iberia. BCSZ, Badajos–Co´rdoba shear zone.
Cambrian Cadomian orogeny, which was associated with the development of an Andean-type continental margin at c. 550– 530 Ma. The suture between these two tectonostratigraphic belts is defined by the Badajoz –Co´rdoba shear zone and adjacent accretionary units, which include ophiolitic and eclogitic rocks (e.g. Quesada & Dallmeyer 1994; Eguiluz et al. 2000; Bandres et al. 2002). When the oroclinal curvature of the late Palaeozoic Iberian – Armorican arc (Fig. 4) is removed (Weil et al. 2001), the NW Iberian domain and Ossa–Morena zone can be correlated with similar units in the South Armorican and North Armorican Massifs, respectively (e.g. Quesada 1990, 1991). Detrital zircon data from Neoproterozoic metasedimentary rocks of the Ossa–Morena zone (Fig. 6) reveal zircon populations that are typical of the West African craton (Ferna´ndez-Suarez et al. 2002a; Gutie´rrez-Alonso et al. 2003) and correlate directly with those found in the North Armorican Massif (Samson et al. 2005). The nature and affinity of the basement underlying the NW Iberian domain, however, is less certain. The presence of c. 1.0 Ga detrital zircons in Neoproterozoic clastic rocks has been taken to suggest an Amazonian source for at least some of the detritus (Ferna´ndez-Sua´rez et al. 2000; Gutie´rrez-Alonso et al. 2003), although a potential source of c. 1.0 Ga zircons exists in the Central Sahara (deWit et al. 2005) beyond the Pan-African Trans-Saharan orogen. However, Sm– Nd isotope data for Neoproterozoic sedimentary rocks (Ugidos et al. 2003) yield 1Nd values between 22.2 and 20.4 that are also consistent with an Amazonian source and outside the range typical of time-equivalent sedimentary rocks with West African craton provenance. This is further supported by detrital muscovite data from Early Cambrian sedimentary strata that are likely to reflect more proximal source areas and cluster in the age ranges c. 550– 640 Ma, c. 920–1060 Ma and c. 1580– 1780 Ma (Gutie´rrez-Alonso et al. 2005). Evidence of a younger, Avalonia-like basement may also be found in ophiolitic amphibolites of the Cabo Ortegal Complex (Variscan suture) of NW Spain, which have yielded an age of c. 1160 Ma that is interpeted to date crystallization of the gabbroic protolith (Sa´nchez Martı´nez et al. 2006). On the other hand, the Cambrian to Silurian stratigraphy and fauna of the southern NW Iberian domain are very similar to time-equivalent successions in the Mauritanide foreland, suggesting a West African connection by the early Palaeozoic (Quesada et al. 1991; Quesada 1997). These similarities include the Arenig Armorican Quartzite and Late Ordovician glaciomarine diamictites, and the presence of shallow-marine Tethyan (African) trilobites. Sm –Nd data for the Cambrian
THE PERI-GONDWANAN TERRANES
sedimentary rocks (Ugidos et al. 2003) also yield more negative 1Nd values (27.0 to 23.8) further supporting a peri-West African source, whereas detrital zircons in the Ordovician sedimentary rocks yield mixed populations (Ferna´ndez-Sua´rez et al. 2002b; Martı´nez Catala´n et al. 2004) with both West African and Amazonian signatures. Neoproterozoic to Cambrian reworked bauxitic sedimentary rocks in the Sierra Albarrana unit of the Ossa –Morena zone that are similar to coeval rocks in the Mauritanide foreland likewise suggest a connection with West Africa (Quesada 1997). Taken together, these data suggest that, in contrast to the Neoproterozoic sources, the Palaeozoic detritus within both the NW Iberian domain and the Ossa –Morena zone have mixed Amazonian and West African components. To account for this observation, Ferna´ndez-Sua´rez et al. (2002b) and Gutie´rrez-Alonso et al. (2003, 2005) proposed that the NW Iberian domain lay off Amazonia – Oaxaquia in the Neoproterozoic, but was transferred along the Gondwanan margin to a position closer to West Africa and stitched to the Ossa – Morena zone around the time of the Precambrian –Cambrian boundary. A Cambrian rifting event is recorded in Iberia with the development of a sandstone–limestone platform at c. 520 –510 Ma. By this time, the NW Iberian domain and Ossa –Morena zone were proximal to each other and contain similar fauna, which is indistinguishable from that of Cadomia (e.g. Lin˜a´n & Ga´mez-Vintaned 1993; Vidal et al. 1994). It is therefore likely that all three regions were juxtaposed by the Cambrian. Voluminous bimodal rift-related magmatism in the Ossa – Morena zone ranges from Early Cambrian to Late Ordovician in age (Quesada 1990; Giese & Buehn 1994; Sa´nchez-Garcı´a et al. 2003) and is thought to reflect protracted continental rifting that culminated in the Early Ordovician with the opening of a significant tract of new ocean and the development of a breakup unconformity (Quesada 1990). In the NW Iberian domain, rift-related magmatism is also very abundant, with mainly Early Ordovician ages (Valverde-Vaquero & Dunning 2000). Widespread Arenig subsidence, recorded in the broad distribution of the Armorican Quartzite across Cadomia and Iberia (e.g. Noblet & Lefort 1990), reflects post-breakup thermal subsidence and the establishment of a new passive margin.
Bohemia The Bohemian Massif of Central Europe (Fig. 4) is divided into several zones. Of these, the Moldanubian and Saxo-Thuringian zones have been shown to have close affinities with Cadomia and West Africa (Fig. 12), whereas the Moravo-Silesian
361
Fig. 12. Interpretive late Neoproterozoic–Palaeozoic tectonostratigraphic column for Bohemia.
zone (Brunovistulian) has Avalonian affinities (Zulauf et al. 1999; Finger et al. 2000; Linnemann & Romer 2002; Linnemann et al. 2004). In the Moldanubian zone, the Tepla´ –Barrandian unit consists of late Neoproterozoic (c. 590–570 Ma) arc-related interbedded metasedimentary and metavolcanic rocks that are interpreted to record southerlydirected (present co-ordinates) subduction beneath northern West Gondwana followed by the accretion of an island arc to the Gondwanan margin (Zulauf et al. 1999; Do¨rr et al. 2002). These sequences are unconformably overlain by unmetamorphosed (?) Early Cambrian sedimentary rocks, which were deposited in a rift-related and/or transtensional regime (Zulauf et al. 1997; Drost et al. 2004), and by an Ordovician sequence of sedimentary and volcanic rocks that reflect rifting along the Gondwanan margin. U –Pb isotopic data suggest that the Neoproterozoic sequence has an age of c. 590– 570 Ma (Do¨rr et al. 2002; Drost et al. 2004). The
362
R. D. NANCE ET AL.
occurrence of isolated fragments of c. 2.1 Ga cratonic basement in the Moldanubian zone (Kro¨ner et al. 1988; Wendt et al. 1993) suggests that the Neoproterozoic sedimentary rocks were deposited on Icartian-like basement. In the Saxo-Thuringian zone, Neoproterozoic rocks include c. 570 –540 Ma interbedded marine sedimentary and arc-related submarine volcanic rocks, and granitoid plutons (Linnemann et al. 2000, 2004; Linnemann & Romer 2002). Inheritance with ages in the range c. 2.1–1.7 Ga is recorded in Pb –Pb analyses of zircon from igneous bodies and is interpreted to reflect West African (Eburnean) basement (Linnemann et al. 2000). Detrital zircon grains in Neoproterozoic greywacke– mudstone successions have yielded sensitive high-resolution ion microprobe (SHRIMP) U/Pb ages in the ranges c. 0.56– 0.75, 1.7–2.25, 2.45–2.75 and 3.0–3.4 Ga (Linnemann et al. 2004) (Fig. 6), further supporting a West African provenance for the Saxo-Thuringian zone. The underlying cratonic basement of these successions is not exposed, but TDM model ages fall in the range c. 1.3–1.9 Ga (Linnemann & Romer 2002), indicating a source area dominated by old cratonic crust. Similar TDM model ages also occur in the Cambian– Ordovician sedimentary rocks of this zone (Linnemann et al. 2004). The older part of the Neoproterozoic succession in the Saxo-Thuringian zone is interpreted to have formed in a back-arc basin at c. 560 –570 Ma (Linnemann et al. 2000). Volcanic –sedimentary sections that formed proximal to the former arc occur in the northern part of the zone (Buschmann 1995; Kemnitz 2007), whereas passive margin sequences of the back-arc basin occur in the southeastern part (Linnemann 1991, 1995). Younger Neoproterozoic to Early Cambrian sedimentary rocks are interpreted to have formed in a retroarc basin that originated at c. 545 –540 Ma as a result of arc –continent collision and the resulting closure of the older back-arc basin (Linnemann et al. 2007b). Late Neoproterozoic to Cambrian strata record a transition from a probably cooler to a warmer palaeoclimate, similar to that recorded in Cadomia and the Ossa – Morena zone of Spain (Buschmann et al. 2006). Cambrian faunas likewise resemble those of other Cadomian-type terranes ´ lvaro et al. 2003), but also show affinities with (A those of Baltica (Havlı´cek 1999) and Carolinia (Samson et al. 1990). Latest Neoproterozoic –earliest Cambrian deformation and basin inversion in Saxo-Thuringia was followed by intense granitoid plutonism at c. 540 Ma, and by c. 530 –500 Ma siliciclastic and carbonate deposition thought to represent sedimentation along a ‘San Andreas-style’ transform margin (Linnemann & Romer 2002). Nd isotopic data and
stratigraphic criteria (e.g. the presence of Late Ordovician glaciomarine diamictite of the Saharan glaciation) support a West African source (Linnemann & Romer 2002). In the Moravo-Silesian zone of the southeastern Bohemian Massif, the Brunovistulian unit consists of metasedimentary rocks that are thought to have been deposited in a back-arc basin and then metamorphosed to the amphibolite facies by arc – continent collision and basin inversion at c. 600 Ma (e.g. Finger et al. 2000). This sequence is intruded by c. 590–580 Ma and c. 550 Ma granitoid rocks. According to Finger et al. (2000), the tectonothermal evolution of the Brunovistulian unit strongly resembles that of Avalonia, an interpretation supported by Sm –Nd isotopic data (Hegner & Kro¨ner 2000) and c. 1 Ga U –Pb detrital zircons and igneous zircon cores (Friedl et al. 2000) that suggest an Amazonian connection in the Neoproterozoic. A Cambrian –Ordovician tectonothermal event is thought to reflect crustal thinning and rifting.
Middle American terranes The occurrence of early Palaeozoic Gondwanan fauna in several terranes in Middle America shows these terranes to be of peri-Gondwanan affinity. They include: (1) the Chortis block of Honduras and Guatemala; (2) Oaxaquia, which underlies much of Mexico (Ortega-Gutie´rrez et al. 1995); and (3) the Maya terrane of the Yucatan Peninsula (Fig. 3). Each of these terranes exposes c. 1 Ga basement, but only the Maya terrane contains rocks of late Neoproterozoic age (Keppie & Ortega-Gutie´rrez 1999). The c. 1 Ga basement of Middle America is isotopically transitional between that of the Grenville Belt of eastern North America and the basement massifs of c. 1 Ga age in the northern Andes of Colombia (Ruiz et al. 1999), and has been related to mixing of juvenile ‘Grenville’ (c. 1 Ga) and Archaean sources (Cameron et al. 2004). The Maya terrane is thought to have been contiguous with the Florida basement (Suwannee terrane) until the opening of the Gulf of Mexico in the Mesozoic (e.g. Pindell et al. 1990; Dickinson & Lawton 2001). In the Neoproterozoic –early Palaeozoic, these terranes are thought to have lain along the northern margin of Amazonia (Keppie & Ramos 1999; Keppie et al. 2003a, 2006) in accordance with the palaeomagnetic data of Ballard et al. (1989). Of the c. 1 Ga basement rocks, those of the Oaxacan Complex in southern Mexico (the largest exposed portion of Oaxaquia) are best known and consist of: (1) a metavolcanic –metasedimentary juvenile arc sequence of uncertain age; (2) a
THE PERI-GONDWANAN TERRANES
c. 1140 Ma, bimodal, within-plate intrusive suite that was deformed and metamorphosed at c. 1100 Ma; (3) a c. 1012 Ma anorthosite–gabbro that was deformed and metamorphosed in the granulite facies at c. 1104–980 Ma; and (4) c. 920 Ma post-tectonic calc-alkaline plutonism (Keppie et al. 2001, 2003b; Ortega-Obrego´n et al. 2003; Solari et al. 2003. This basement complex is unconformably overlain by latest Cambrian– Ordovician sedimentary rocks that contain trilobites of Gondwanan affinity (Robison & Pantoja-Alor 1968: taxonomy revised by Shergold 1975; Landing et al. 2006). Correlative rocks in northeastern Mexico (Novillo Gneiss; Cameron et al. 2004) are unconformably overlain by Silurian rocks containing brachiopods most similar to those in the Merida Andes of Venezuela (Boucot et al. 1997). These data, together with the distribution of Ordovician facies belts along the margin of Amazonia, suggest that Chortis –Oaxaquia– Maya may have been derived from a gap in the Ordovician facies belts that exists north of Colombia (e.g. Keppie 1977; Cocks & Fortey 1988; Keppie & Ortega-Gutie´rrez 1995; Boucot et al. 1997; Keppie et al. 2001). Detrital zircon ages in the Ordovician sedimentary sequences range from 980 to 1230 Ma, matching the age provinces of the Oaxacan Complex and the c. 1 Ga basement massifs in the northern Andes (Gillis et al. 2001). A variety of data suggest that an ocean lay between Chortis– Oaxaquia–Maya and Laurentia until the Permo-Carboniferous (Keppie & Ramos 1999), implying that all of these terranes were transferred to Laurentia during the amalgamation of Pangaea. The first appearance of fauna with Laurentian affinities in Chortis –Oaxaquia occurs in Mississippian rocks that unconformably overlie those of Early Palaeozoic age (Sour-Tovar et al. 1996; Stewart et al. 1999; Navarro-Santillan et al. 2002). The detrital zircon record, which indicates that Oaxaquia was isolated from the southern margin of Laurentia until the Carboniferous (Gillis et al. 2001), is consistent with palaeomagnetic data from c. 1 Ga rocks of the Oaxacan Complex that would locate Oaxaquia between Amazonia and Baltica (Ballard et al. 1989; Keppie & Ortega-Gutie´rrez 1999). The Maya terrane contains c. 1.23 Ga orthogneisses that underwent granulite-facies metamorphism at 990 – 975 Ma (Weber & Ko¨hler 1999; Ruiz et al. 1999), a history that suggests a genetic relationship to Oaxaquia (Keppie & Ramos 1999). Zircons from plutonic rocks found in boreholes in the Yucatan Peninsula have yielded late Neoproterozoic ages (c. 545 Ma; Krogh et al. 1993) and Late Silurian ages have been recorded in plutons in the Maya Mountains (Steiner & Walker 1996). Although
363
the relationship between these units is not exposed, restoration of the c. 608 anticlockwise rotation that occurred during the early Mesozoic opening of the Gulf of Mexico (Molina-Garza et al. 1992; Dickinson & Lawton 2001) supports the former continuity of the Maya terrane with northern Oaxaquia and Florida. Although the above data imply a position for Chortis –Oaxaquia –Maya along the periphery of the Amazonian craton throughout the Neoproterozoic, the paucity of Neoproterozoic subductionrelated magmatism in Oaxaquia and the Chortis block suggests that these terranes lay inboard of the late Neoproterozoic Avalonia and Carolina arcs. They may also have formed the basement of these arcs, the felsic magmas of which were produced largely as a result of recycling c. 1 Ga crust (Murphy et al. 2004a; Samson et al. 1995).
Other peri-Gondwanan terranes Other peri-Gondwanan terranes occur in southern and eastern Europe, the Middle East and Brazil. Those in the Middle East and Brazil resemble Avalonia in that they are juvenile and preserve remnants of primitive island arcs developed within the Panthalassa-like Mirovoi Ocean. These include the c. 900– 700 Ma island arc rocks of the Arabian– Nubian Shield (e.g. Stern 1994; Blasband et al. 2000) and the c. 950–900 Ma calc-alkaline granitoid orthogneisses and metarhyolites of the Tocantins province (Goia´s magmatic arc) in central Brazil. Rocks of the Tocantins province yield a similar envelope of 1Nd growth lines and almost identical (c. 0.9–1.2 Ga) depleted mantle model ages to those of Avalonia and Carolinia (Pimental & Fuck 1992; Nance & Murphy 1996). However, the position of this juvenile crust within the peri-Rodinian Mirovoi Ocean and its location relative to Avalonia and Carolinia at this time is unknown. Peri-Gondwanan remnants in southern and eastern Europe resemble Cadomia and occur as inliers in the Alps (e.g. Neubauer 2002; von Raumer et al. 2002), the South Carpathians (e.g. Kra¨utner 1993), the Serbo-Macedonian Massif of Bulgaria (Graf et al. 1998) and northern Greece (Himmerkus et al. 2007), and the Menderes Massif (e.g. Kro¨ner & Sengo¨r 1990; Canden et al. 2000) and western Pontides (e.g. Ustao¨mer 1999; Chen et al. 2002) in Turkey. They are interpreted to record accretion (c. 650 Ma), Andean-type subduction (c. 570 – 520 Ma), back-arc rifting (c. 520 – 500 Ma) and back-arc spreading (c. 485 Ma) along the northern margin of Gondwana (Neubauer 2002), but their positions relative to other peri-Gondwanan terranes are poorly constrained.
364
R. D. NANCE ET AL.
Continental reconstructions Neoproterozoic Available palaeogeographical constraints clearly separate the peri-Gondwanan terranes from eastern Laurentia during the late Neoproterozoic and earliest Palaeozoic, and link them, instead, with West Gondwana. The northern margin of West Gondwana was also an active margin at this time, consistent with the widespread record of subduction in the peri-Gondwanan terranes, whereas that of eastern Laurentia was a developing rift – passive margin (e.g. Cawood et al. 2001). Although the data do not permit the peri-Gondwanan terranes to be accurately positioned along the northern margin of West Gondwana, taken together, they show the terranes to fall into four broad categories that have palaeogeographical implications (Fig. 1). These categories, which are most applicable to the late Neoproterzoic, are based on the isotopic character of their basement. They are: (1) Avalonian-type terranes (e.g. West Avalonia, East Avalonia, Meguma, Carolinia, Moravia –Silesia) that originated from c. 1.3–1.0 Ga juvenile crust within the Panthalassa-like, peri-Rodinian Mirovoi Ocean and were later accreted to the Amazonian margin of Gondwana, dated in Avalonia at c. 665 –650 Ma; (2) Cadomian-type terranes (e.g. North Armorican Massif, Ossa –Morena, SaxoThuringia and Moldanubia) that formed along the West African margin of Gondwana by recycling ancient (c. 2.0– 2.2 Ga) West African crust; (3) Ganderian-type terranes (e.g. Ganderia, Florida, the Maya terrane and possibly the NW Iberian domain and South Armorican Massif) that formed along the Amazonian margin of Gondwana by recycling Avalonian and older Amazonian basement; and (4) cratonic terranes (e.g. Oaxaquia and the Chortis block) that represent displaced Amazonian portions of cratonic Gondwana that lack Neoproterozoic arc magmatism. In addition, evidence exists for movement of some terranes along this margin during the late Neoproterozoic–early Palaeozoic. Data from the Meguma terrane, for example, suggest that it is floored by Avalonian (peri-Amazonian) basement (Keppie et al. 1997) but that it was promixal to West Africa by the Cambrian (Krogh & Keppie 1990). Similarly, detrital zircon and muscovite data from the NW Iberian domain suggest a peri-Amazonian location in the late Neoproterozoic, but a peri-West African position in the early Palaeozoic (Ferna´ndez-Sua´rez et al. 2002b; Gutie´rrez-Alonso et al. 2003, 2005). In our palaeogeographical reconstructions (Fig. 13), we consequently position the Avaloniantype, Ganderian-type and cratonic terranes adjacent to Amazonia in the late Neoproterozoic, whereas
the Cadomian-type terranes are placed adjacent to West Africa. Avalonia is tentatively shown in the inverted orientation of Keppie et al. (2003a) based on the inferred northwestward polarity of subduction (present coordinates) proposed for West Avalonia in Nova Scotia (Dostal et al. 1996). Following Murphy et al. (2004a), the Avaloniatype terranes are considered to have lain some unknown distance outboard of West Gondwana prior to c. 650 Ma in the form of one or more juvenile arcs or oceanic plateau (c. 1 Ga juvenile crust; Fig. 14a) within the Panthalassa-like ocean (Mirovoi) that would have surrounded the supercontinent Rodinia following its amalgamation at c. 1.1 Ga (e.g. Dalziel et al. 2000). The early arc phase in these terranes (Fig. 14b) is likewise attributed to renewed subduction following the breakup of Rodinia at c. 0.75 Ga (Powell et al. 1993; Wingate & Giddings 2000). The collision of the Avalonian-type terranes with the northern margin of West Gondwana by c. 665–650 Ma (Fig. 15a) brings the so-called ‘Avalonian– Cadomian belt’ (e.g. Murphy & Nance 1989) into existence for the first time and broadly coincides with metamorphism and a brief hiatus in arc magmatism in both the Avalonianand Cadomian-type terranes. In this way, the formation of the Avalonian–Cadomian belt is analogous to the Mesozoic –Cenozoic evolution of western North America (e.g. Moores 1998) in that proximal and exotic terranes were incorporated into a single belt (Fig. 15b) that, until the opening of the Rheic Ocean, shared a similar history (e.g. Nance et al. 2002; Keppie et al. 2003a). By c. 640 Ma, the occurrence of abundant ensialic arc-related magmatism in most peri-Gondwanan terranes, together with the presence of Gondwanan detrital zircons, indicates that a subduction zone had been established outboard of the peri-Gondwanan terranes and was angled beneath these accreted terranes and the cratonic margin of West Gondwana (Fig. 13a). According to Murphy et al. (1999a), Nance et al. (2002), and Keppie et al. (2003a), it is the collision of a spreading ridge with the northern margin of West Gondwana, and the consequent migration of a triple point, that is likely to have been responsible for the diachronous cessation of arc-related magmatism and the onset of strike-slip tectonics in the Avalonian- and Cadomian-type terranes between c. 600 and c. 540 Ma (Fig. 13b). Evidence for continuing arc magmatism in the Ganderian-type terranes (and the migrating Avalonian ones like the NW Iberian domain), however, suggests that these were positioned in such a fashion that they experienced continued subduction into the Cambrian. Ridge –trench collision as a mechanism for the Avalonian –Cadomian arc-platform transition
THE PERI-GONDWANAN TERRANES 365
Fig. 13. Late Neoproterozoic– early Palaeozoic reconstructions showing locations of peri-Gondwanan terranes along the northern margin of Gondwana at (a) c. 640–590 Ma, (b) c. 590– 540 Ma and (c) c. 510 –480 Ma. Figure modified from Murphy et al. (2004a) with Avalonia inverted following Keppie et al. (2003a) based on inferred subduction polarity for West Avalonia (Dostal et al. 1996). B, Bohemia; C, Carolinia; Ch, Chortis block; EA, East Avalonia; F, Florida; G, Ganderia; M, Meguma; NAM, North Armorica Massif; NWI, NW Iberian domain; OMZ, Ossa-Morena zone; Ox, Oaxaquia; SAM, South Armorican Massif; WA, West Avalonia; Y, Yucatan block (Maya terrane).
366
R. D. NANCE ET AL.
Fig. 14. Reconstructions of Laurentia – Gondwana – Baltica. (a) Rodinia reconstruction (from Murphy et al. 2004a). Am, Amazonia; Au, Australia; B, Barentsia; Ba, Baltica; Co, Congo-Sa˜o Francisco; Gr, Greenland; In, India; La, Laurentia; O, Oaxaquia; P, Pampean terrane; R, Rockall, RP, Rı´o de la Plata; Ma, Mawson continent (East Antarctica); Ka, Kalahari; SC, South China; Si, Siberia; WA, West Africa. (b) Reconstruction at c. 750 Ma emphasizing the history of the peri-Gondwanan terranes and assigning minimum movement to the continents required to satisfy palaeomagnetic data. Figure modified after Murphy et al. (2001, 2004a) with Avalonia inverted following Keppie et al. (2003a) based on inferred subduction polarity for West Avalonia (Dostal et al. 1996).
THE PERI-GONDWANAN TERRANES
(a)
(b)
Fig. 15. Reconstructions of Laurentia–Gondwana–Baltica at (a) c. 650 Ma and (b) c. 600 Ma, emphasizing the history of the peri-Gondwanan terranes and assigning minimum movement to the continents required to satisfy palaeomagnetic data. Figure modified after Murphy et al. (2001, 2004a) with Avalonia inverted following Keppie et al. (2003a) based on inferred subduction polarity for West Avalonia (Dostal et al. 1996). (See Fig. 13 for identification of major cratons.)
367
368
R. D. NANCE ET AL.
would additionally explain the change from sinistral to dextral motion along basin-bounding faults within Avalonia that occurs at about this time (e.g. Nance & Murphy 1990). It would also permit the local continuation of subduction in areas such as Anglesey in the UK (Gibbons & Hora´k 1996).
(a)
Palaeozoic Although there is broad consensus regarding the proximity of the Cadomian-type terranes to West Africa and that of the Avalonian- and Ganderiantype terranes to Amazonia in the late Neoproterozoic (Fig. 16a), there is considerable disagreement with regard to: (1) the nature of their separation from the northern margin of Gondwana with the opening of the Rheic Ocean in the early Palaeozoic; (2) the manner in which the Avalonian- and Ganderian-type terranes traversed the Iapetus Ocean as the Rheic Ocean widened; and (3) the timing of the subsequent accretion of these terranes to Laurussia with the closure of these oceans in the mid- to late Palaeozoic. By (1) the Early Cambrian in most of the Avalonian- and Cadomian-type terranes, (2) the Middle Cambrian in the Ganderian-type terranes, and (3) the latest Cambrian to mid-Silurian in Oaxaquia, the tectonostratigraphy of the periGondwanan terranes is dominated by stable platform and passive margin assemblages and local rift-related magmatism (e.g. Greenough & Papezik 1986). Faunal data (e.g. Landing 1996) suggest that Avalonia was insular by the Early Cambrian, but that it remained proximal to West Gondwana (Fig. 16b) until the Early Ordovician (Murphy et al. 2006). An Early Ordovician age for the separation of Avalonia from Gondwana (Fig. 13c) is also supported by subsidence data (Prigmore et al. 1997). Some uncertainty exists regarding Cadomia– Gondwana relations in the early Palaeozoic (e.g. Robardet 2003), but isotopic data and detrital zircon ages from Bohemia (e.g. Linnemann 2004; Linnemann et al. 2004) and Iberia (e.g. Martı´nez Catal´n et al. 2004) suggest that Cadomia remained attached to Gondwana until its Variscan collision with Laurussia with the closure of the Rheic Ocean in the Devonian– Carboniferous. The timing of the separation of the Ganderian-type terranes from Gondwana is poorly constrained, but may be recorded in Ganderia by the Middle Cambrian– Early Ordovician arc magmatism and concomitant passive margin sedimentation interpreted to record the development of an arc – back-arc system along its leading edge (van Staal et al. 1998). Given the similarity in their basement isotopic signatures and tectonothermal records prior to c. 570 Ma, it is likely that Avalonia and Carolinia
(b)
Fig. 16. Reconstructions of Laurentia–Gondwana– Baltica at (a) c. 550 Ma and (b) c. 500 Ma, emphasizing the history of the peri-Gondwanan terranes and assigning minimum movement to the continents required to satisfy palaeomagnetic data. Figure modified after Murphy et al. (2001, 2004a) with Avalonia inverted following Keppie et al. (2003a) based on inferred subduction polarity for West Avalonia (Dostal et al. 1996). (See Fig. 13 for identification of major cratons.)
were proximal to each other until this time. This is also likely to be the case for Ganderia, given its tectonothermal history prior to c. 570 Ma, although its somewhat more evolved isotopic signature suggests a basement containing a larger component of older continental crust. In the latest Neoproterozoic and early Palaeozoic, however, contrasts in the tectonic evolution of Ganderia and Avalonia have been taken to imply their existence as independent crustal blocks that traversed Iapetus separately and were accreted to Laurentia at different times (e.g. van Staal et al. 1998; Barr et al. 2002). The tectonic evolution of Carolinia has been similarly
THE PERI-GONDWANAN TERRANES
taken to indicate a linkage to Ganderia rather than Avalonia after c. 570 Ma (Hibbard & van Staal 2005; Hibbard et al. 2007). For Avalonia in an inverted orientation (Fig. 13), the journey across Iapetus must have additionally involved significant clockwise rotation. The accretion of Ganderia and Carolinia to Laurentia is thought to have occurred in the Late Ordovician–Early Silurian by way of sinistrally oblique subduction beneath Laurentia (Hibbard 2000), which subsequently resulted in telescoping of the Ganderian arc– back-arc system (van Staal et al. 1998). However, because Avalonia remained unaffected by orogenesis until the the Acadian orogeny in the Late Silurian –Early Devonian, as described by, for example Dallmeyer & Nance (1994) and Waldron et al. (1996), those workers suggested that the accretion of Avalonia did not occur until this time, the subduction of oceanic crust between Avalonia and Ganderia accounting for the development of the Ganderian arc – back-arc system. Conversely, Murphy et al. (1996b) argued that the accretion of West Avalonia to Laurentia had occurred by the Early Silurian on the basis of sediment isotopic signatures. The closure of Iapetus in Britian is thought to have taken place by c. 420 Ma (e.g. Soper & Woodcock 1990). Available palaeomagnetic data likewise suggest that any palaeolatitudinal separation between Avalonia and Laurentia had disappeared by the mid-Silurian (e.g. Miller & Kent 1988; Trench & Torsvik 1992; Potts et al. 1993; Hodych & Buchan 1998). A potential resolution to this debate is possible if Ganderia is taken to represent the leading edge of Avalonia during its early Palaeozoic traverse across Iapetus (e.g. Keppie et al. 2003a). As such, Ganderia might be expected to develop a distinctive Palaeozoic record of subduction and deformation, and an earlier record of collision, despite travelling with Avalonia as a single crustal block. This interpretation would also be consistent with the development of Ganderia’s extensive early Palaeozoic siliciclastic passive margin succession, which suggests that it was part of a much larger crustal block. In this view, Carolinia would likewise represent part of the leading edge of this Avalonian crustal block, consistent with Late Ordovician palaeomagnetic data that place it at significantly lower palaeolatitudes (2158, Vick et al. 1987; 288, Noel et al. 1988) than contemporary Avalonia (2458, Seguin et al. 1987; 2418, Johnson & Van der Voo 1990; 2328, Hodych & Buchan 1998). The close magmatic and basement isotopic linkage between Carolinia and Avalonia prior to c. 570 Ma but its subsequent tectonothermal similarity to Ganderia would then be accounted for. Indeed, Carolinia may traverse the transition
369
between Ganderian- and Avalonian-type terranes. As the trailing edge of this crustal block, Avalonia would have been accreted to Laurentia with Ganderia and Carolinia by the Early Silurian as part of the dominantly Early Silurian Salinic orogeny rather than during the Late Silurian –Early Devonian Acadian orogeny, which Murphy et al. (1999b) have argued was mantle plume generated in Maritime Canada rather than the result of terrane accretion. Uncertainty also exists with regard to the accretion of the Meguma terrane. This event is generally thought to have been accommodated by dextral transpression on the Minas fault zone in the Middle Devonian –Early Carboniferous (e.g. Culsham & Reynolds 1997; Hicks et al. 1999). However, Murphy et al. (2004b) have argued, based on the similarity between detrital zircon populations in Upper Ordovician –Lower Devonian strata of the Meguma terrane and those of contemporary sedimentary rocks in Avalonia, that the two terranes were attached throughout the Palaeozoic and that Meguma was accreted to Laurentia with Avalonia in the Early Silurian. In this scenario, late Palaeozoic movement on the Minas fault zone would be post-accretionary, perhaps related to oblique subduction along the northern margin of the Rheic Ocean. Florida, Oaxaquia, the Chortis block and Maya terrane probably arrived with Gondwana during the Alleghanian–Ouachita orogeny in the late Palaeozoic, the Maya terrane subsequently rotating southward with the opening of the Gulf of Mexico (e.g. Dickinson & Lawton 2001).
Conclusions The evolution of the peri-Gondwanan terranes of the Appalachian – Variscan orogen provide important constraints on the palaeogeography of the northern (Amazonian – West African) margin of West Gondwana and on palaeocontinental reconstructions for the late Neoproterozoic – early Palaeozoic. Differences in basement isotopic compositions between these terranes constrain their relative palaeogeography in the late Neoproterozoic and allow the geometry of the margin to be reconstructed for this time interval. These differences allow the terranes to be subdivided into four main ‘palaeogeographical’ categories: (1) Avalonian-type terranes (e.g. West Avalonia, East Avalonia, Meguma, Carolinia, Moravia – Silesia) that evolved upon c. 1.3 – 1.0 Ga juvenile crust and were accreted to the Amazonian margin of Gondwanan by c. 665 – 650 Ma; (2) Cadomian-type terranes (e.g. North Armorica, Ossa – Morena, Saxo-Thuringia,
370
R. D. NANCE ET AL.
Moldanubia) that recycled ancient (c. 2.0 – 2.2 Ga) crust along the West African margin of Gondwana; (3) Ganderian-type terranes (e.g. Ganderia, Florida, Maya, (?)NW Iberia, (?) South Armorica) that recycled Avalonian and older crust along the Amazonian margin of Gondwana and presumably lay inboard of the Avalonian-type terranes in the late Neoproterozoic; and (4) cratonic terranes (e.g. Oaxaquia, Chortis) that represent displaced Amazonian portions of cratonic Gondwana. In addition, evidence exists for movement of some terranes (e.g. Meguma, NW Iberia) along this margin during the latest Neoproterozoic – early Palaeozoic, consistent with its interpretation as a transform margin. These differences in crustal signatures are matched in some terranes by palaeogeographical contrasts in Cambrian fauna and sedimentary provenance data. Thus Avalonia has cool-water, high-latitude fauna and detrital zircon signatures consistent with proximity to the Amazonian craton, whereas Ossa – Morena and Cadomia show a transition from tropical to temperate waters and detrital zircon signatures that suggest continuing proximity to West Africa. Both the Ganderian-type and cratonic terranes show linkages to Amazonia. Avalonia probably separated from Gondwana in the Early Ordovician, with or without Ganderia, Carolinia and the Meguma terrane. The Cadomiantype terranes, on the other hand, appear to have remained attached to Gondwana until its collision with Laurussia during the Devonian –Carboniferous Variscan orogeny. Ganderia and Carolinia are thought to have accreted to Laurentia in the Late Ordovician – Early Silurian and this may also be the case for Avalonia, differences in their tectonothermal records reflecting their positions on the leading and trailing edges, respectively, of the same crustal block. However, orogenesis in Avalonia that might reflect its accretion occurred in the Late Silurian – Early Devonian. The accretion of the Meguma terrane may likewise have accompanied that of Avalonia by the Early Silurian, but could have been responsible for the Neo-Acadian orogeny in the Middle Devonian – Early Carboniferous. Florida and the Middle American terranes are thought to have arrived with Gondwana during the late Palaeozoic Alleghanian – Ouachita orogeny, the Maya terrane subsequently rotating southward with the opening of the Gulf of Mexico. Funding for this project was provided by NSF grants (EAR-0308105) to R.D.N. and (EAR-0308437) to B.V.M., NSERC Discovery grants to J.B.M., PAPIIT
(IN103003) and CONACyT grants to J.D.K., and Spanish Education and Science Ministry Project grant (CGL2006-00902 O.D.R.E.) to G.G.A. J.B.M. is also grateful for the support of the University Council of Research, St. Francis Xavier University, and the University of Salamanca during a sabbatical leave in 2005. Constructive comments by P. Cawood and J. Hibbard greatly improved the manuscript. This paper (CTT-TADT 06/1) is a contribution to IGCP Projects 453 and 497. TSRC Publication 394.
References A LLE` GRE , C. J. & B EN O THMAN , D. 1980. Nd– Sr isotopic relationship in granitoid rocks and continental crust development: a chemical approach to orogenesis. Nature, 286, 335– 342. ´ LVARO , J. J., E LICKI , O., G EYER , G., R USHTON , A A. W. A. & S HERGOLD , J. H. 2003. Palaeogeographical controls on the Cambrian trilobite immigration and evolutionary patterns reported in the western Gondwanan margin. Palaeogeography, Palaeoclimatology, Palaeoecology, 195, 5 –35. B ALLARD , M. M., V AN DER V OO , R. & U RRUTIA F UCUGAUCHI , J. 1989. Paleomagnetic results from Grenvillian-aged rocks from Oaxaca, Mexico: evidence for a displaced terrane. Precambrian Research, 42, 343–352. B ANDRES , A., E GUI´ LUZ , L., G IL I BARGUCHI , J. I. & P ALACIOS , T. 2002. Geodynamic evolution of a Cadomian arc region: the northern Ossa– Morena zone, Iberian massif. Tectonophysics, 352, 105 –120. B ARKER , C., S ECOR , D. T., J R ., P RAY , J. & W RIGHT , J. 1998. Age and deformation of the Longtown metagranite, South Carolina Piedmont: A possible constraint on the origin of the Carolina terrane. Journal of Geology, 106, 713–725. B ARR , S. M. & H EGNER , E. 1992. Nd isotopic composition of felsic igneous rocks in Cape Breton Island, Nova Scotia. Canadian Journal of Earth Sciences, 29, 650–657. B ARR , S. M., D UNNING , G. R., R AESIDE , R. P. & J AMIESON , R. A. 1990. Contrasting U–Pb ages from plutons in the Bras d’Or and Mira terranes of Cape Breton Island, Nova Scotia. Canadian Journal of Earth Sciences, 27, 1200–1208. B ARR , S. M., B EVIER , M. L., W HITE , C. E. & D OIG , R., 1994. Magmatic history of the Avalon terrane in southern New Brunswick, Canada, based on U– Pb (zircon) geochronology. Journal of Geology, 102, 399–409. B ARR , S. M., R AESIDE , R. P. & W HITE , C. E. 1998. Geological correlations between Cape Breton Island and Newfoundland, northern Appalachian Orogen. Canadian Journal of Earth Sciences, 35, 1252–1270. B ARR , S. M., W HITE , C. E. & M ILLER , B. V. 2002. The Kingston Terrane, southern New Brunswick, Canada: evidence for an Early Silurian volcanic arc. Geological Society of America Bulletin, 114, 964 –982. B ARR , S. M., D AVIS , D. W., K AMO , S. & W HITE , C. E. 2003a. Significance of U–Pb detrital zircon ages in quartzite from peri-Gondwanan terranes, New
THE PERI-GONDWANAN TERRANES Brunswick and Nova Scotia, Canada. Precambrian Research, 126, 123– 145. B ARR , S. M., W HITE , C. E. & M ILLER , B. V. 2003b. Age and geochemistry of Late Neoproterozoic and Early Cambrian igneous rocks in southern New Brunswick: similarities and contrasts. Atlantic Geology, 39, 55– 73. B ARTHOLOMEW , M. J. & T OLLO , R. P. 2004. Northern ancestry for the Goochland terrane as a displaced fragment of Laurentia. Geology, 32, 669– 672. B EVIER , M. L. & B ARR , S.M. 1990. U– Pb age constraints on the stratigraphy and tectonic history of the Avalon terrane, New Brunswick, Canada. Journal of Geology, 98, 53– 63. B EVIER , M. L., W HITE , C. E. & B ARR , S. M. 1990. Late Precambrian U–Pb ages for the Brookville Gneiss, southern New Brunswick. Journal of Geology, 98, 955–965. B EVIER , M. L., B ARR , S. M., W HITE , C. E. & M ACDONALD , A. S. 1993. U–Pb geochronologic constraints on the volcanic evolution of the Mira (Avalon) terrane, southeastern Cape Breton Island, Nova Scotia. Canadian Journal of Earth Sciences, 30, 1– 10. B LASBAND , B., W HITE , S., B ROOIJMANS , P., D E B OORDER , H. & V ISSER , W. 2000. Late Proterozoic extensional collapse in the Arabian–Nubian Shield. Journal of the Geological Society, London, 157, 615–628. B OHER , M., A BOUCHAMI , W., M ICHARD , A. N., 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 OUCOT , A. J., B LODGETT , R. B. & S TEWART , J. H. 1997. European Province Late Silurian brachiopods from the Ciudad Victoria area, Tamaulipas, northeastern Mexico. In: K LAPPER , G., M URPHY , M. A. & T ALENT , J. A. (eds) Paleozoic Sequence Stratigraphy, Biostratigraphy, and Biogeography: Studies in Honor of J. Grenville (‘Jess’) Johnson. Geological Society of America, Special Papers, 321, 273– 293. B ROWN , M., P OWER , G. M., T OPLEY , C. G. & D’L EMOS , R. S. 1990. Cadomian magmatism in the North Armorican Massif. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 181– 213. B USCHMANN , B. 1995. Geotectonic facies analysis of the Rothstein Formation (Neoproterozoic, Saxothuringian Zone, Germany). PhD thesis, TU Bergakademie Freiburg. B USCHMANN , B., E LICKI , O. & J ONAS , P. 2006. The Cadomian unconformity in the Saxo-Thuringian Zone, Germany: Palaeogeographic affinities of Ediacaran (terminal Neoproterozoic) and Cambrian strata. Precambrian Research, 147, 387– 403. C AMERON , K. L., L OPEZ , R., O RTEGA -G UTIE´ RREZ , F., S OLARI , L. A., K EPPIE , J. D. & S CHULZE , C. 2004. U–Pb constraints and Pb isotopic compositions of leached feldspars: Constraints on the origin and evolution of Grenvillian rocks from eastern and southern Mexico. In: T OLLO , R. P., C ORRIVEAU , L., M C L ELLAND , J. & B ARTHOLOMEW , M. J. (eds) Proterozoic Tectonic Evolution of the Grenville
371
Orogen in North America. Geological Society of America, Memoirs, 197, 755 –770. ¨ , O BERHA¨ NSLI , R., C ANDEN , O., D ORA , O. O C TINKAPLAN , M., P ARTZSCH , J. H., W ARKUS , F. C. & D U¨ RR , S. 2000. Pan-African high-pressure metamorphism in the Precambrian basement of the Menderes Massif, western Anatolia, Turkey. International Journal of Earth Sciences, 89, 793–811. C AWOOD , P. A. & P ISAREVSKY , S. A. 2006. Was Baltica right-way-up or upside-down in the Neoproterozoic? Journal of the Geological Society, London, 163, 753– 759. C AWOOD , P. A., M C C AUSLAND , P. J. A. & D UNNING , G. R. 2001. Opening Iapetus: constraints from the Laurentian margin of Newfoundland. Geological Society of America Bulletin, 113, 443–453. C HANTRAINE , J., A UVRAY , B., B RUN , J. P., C HAUVEL , J. J. & R ABU , D. 1994. The Cadomian Orogeny in the Armorican Massif. In: K EPPIE , J. D. (ed.) PreMesozoic Geology in France and Related Areas. Springer, Berlin, 75–128. C HEN , F., S IEBEL , W., S ATIR , M., T ERZIOLU , N. & S AKA , K. 2002. Geochronology of the Karadere basement (NW Turkey) and implications for the geological evolution of the Istanbul zone. International Journal of Earth Sciences, 91, 469–481. C LARKE , D. B., M AC D ONALD , M. A. & T ATE , M. C. 1997. Late Devonian mafic– felsic magmatism in the Meguma Zone, Nova Scotia. In: S INHA , A.K., W HALEN , J. B. & H OGAN , J. P. (eds) The Nature of Magmatism in the Appalachian Orogen. Geological Society of America, Memoirs, 191, 107–127. C OCHERIE , A., C HANTRAINE , J., F ANNING , C. M., D ABARD , M. -P., P ARIS , F., L E H E´ RISSE , A. & E GAL , E. 2001. Datation U–Pb: aˆge Briove´riaen de la se´rie d’Erquy (Massif armoricain, France). Comptes Rendus de l’Acade´mie des Sciences, 333, 427– 434. C OCKS , L. R. M. 2000. The Early Paleozoic geography of Europe. Journal of the Geological Society, London, 157, 1–10. C OCKS , L. R. M. & F ORTEY , R. A. 1988. Lower Paleozoic facies and faunas around Gondwana. In: A UDLEY -C HARLES , M. G. & H ALLAM , A. (eds) Gondwana and Tethys. Geological Society, London, Special Publications, 37, 183–200. C OCKS , L. R. M. & F ORTEY , R. A. 1990. Biogeography of Ordovician and Silurian faunas. In: M C K ERROW , W. S. & S COTESE , C. R. (eds) Paleozoic Paleogeography and Biogeography. Geological Society, London, Memoirs, 12, 97–104. C OMPSTON , W., W RIGHT , A. E. & T OGHILL , P. 2002. Dating the Late Precambrian volcanicity of England and Wales. Journal of the Geological Society, London, 159, 323– 339. C OWIE , J. R. 1974. The Cambrian of Spitzbergen and Scotland. In: H OLLAND , C. H. (ed.) Cambrian of the British Isles, Norden and Spitzbergen: Lower Paleozoic Rocks of the World, Vol. 2. Wiley, Chichester, 123– 156. C ULSHAW , N. & R EYNOLDS , P. 1997. 40Ar/39Ar age of shear zones in the southwest Meguma Zone between Yarmouth and Meteghan, Nova Scotia. Canadian Journal of Earth Sciences, 34, 848– 853.
372
R. D. NANCE ET AL.
C URRIE , K. L. & M C N ICOLL , V. J. 1999. New data on the age and geographic distribution of Neoproterozoic plutons near Saint John, New Brunswick. Atlantic Geology, 35, 157–166. D ALLMEYER , R. D. 1989. Contrasting accreted terranes in the southern Appalachian orogen, basement beneath the Atlantic and Gulf Coastal Plains, and West African orogens. Precambrian Research, 42, 387– 409. D ALLMEYER , R. D. & K EPPIE , J. D. 1987. Polyphase late Palaeozoic tectonothermal evolution of the southwestern Meguma Terrane, Nova Scotia: evidence from 40 Ar/39Ar mineral ages. Canadian Journal of Earth Sciences, 24, 1242–1254. D ALLMEYER , R. D. & N ANCE , R. D. 1992. Tectonic implications of 40Ar/39Ar mineral ages from late Precambrian–Cambrian plutons, Avalon composite terrane, southern New Brunswick, Canada. Canadian Journal of Earth Sciences, 29, 2445–2462. D ALLMEYER , R. D. & N ANCE , R. D. 1994. 40Ar/39Ar whole-rock phyllite ages from late Precambrian rocks of the Avalon composite terrane, New Brunswick: evidence of Silurian–Devonian thermal rejuvenation. Canadian Journal of Earth Sciences, 31, 818– 824. D ALLMEYER , R. D., S TRACHAN , R. A. & D’L EMOS , R. S. 1991. Chronology of Cadomian tectonothermal activity in the baie de Saint-Brieuc (north Brittany), France: evidence from 40Ar/39Ar mineral ages. Canadian Journal of Earth Sciences, 28, 762– 773. D ALZIEL , I. A. D. 1991. Pacific margins of Laurentia and East Antarctica–Australia as a conjugate rift pair; evidence and implications for an Eocambrian supercontinent. Geology, 19, 598– 601. D ALZIEL , I. A. D. 1992. On the organization of American plates in the Neoproterozoic and the breakout of Laurentia. GSA Today, 2, 238 –241. D ALZIEL , I. W. D. 1994. Precambrian Scotland as a Laurentia–Gondwana link; origin and significance of cratonic promontories. Geology, 22, 589–592. D ALZIEL , I. W. D. 1997. Overview: Neoproterozoic– Paleozoic geography and tectonics: review, hypotheses and environmental speculations. Geological Society of America Bulletin, 109, 16–42. D ALZIEL , I. W. D., M OSHER , S. & G AHAGAN , L. M. 2000. Laurentia–Kalahari collision and the assembly of Rodinia. Journal of Geology, 108, 499–513. D ENNIS , E. & D ABARD , M. P. 1988. Sandstone petrography and geochemistry of late Proterozoic sediments of the Armorican Massif (France)—a key to basin development during the Cadomian Orogeny. Precambrian Research, 42, 189 –206. D ENNIS , A. & S HERVAIS , J. 1991. Arc rifting of the Carolina terrane in northwestern South Carolina. Geology, 19, 226–229. D ENNIS , A. & S HERVAIS , J. 1996. The Carolina terrane in northwestern South Carolina: Insights into the development of an evolving island arc. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society America Special Papers, 304, 237–256. D ENNIS , A. & W RIGHT , J. E. 1997. The Carolina terrane in northwestern South Carolina: age of deformation
and metamorphism in an exotic arc. Tectonics, 16, 460–473. D ENNIS , A. J., S HERVAIS , J. W., M AULDIN , J., M AHER , H. D., J R & W RIGHT , J. E. 2004. Petrology and geochemistry of Neoproterozoic volcanic arc terranes beneath the Atlantic Coastal Plain, Savannah River Site, South Carolina. Geological Society of America Bulletin, 116, 572– 593. DE W IT , M. J., B OWRING , S., D UDAS , F. & K AMGA , G. 2005. The great Neoproterozoic Central Saharan arc and the amalgamation of the North African Shield. GAC– MAC– CSPG– CSSS Joint Meeting, Halifax, Nova Scotia, Abstracts, 30, 42–43. D EYNOUX , M., A FFATON , P., T ROMPETTE , R. & V ILLENEUVE , M. 2006. Pan-African tectonic evolution and glacial events registered in Neoproterozoic to Cambrian cratonic and foreland basins of West Africa. Journal of African Earth Sciences, 46, 397–426. D ´I AZ G ARCI´ A , F. 2006. Geometry and regional significance of Neoproterozoic (Cadomian) structures of the Narcea Antiform, NW Spain. Journal of the Geological Society, London, 163, 499 –508. D ICKINSON , W. R. & L AWTON , T. F. 2001. Carboniferous to Cretaceous assembly and fragmentation of Mexico. Geological Society of America Bulletin, 113, 1142–1160. D’L EMOS , R. S. & B ROWN , M. 1993. Sm–Nd isotope characteristics of late Cadomian granite magmatism in northern France and the Channel Islands. Geological Magazine, 130, 797– 804. D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) 1990. The Cadomian Orogeny. Geological Society, London, Special Publications, 51. D’L EMOS , R. S., B ROWN , M. & S TRACHAN , R. A. 1992. Granite magma generation, ascent and emplacement within a transpressional orogen. Journal of the Geological Society, London, 149, 487 –490. D’L EMOS , R. S., M ILLER , B. V. & S AMSON , S. D. 2001. Precise U–Pb zircon ages from Alderney, Channel Islands: growing evidence for discrete Neoproterozoic magmatic episodes in northern Cadomia. Geological Magazine, 138, 719– 726. 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. D OIG , R., M URPHY , J. B. & N ANCE , R. D. 1991. U– Pb geochronology of Late Proterozoic rocks of the eastern Cobequid Highlands, Avalon Composite Terrane, Nova Scotia. Canadian Journal of Earth Sciences, 28, 504–511. D OIG , R., M URPHY , J. B. & N ANCE , R. D. 1993. Tectonic significance of the Late Proterozoic Economy River Gneiss, Cobequid Highlands, Avalon composite terrane, Nova Scotia. Canadian Journal of Earth Sciences, 30, 474–479. D O¨ RR , W., Z ULAUF , G., F IALA , J., F RANKE , W. & V EJNAR , Z. 2002. Neoproterozoic to Early Cambrian history of an active plate margin in the Tepla´ – Barrandian unit—a correlation of U– Pb isotopicdilution– TIMS ages (Bohemia, Czech Republic). Tectonophysics, 352, 65– 85.
THE PERI-GONDWANAN TERRANES D O¨ RR , W., F INGER , F., L INNEMANN , U. & Z ULAUF , G. 2004. The Avalonian–Cadomian Belt and related periGondwanan terranes. International Journal of Earth Sciences, 93, 657–658. D OSTAL , J., K EPPIE , J. D., C OUSENS , B. L. & M URPHY , J. B. 1996. 550– 580 Ma magmatism in Cape Breton Island (Nova Scotia, Canada): the product of NW-dipping subduction during the final stage of assembly of Gondwana. Precambrian Research, 76, 96– 113. D ROST , K., L INNEMANN , U., M C N AUGHTON , N. ET AL . 2004. New data on the Neoproterozoic– Cambrian geotectonic setting of the Tepla´ –Barrandian volcanosedimentary successions: geochemistry, U– Pb zircon ages, and provenance (Bohemian Massif, Czech Republic). International Journal of Earth Sciences, 93, 742–757. D UNNING , G. R. & K ROGH , T. E. 1985. Geochronology of ophiolites in the Newfoundland Appalachians. Canadian Journal of Earth Sciences, 22, 1659–1670. D UNNING , G. R. & O’B RIEN , S. J. 1989. Late Proterozoic–early Paleozoic crust in the Hermitage flexure, Newfoundland Appalachians: U–Pb ages and tectonic significance. Geology, 17, 548 –551. D UNNING , G. R., B ARR , S. M., R AESIDE , R. P. & J AMIESON , R. A. 1990. U– Pb zircon, titanite and monazite ages in the Bras d’Or and Aspy terranes of Cape Breton Island, Nova Scotia: implications for magmatic and metamorphic history. Geological Society of America Bulletin, 102, 322– 330. D UNNING , G. R., S WINDEN , H. S., K EAN , B. F., E VANS , D. T. & J ENNER , G. A. 1991. A Cambrian island arc in Iapetus; geochronology and geochemistry of the Lake Ambrose volcanic belt, Newfoundland Appalachians. Geological Magazine, 128, 1 –17. E GAL , E., G UERROT , C., L E G OFF , D., T HIE´ BLEMONT , D. & C HANTRAINE , J. 1996. The Cadomian orogeny revisited in northern France. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 281–318. E GUI´ LUZ , L., G IL I BARGUCHI , J. I., A´ BALOS , B. & A PRAIZ , A. 2000. Superposed Hercynian and Cadomian orogenic cycles in the Ossa–Morena zone and related areas of the Iberian Massif. Geological Society of America Bulletin, 112, 1398–1413. E VANS , D. T. W., K EAN , B. F. & D UNNING , G. R. 1990. Geological studies, Victoria Lake Group, Central Newfoundland. Current Research, Newfoundland Department of Mines and Energy, Geological Survey Branch, Report, 90-1, 131 –144. F ARRAR , S. S., 1984. The Goochland granulite terrane; remobilized Grenville basement in the eastern Virginia Piedmont. In: B ARTHOLOMEW , M. J. (ed.) Grenville Event in the Appalachians and Related Topics. Geological Society of America, Special Papers, 194, 215–227. F ERNA´ NDEZ -S UA´ REZ , J., G UTIE´ RREZ -A LONSO , G., J ENNER , G. A. & J ACKSON , S. 1998. Geochronology and geochemistry of the Pola de Allande granitoids (northern Spain). The bearing on the Cadomian/ Avalonian evolution of NW Iberia. Canadian Journal of Earth Sciences, 35, 1–15.
373
F ERNA´ NDEZ -S UA´ REZ , J., G UTIE´ RREZ -A LONSO , G., J ENNER , G. A. & T UBRETT , M. N. 2000. New ideas on the Proterozoic –early Palaeozoic evolution of NW Iberia; insights from U–Pb detrital zircon ages. Precambrian Research, 102, 185–206. F ERNA´ NDEZ -S UA´ REZ , J., G UTIE´ RREZ A LONSO , G. & J EFFRIES , T. E. 2002a. The importance of alongmargin terrane transport in northern Gondwana: insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany. Earth and Planetary Science Letters, 204, 75– 88. F ERNA´ NDEZ -S UA´ REZ , J., G UTIE´ RREZ -A LONSO , G., C OX , R. & J ENNER , G. A. 2002b. Assembly of the Armorica Microplate: A strike-slip terrane delivery? Evidence from U–Pb ages of detrital zircons. Journal of Geology, 110, 619 –626. F INGER , F., H ANZL , P., P IN , C., VON Q UADT , A. & S TEYRER , H. P. 2000. The Brunovistulian: Avalonian Precambrian sequence at the eastern end of the Central European Variscides? In: F RANKE , W., H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Orogenic Processes: Quantification and Modelling in the Variscan Belt. Geological Society, London Special Publications, 179, 103– 112. F ORTEY , R. A. & C OCKS , L. R. M. 2003. Palaeontological evidence bearing on global Ordovician–Silurian continental reconstructions. Earth-Science Reviews, 61, 245– 307. F RANKE , W. 1989. Tectonostratigraphic units in the Variscan belt of Central Europe. In: D ALLMEYER , R. D. & K EPPIE , J. D. (eds) Terranes in the Circum-Atlantic Paleozoic Orogens. Geological Society of America, Special Papers, 230, 67– 90. F RIEDL , G., F INGER , F., M C N AUGHTON , N. J. & F LETCHER , I. R. 2000. Deducing the ancestry of terranes: SHRIMP evidence for South America-derived Gondwana fragments in central Europe. Geology, 28, 1035– 1038. F RIEDL , G., F INGER , F., P AQUETTE , J. L., VON Q UADT , A., M C N AUGHTON , N. J. & F LETCHER , I. R. 2004. Pre-Variscan geological events in the Austrian part of the Bohemian Massif deduced from U– Pb zircon ages. International Journal of Earth Sciences, 93, 802– 823. G APAIS , D. & B ALE´ , P. 1990. Shear zone pattern and granite emplacement within a Cadomian sinistral wrench zone at St. Cast, N. Brittany. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 169–180. G EBAUER , D. & F RIEDL , G. 1994. A 1.38 Ga protolith age for the Dobra orthogneiss (Moldanubian zone of the southern Bohemian massif, NE Austria): Evidence from ion-microprobe (SHRIMP)-dating of zircon. European Journal of Mineralogy, 5, 115. G EYER , G. & L ANDING , E. 2001. Middle Cambrian of Avalonian Massachusetts: Stratigraphy and correlation of the Braintree trilobites. Journal of Paleontology, 75, 116– 135. G IBBONS , W. & H ORA´ K , J. M. 1996. The evolution of the Neoproterozoic Avalonian subduction system: Evidence from the British Isles. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic.
374
R. D. NANCE ET AL.
Geological Society of America, Special Papers, 304, 269– 280. G IESE , U. & B UEHN , B. 1994. Early Paleozoic rifting and bimodal volcanism in the Ossa–Morena Zone of south-west Spain. Geologische Rundschau, 83, 143– 160. G ILLIS , R. J., G EHRELS , G. E., F LORES DE D IOS , A. & R UIZ , J. 2001. Paleogeographic implications of detrital zircon ages from the Oaxaca terrane of southern Mexico. Geological Society of America, Abstracts with Programs, 33, A-428. G OWER , C. F., R YAN , A. B. & R IVERS , T. 1990. MidProterozoic Laurentia– Baltica: An overview of its geological evolution and summary of the contributions by this volume. In: G OWER , C. F., R IVERS , T. & R YAN , A. B. (eds) Mid-Proterozoic Laurentia – Baltica. Geological Association of Canada, Special Papers, 38, 1– 20. G RADSTEIN , F. M., O GG , J. G., S MITH , A. G., B LEEKER , W. & L OURENS , L. J. 2004. A new geologic time scale, with special reference to Precambrian and Neogene. Episodes, 27, 83–100. G RAF , J., VON Q UADT , A., B ERNOULLI , D. & B URG , J-P. 1998. Geochemistry and geochronology of igneous rocks of the central Serbo-Macedonian Massif (Western Bulgaria). Abstracts, Carpathian –Balkan Geological Association XVI Congress. Geological Survey of Austria, Vienna, 191. G REENOUGH , J. G. & P APEZIK , V. S. 1986. Acado-Baltic volcanism in eastern North America and western Europe: Implications for Cambrian tectonism. Maritime Sediments and Atlantic Geology, 22, 240– 251. G UERROT , C. & P EUCAT , J. J. 1990. U–Pb geochronology of the Upper Proterozoic Cadomian orogeny in the northern Armorican Massif, France. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 13– 26. G UTIE´ RREZ -A LONSO , G., F ERNA´ NDEZ -S UA´ REZ , J., J EFFRIES , T. E., J ENNER , G. A., T UBRETT , M. N., C OX , R. & J ACKSON , S. E. 2003. Terrane accretion and dispersal in the northern Gondwana margin. An Early Paleozoic analogue of a long-lived active margin. Tectonophysics, 365, 221–232. G UTIE´ RREZ -A LONSO , G., F ERNA´ NDEZ -S UA´ REZ , J., C OLLINS , A. S., A BAD , I. & N IETO , F. 2005. Amazonian Mesoproterozoic basement in the core of the Ibero-Armorican Arc: 40Ar/39Ar detrital mica ages complement the zircon’s tale. Geology, 33, 637– 640. H AINES , P. W., W ARTHO , J.-A., S HERLOCK , S. C., T URNER , S. P. & K ELLEY , S. P. 2004. 40Ar– 39Ar dating of detrital muscovite in provenance investigations: A case study from the Adelaide Rift Complex, South Australia. Earth and Planetary Science Letters, 227, 297– 311. H AVLI´ CEK , V. 1999. Perunica microplate: relation to Ukranian Shield, mid-Bohemian rift, and hypothetic large-scale overthrusts in central Bohemia. Bulletin of the Czech Geological Survey, 74, 75– 81. H EATHERINGTON , A. L. & M UELLER , P. A. 1999. Lithospheric sources of North Florida, USA tholeiites and implications for the origin of the Suwannee terrane. Lithos, 46, 215–233.
H EATHERINGTON , A. L. & M UELLER , P. A. 2005. Common Pb and Sm–Nd isotopic signatures of the Suwannee terrane and adjacent terranes of the southeastern US: Comparison and speculations on interterrane correlations. Geological Society of America, Abstracts with Programs, 37, 7. H EATHERINGTON , A. L., M UELLER , P. A. & N UTMAN , A. P. 1996. Neoproterozoic magmatism in the Suwannee terrane: Implications for terrane correlation. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related Peri-Gondwanan Terranes of the CircumNorth Atlantic. Geological Society of America, Special Papers, 304, 257– 268. H EGNER , E. & K RO¨ NER , A. 2000. Review of Nd isotopic data and xenocrystic and detrital zircon ages from the pre-Variscan basement in the eastern Bohemian Massif: speculations on palinspastic reconstructions. In: F RANKE , W., H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Orogenic Processes: Quantification and Modelling in the Variscan Belt. Geological Society, London, Special Publications, 179, 113– 129. H ERMES , O. D. & Z ARTMAN , R. E. 1985. Late Proterozoic and Devonian plutonic terrane within the Avalon zone of Rhode Island. Geological Society of American Bulletin, 96, 272– 282. H ERMES , O. D. & Z ARTMAN , R. E. 1992. Late Proterozoic and Silurian alkaline plutons within the southeastern New England Avalon zone. Journal of Geology, 100, 477– 486. H IBBARD , J. 2000. Docking Carolina: Mid-Paleozoic accretion in the southern Appalachians. Geology, 28, 127–130. H IBBARD , J. & S AMSON , S. 1995. Orogenesis exotic to the Iapetan cycle in the southern Appalachians. In: H IBBARD , J., VAN S TAAL , C. & C AWOOD , P. (eds) Current Perspectives in the Appalachian –Caledonian Orogen. Geological Association of Canada, Special Papers, 41, 191–205. H IBBARD , J. & VAN S TAAL , C. 2005. A Carolina– Ganderia link in the Appalachian peri-Gondwanan realm: Implications for the terminal Taconic plate configuration. Geological Society of America, Abstracts with Programs, 37, 16. H IBBARD , J. P., S TODDARD , E. F., S ECOR , D. T. & D ENNIS , A. J. 2002. The Carolina Zone: overview of Neoproterozoic to Early Paleozoic peri-Gondwanan terranes along the eastern flank of the southern Appalachians. Earth-Science Reviews, 57, 299– 339. H IBBARD , J. P., VAN S TAAL , C. R. & M ILLER , B. V. 2007. Links between Carolinia, Avalonia and Ganderia in the Appalachian peri-Gondwanan realm. In: S EARS , J. W., H ARMS , T. A. & E VENCHICK , C. A. (eds) Whence the Mountains? Inquiries into the Evolution of Orogenic Systems: A Volume in Honor of Raymond A. Price. Geological Society of America, Special Papers, 433, 291– 312. H ICKS , R. J., J AMIESON , R. A. & R EYNOLDS , P. 1999. Detrital and metamorphic 40Ar/39Ar ages from muscovite and whole-rock samples, Meguma Supergroup, southern Nova Scotia. Canadian Journal of Earth Sciences, 36, 23–32. H IMMERKUS , F., A NDRES , B., R EISCHMANN , T. & K OSTOPOULOS , D. 2007. Gondwana-derived terranes in the northern Hellenides. In: H ATCHER J R , R. D.,
THE PERI-GONDWANAN TERRANES C ARLSON , M. P., M C B RIDE , J. H. & M ARTI´ NEZ C ATALA´ N , J. R. (eds) 4-D Framework of Continental Crust. Geological Society of America, Memoirs, 200, 379–390. H IRDES , W. & D AVIS , D.W. 2002. U– Pb geochronology of Paleoproterozoic rocks in the southern part of the Kedougou –Ke´nie´ba inlier, Senegal, West Africa: Evidence for diachronous accretionary development of the Eburnian province. Precambrian Research, 118, 83–99. H ODYCH , J. & B UCHAN , K. 1998. Palaeomagnetism of the ca. 440 Ma Cape St. Mary’s sills of the Avalon Peninsula, Newfoundland: implications for Iapetus Ocean closure. Geophysical Journal International, 135, 155– 164. H ODYCH , J. P., C OX , R. A. & K OSLER , J. 2004. An equatorial Laurentia at 550 Ma confirmed by Grenvillian inherited zircons dated by LAM ICP-MS in the Skinner Cove volcanics of western Newfoundland: implications for inertial interchange true polar wander. Precambrian Research, 129, 93– 113. H OFFMAN , P. F. 1991. Did the breakout of Laurentia turn Gondwanaland inside out? Science, 252, 1409–1412. H OFFMAN , P. F., K AUFMAN , A. J., H ALVERSON , G. P. & S CHRAG , D. P. 1998. A Neoproterozoic snowball Earth. Science, 281, 1342–1346. H ORA´ K , J. M. 1993. The late Precambrian Coedana and Sarn Complexes, northwest Wales—A geochemical and petrologic study. PhD thesis, University of Wales, Cardiff. I NGLE , S., M UELLER , P., H EATHERINGTON , A. & K OZUCH , M. 2003. Isotopic evidence for the magmatic and tectonic histories of the Carolina terrane: implications for stratigraphy and terrane affiliation. Tectonophysics, 371, 187 –211. I NGLIS , J. D., S AMSON , S. D., D’L EMOS , R. S. & H AMILTON , M. 2004. U– Pb geochronological constraints on the tectonothermal evolution of the Paleoproterozoic basement of Cadomia, La Hague, NW France. Precambrian Research, 134, 293 –315. 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 AMES , N. P., N ARBONNE , G. M., D ALRYMPLE , R. W. & K YSER , T. K. 2005. Glendonites in Neoproterozoic low-latitude, interglacial sedimentary rocks, northwest Canada: Insights into the Cryogenian ocean and Precambrian cold-water carbonates. Geology, 33, 9– 12. J OHNSON , S. 2001. Contrasting geology in the Pocologan River and Long Reach areas: implications for the New River belt and correlations in southern New Brunswick and Maine. Atlantic Geology, 37, 61– 79. J OHNSON , R. J. E. & V AN DER V OO , R. 1986. Paleomagnetism of the Late Precambrian Fourchu Group, Cape Breton Island, Nova Scotia. Canadian Journal of Earth Sciences, 23, 1673– 1685. J OHNSON , R. J. E. & V AN DER V OO , R. 1990. Pre-folding magnetization reconfirmed for the Late Ordovician– Early Silurian Dunn Point volcanics, Nova Scotia. Tectonophysics, 178, 193 –205.
375
K ARABINOS , P. & G ROMET , L. P. 1993. Application of single-grain zircon evaporation analyses to detrital zircon grain studies and age discrimination in ignous suites. Geochimica et Cosmochimica Acta, 57, 4257– 4267. K AYE , C. A. & Z ARTMAN , R. E. 1980. A late Proterozoic Z to Cambrian age for the stratified rocks of the Boston Basin, Massachusetts, U.S.A. In: W ONES , D. R. (ed.) The Caledonides of the U.S.A. IGCP Project 27, Caledonide Orogen, 1979 Meeting Proceedings, Blacksburg, VA. Department of Geological Sciences, Virginia Polytechnical Institute and State University Memoir, 2, 257–261. K EEN , C. E., K EEN , M. J., N ICHOLS , B. ET AL . 1986. Deep seismic reflection profile across the northern Appalachians. Geology, 14, 141–145. K EMNITZ , H. 2007. The Lausitz greywackes, Saxo-Thuringia, Germany—witness to the Cadomian orogeny. In: L INNEMANN , U., K RAFT , P., N ANCE , R. D. & Z ULAUF , G. (eds) The Geology of PeriGondwana: The Avalonian– Cadomian Belt, Adjoining Cratons and the Rheic Ocean. Geological Society of America, Special Papers, 423, 97–142. K EPPIE , J. D. 1977. Plate tectonic interpretation of Paleozoic world maps (with emphasis on circum-Atlantic orogens and southern Nova Scotia). Nova Scotia Department of Mines Paper, 77–3. K EPPIE , J. D. 1993. Synthesis of Paleozoic deformational events and terrane accretion in the Canadian Appalachians. Geologische Rundschau, 82, 381 –431. K EPPIE , J. D. 2004. Terranes of Mexico revisited: A 1.3 billion year odyssey. International Geology Review, 46, 765– 794. K EPPIE , J. D. & D ALLMEYER , R. D. 1995. Late Paleozoic collision, delamination, short-lived magmatism and rapid denudation in the Meguma terrane (Nova Scotia, Canada): constraints from 40Ar/39Ar isotopic data. Canadian Journal of Earth Sciences, 2, 644– 659. K EPPIE , J. D. & D OSTAL , J. 1998. Birth of the Avalonian arc in Nova Scotia, Canada: geochemical evidence for 700–630 Ma back-arc rift volcanism off Gondwana. Geological Magazine, 135, 171–181. K EPPIE , J. D. & K ROGH , T. E. 2000. 440 Ma igneous activity in the Meguma terrane, Nova Scotia, Canada: part of the Appalachian overstep sequence? American Journal of Science, 300, 528– 538. K EPPIE , J. D. & O RTEGA -G UTIE´ RREZ , F. 1995. Provenance of Mexican terranes: Isotopic constraints. International Geology Review, 37, 813– 824. K EPPIE , J. D. & O RTEGA -G UTIE´ RREZ , F. 1999. Middle American Precambrian basement: A missing piece of the reconstructed 1-Ga orogen. In: R AMOS , V. A. & K EPPIE , J. D. (eds) Laurentia–Gondwana Connections Before Pangea. Geological Society of America, Special Papers, 336, 199–210. K EPPIE , J. D. & R AMOS , V. A. 1999. Odyssey of terranes in the Iapetus and Rheic oceans during the Paleozoic. In: R AMOS , V. A. & K EPPIE , J. D. (eds) Laurentia– Gondwana Connections Before Pangea. Geological Society of America, Special Papers, 336, 267– 276. K EPPIE , J. D., N ANCE , R. D., M URPHY , J. B. & D OSTAL , J. 1991. Northern Appalachians: Avalon and Meguma
376
R. D. NANCE ET AL.
Terranes. In: D ALLMEYER , R. D. & L E´ CORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, New York, 315–334. K EPPIE , J. D., C OUSENS , B. L., D OSTAL , J. & M URPHY , J. B. 1997. Palaeozoic within-plate volcanic rocks in Nova Scotia (Canada) reinterpreted; isotopic constraints on magmatic source and palaeocontinental reconstructions. Geological Magazine, 134, 425–447. K EPPIE , J. D., D AVIS , D. W. & K ROGH , T. E. 1998. U– Pb geochronological constraints on Precambrian stratified units in the Avalon composite terrane of Nova Scotia, Canada: Tectonic implications. Canadian Journal of Earth Sciences, 35, 222– 236. K EPPIE , J. D., D OSTAL , J., D ALLMEYER , R. D. & D OIG , R. 2000. Superposed Neoproterozoic and Silurian magmatic arcs in central Cape Breton Island, Canada: geochemical and geochronological constraints. Geological Magazine, 137, 137–153. K EPPIE , J. D., D OSTAL , J., O RTEGA -G UTIE´ RREZ , F. & L OPEZ , R. 2001. A Grenvillian arc on the margin of Amazonia: evidence from the southern Oaxacan Complex, southern Mexico. Precambrian Research, 112, 165–181. K EPPIE , J. D., N ANCE , R. D. & M URPHY , J. B. & Dostal J. 2003a. Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic –Paleozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia. Tectonophysics, 365, 195–219. K EPPIE , J. D., D OSTAL , J., C AMERON , K. L., S OLARI , L. A., O RTEGA -G UTIE´ RREZ , F. & L OPEZ , R. 2003b. Geochronology and geochemistry of Grenvillian igneous suites in the northern Oaxacan Complex, southern Mexico: tectonic implications. Precambrian Research, 120, 365– 389. K EPPIE , J. D., D OSTAL , J., N ANCE , R. D., M ILLER , B. V., O RTEGA -R IVERA , A. & L EE , J. K. W. 2006. Circa 546 Ma plume-related dykes in the 1 Ga Novillo Gneiss (east–central Mexico): Evidence for the initial separation of Avalonia. Precambrian Research, 147, 342–353. K ERR , A., J ENNER , G. A. & F RYER , B. J. 1995. Sm–Nd isotopic geochemistry of Precambrian to Paleozoic granitoid suites and the deep-crustal structure of the southeast margin of the Newfoundland Appalachians. Canadian Journal of Earth Sciences, 32, 224– 245. K NOLL , A. H. 1992. The early evolution of eukaryotes, a geological perspective. Science, 256, 622–637. K RA¨ UTNER , H. G. 1993. Pre-Alpine evolution in the Southern Carpathians and adjacent areas. Geologica Carpathica, 44, 203–212. K ROGH , T. E. & K EPPIE , J. D. 1990. Age of detrital zircon and titanite in the Meguma Group, southern Nova Scotia, Canada: clues to the origin of the Meguma Terrane. Tectonophysics, 177, 307– 323. K ROGH , T. E., S TRONG , D. F., O’B RIEN , S. J. & P APEZIK , V. S. 1988. Precise U– Pb zircon dates from the Avalon Terrane in Newfoundland. Canadian Journal Earth Sciences, 25, 442– 453. K ROGH , T. E., K AMO , S. L., S HARPTON , B., M ARIN , L. & H ILDEBRAND , A. R. 1993. U– Pb ages of single shocked zircons linking distal K/T ejecta to the Chicxulub crater. Nature, 366, 731– 734.
K RO¨ NER , A. & S ENGO¨ R , A. M. C. 1990. Archean and Proterozoic ancestry in late Precambrian to early Paleozoic crustal elements of southern Turkey as revealed by single-zircon dating. Geology, 18, 1186– 1190. K RO¨ NER , A., W ENDT , I. J., L IEW , T. C. ET AL . 1988. U–Pb zircon and Sm– Nd model ages of high grade Moldanubian sediments, Bohemian massif, Czechoslovakia. Contributions to Mineralogy and Petrology, 99, 257– 266. L ANDING , E. 1996. Avalon: Insular continent by the latest Precambrian. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 29– 63. L ANDING , E. 2004. Precambrian –Cambrian boundary interval deposition and the marginal platform of the Avalon microcontinent. Journal of Geodynamics, 37, 411–435. L ANDING , E. 2005. Early Paleozoic Avalon–Gondwana unity: An obituary—Response to ‘Palaeontological evidence bearing on global Ordovician– Silurian continental reconstructions’ by R. A. Fortey and L. R. M. Cocks. Earth-Science Reviews, 69, 169– 175. L ANDING , E., K EPPIE , J. D. & W ESTROP , S. R. 2006. Lower Paleozoic of northwest Gondwana—terminal Cambrian and lowest Ordovician Tin˜u Formation in southern Mexico. Geological Society of America, Abstracts with Programs, 38, 21. L IN˜ A´ N , E. & G A´ MEZ -V INTANED , J. A. 1993. Lower Cambrian palaeogeography of the Iberian Peninsula and its relations with some neighbouring European areas. Bulletin de la Socie´te´ Ge´ologique de France, 164, 831– 842. L INNEMANN , U. 1991. Glazioeustatisch kontrollierte Sedimentationsprozesse im Oberen Proterozoikum der Elbezone (Weesensteiner Gruppe/Sachsen). Zentralblatt fu¨r Geologie und Pala¨ontologie, Teil I, 12, 2907– 2934. L INNEMANN , U. 1995. The Neoproterozoic terranes of Saxony (Germany). Precambrian Research, 73, 235–250. L INNEMANN , U. 2004. Sedimentation und geotektonischer Rahmen der Beckenentwicklung im Saxothuringikum (Neoproterozoikum–Unterkarbon). In: L INNEMANN , U. (ed.) Das Saxothuringikum. Geologica Saxonica, 48–49, 71–110. L INNEMANN , U. & R OMER , R. L. 2002. The Cadomian Orogeny in Saxo-Thuringia, Germany: geochemical and Nd–Sr–Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance. Tectonophysics, 352, 33–64. L INNEMANN , U., G EHMLICH , M., T ICHOMIROWA , M. ET AL . 2000. From Cadomian subduction to Early Palaeozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (Central European Variscides, Germany). In: F RANKE , W., A LTHERR , R., H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Orogenic Processes—Quantification and Modelling in the Variscan Belt. Geological Society, London Special Publications, 179, 131–153. L INNEMANN , U., M C N AUGHTON , N. J., R OMER , R. L., G EHMLICH , M., D ROST , K. & T ONK , C. 2004. West
THE PERI-GONDWANAN TERRANES African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana?—U–Pb-SHRIMP zircon evidence and the Nd-isotopic record. International Journal of Earth Sciences, 93, 683–705. L INNEMANN , U., N ANCE , R. D., Z ULAF , G. & K RAFT , P. (eds) 2007a. The Geology of Peri-Gondwana, The Avalonian –Cadomian Belt, Adjoining Cratons and the Rheic Ocean. Geological Society of America, Special Papers, 423. L INNEMANN , U., G ERDES , A., D ROST , K. & B USCHMANN , B. 2007b. Cadomian orogenic processes—The ultimate cause for the opening of the Rheic Ocean: Constraints by Laser Ablation-ICP-MS U–Pb zircon dating and analysis of the geotectonic setting (Saxo-Thuringian zone, Bohemian massif, Germany). In: L INNEMANN , U., K RAFT , P., N ANCE , R. D. & Z ULAUF , G. (eds) The Geology of Peri-Gondwana: The Avalonian– Cadomian Belt, Adjoining Cratons and the Rheic Ocean. Geological Society of America, Special Papers, 423, 61– 96. L OEWY , S. L., C ONNELLY , J. N. & D ALZIEL , I. W. D. 2004. An orphaned basement block: The Arequipa– Antofalla Basement of the central Andean margin of South America. Geological Society of America Bulletin, 116, 171– 187. M AC N IOCAILL , C., VAN DER P LUIJM , B. A., V AN DER V OO , R. & M C N AMARA , A. K. 2002. Reply: West African proximity of Avalon in the latest Precambrian. Geological Society of America Bulletin, 114, 1051–1052. M ANCUSCO , C. I., G ATES , A. E. & P UFFER , J. H. 1996. Geochemical and petrologic evidence of Avalonian arc to rift transition from granitoids in southeastern Rhode Island. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 29–63. M ARTI´ NEZ C ATALA´ N , J. R., F ERNA´ NDEZ -S UA´ REZ , J., J ENNER , G. A., B ELOUSOVA , E. & D ´I EZ M ONTES , A. 2004. Provenance constraints from detrital zircon U–Pb ages in the NW Iberian Massif: implications for Palaeozoic plate configuration and Variscan evolution. Journal of the Geological Society, London, 161, 463– 476. M AWER , C. K. & W HITE , J. C. 1987. Sense of displacement on the Cobequid –Chedabucto fault system, Nova Scotia, Canada. Canadian Journal of Earth Sciences, 24, 217–223. M C C AUSLAND , P. J. A. & H ODYCH , J. P. 1998. Palaeomagnetism of the 550 Ma Skinner Cove volcanics of western Newfoundland and the opening of the Iapetus Ocean. Earth and Planetary Science Letters, 163, 15–29. M C K ERROW , W. S., S COTESE , C. R. & B RASIER , M. D. 1992. Early Cambrian continental reconstructions. Journal of the Geological Society, London, 149, 599–606. M C L EOD , M. J., R UITENBERG , A. A. & K ROGH , T. E. 1992. Geology and U– Pb geochronology of the Annidale Group, southern New Brunswick: Lower Ordovician volcanic and sedimentary rocks formed near the southeastern margin of Iapetus Ocean. Atlantic Geology, 28, 181–192.
377
M C M ENAMIN , M. A. S. & M C M ENAMIN , D. L. 1990. The Emergence of Animals: The Cambrian Breakthrough. Columbia University Press, New York. M C N AMARA , A. K., M AC N IOCAILL , C., VAN DER P LUIJM , B. A. & V AN DER V OO , R. 2001. West African proximity of Avalon in the latest Precambrian. Geological Society of America Bulletin, 113, 1161– 1170. M EERT , J. G. & P OWELL , C. M. 2001. Assembly and break-up of Rodinia: introduction to the special volume. Precambrian Research, 110, 1– 8. M ILLER , B. V., S AMSON , S. D. & D’L EMOS , R. S. 1999. Time span of plutonism, fabric development, and cooling in a Neoproterozoic magmatic arc segment: U– Pb age constraints from syntectonic plutons, Sark, Channel Islands, U.K. Tectonophysics, 312, 79–95. M ILLER , B. V., S AMSON , S. D. & D’L EMOS , R. S. 2001. U– Pb geochronological constraints on the timing of plutonism, volcanism, and sedimentation, Jersey, Channel Islands, UK. Journal of the Geological Society, London, 158, 243–252. M ILLER , J. D. & K ENT , D. V. 1988. Paleomagnetism of the Siluro-Devonian Andreas redbeds: evidence of a Devonian supercontinent? Geology, 16, 195–198. M OLINA -G ARZA , R. S., V AN DER V OO , R. & U RRUTIA F UCUGAUCHI , J. 1992. Paleomagnetism of the Chiapas massif, southern Mexico: evidence for rotation of the Maya block and implications for the opening of the Gulf of Mexico. Geological Society of America Bulletin, 104, 1156– 1168. M OORES , E. M. 1998. Ophiolites, the Sierra Nevada, ‘Cordillera,’ and orogeny along the Pacific and Caribbean margins of North and South America. International Geology Review, 40, 40–54. M UELLER , P. A., H EATHERINGTON , A. L., W OODEN , J. L., S HUSTER , R. D., N UTMAN , A. P. & W ILLIAMS , I. S. 1994. Precambrian zircons from the Florida basement; a Gondwanan connection. Geology, 22, 119– 122. M UELLER , P. A., K OZUCH , M., H EATHERINGTON , A. L. ET AL . 1996. Evidence for Mesoproterozoic basement in the Carolina terrane and speculations on its origin. In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian, Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 207–217. M URPHY , J. B. 2000. Tectonic influence on sedimentation along the southern flank of the Late Paleozoic Magdalen Basin in the Canadian Appalachians: geochemical and isotopic constraints on the Horton Group in the St. Marys Basin, Nova Scotia. Geological Society of America Bulletin, 112, 997 –1011. M URPHY , J. B. 2002. Geochemistry of the Neoproterozoic metasedimentary Gamble Brook Formation, Avalon terrane, Nova Scotia: evidence for a rifted arc environment along the west Gondwanan margin of Rodinia. Journal of Geology, 110, 407 –420. M URPHY , J. B. 2003. Late Paleozoic formation and development of the St. Marys Basin, mainland Nova Scotia, Canada: a prolonged record of intra-continental strikeslip deformation during the assembly of Pangaea. In: S TORTI , F., H OLDSWORTH , R. E. & S ALVINI , F. (eds) Intraplate Strike-Slip Deformation Belts.
378
R. D. NANCE ET AL.
Geological Society, London, Special Publications, 210, 185–196. M URPHY , J. B. & K EPPIE , J. D. 2005. The Acadian orogeny in the northern Appalachians. International Geology Review, 47, 663–687. M URPHY , J. B. & N ANCE , R. D. 1989. Model for the evolution of the Avalonian–Cadomian belt. Geology, 17, 735 –738. M URPHY , J. B. & N ANCE , R. D. 2002. Nd –Sm isotopic systematics as tectonic tracers: an example from West Avalonia, Canadian Appalachians. EarthScience Reviews, 59, 77– 100. M URPHY , J. B., C AMERON , K., D OSTAL , J., K EPPIE , J. D. & H YNES , A. J. 1985. Cambrian volcanism in Nova Scotia. Canadian Journal of Earth Sciences, 22, 599 –606. M URPHY , J. B., K EPPIE , J. D., D OSTAL , J. & H YNES , A. J. 1990. The geochemistry and petrology of the late Precambrian Georgeville Group: a volcanic arc– rift succession in the Avalon terrane of Nova Scotia. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 383– 393. M URPHY , J. B., K EPPIE , J. D., D OSTAL , J. & C OUSINS , B. L. 1996a. Repeated late Neoproterozoic–Silurian lower crustal melting beneath the Antigonish Highlands, Nova Scotia: Nd isotopic evidence and tectonic interpretations. In: N ANCE , R. D. & T HOMPSON , M. D. (eds.) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 109–120. M URPHY , J. B., K EPPIE , J. D., D OSTAL , J., W ALDRON , J. W. F. & C UDE , M. P. 1996b. Geochemical and isotopic characteristics of Early Silurian clastic sequences in the Antigonish Highlands, Nova Scotia, Canada: constraints on the accretion of Avalonia in the Appalachian–Caledonide Orogen. Canadian Journal of Earth Sciences, 33, 379–388. M URPHY , J. B., K EPPIE , J. D., D AVIS , D. & K ROGH , T. E. 1997. Regional significance of new U–Pb age data for Neoproterozoic igneous units in Avalonian rocks of northern mainland Nova Scotia, Canada. Geological Magazine, 134, 113–120. M URPHY , J. B., K EPPIE , J. D., D OSTAL , J. & N ANCE , R. D. 1999a. Neoproterozoic– early Paleozoic evolution of Avalonia. In: R AMOS , V. A. & K EPPIE , J. D. (eds) Laurentia–Gondwana Connections Before Pangea. Geological Society of America, Special Papers, 336, 253–266. M URPHY , J. B., VAN S TAAL , C. R. & K EPPIE , J. D. 1999b. Is the mid Paleozoic Acadian Orogeny a plume-modified Laramide-style orogeny? Geology, 27, 653 –656. M URPHY , J. B., S TRACHAN , R. A., N ANCE , R. D., P ARKER , K. D. & F OWLER , M. B. 2000. ProtoAvalonia: a 1.2–1.0 Ga tectonothermal event and constraints for the evolution of Rodinia. Geology, 28, 1071–1074. M URPHY , J. B., P ISAREVSKY , S. A., N ANCE , R. D. & K EPPIE , J. D. 2001. Animated history of Avalonia in Neoproterozoic– Early Paleozoic. Journal of Virtual Explorer, 3, 45– 58.
M URPHY , J. B., E GUILUZ , L. & Z ULAUF , G. 2002a. Cadomian orogens, peri-Gondwanan correlatives and Laurentia–Baltica connections. Tectonophysics, 352, 1–9. M URPHY , J. B., N ANCE , R. D. & K EPPIE , J. D. 2002b. Discussion: West African proximity of Avalon in the latest Precambrian. Geological Society of America Bulletin, 114, 1049–1050. M URPHY , J. B., P ISAREVSKY , S. A., N ANCE , R. D. & K EPPIE , J. D. 2004a. Neoproterozoic– Early Paleozoic evolution of peri-Gondwanan terranes: implications for Laurentia– Gondwana connections. International Journal of Earth Sciences, 93, 659–682. M URPHY , J. B., F ERNA´ NDEZ -S UA´ REZ , J., K EPPIE , J. D. & J EFFRIES , T. E. 2004b. Contiguous rather than discrete Paleozoic histories for the Avalon and Meguma terranes based on detrital zircon data. Geology, 32, 585–588. M URPHY , J. B., F ERNA´ NDEZ -S UA´ REZ , J. & J EFFRIES , T. E. 2004c. Contiguous rather than discrete Paleozoic histories for the Avalon and Meguma terranes based on detrital zircon data. Geology, 32, 585–588. M URPHY , J. B., G UTIE´ RREZ -A LONSO , G., N ANCE , R. D. ET AL . 2006. Origin of the Rheic Ocean: Rifting along a Neoproterozoic suture? Geology, 34, 325–328. N AGY , E. A., S AMSON , S. D. & D’L EMOS , R. S. 2002. U–Pb geochronological constraints on the timing of Brioverian sedimentation and regional deformation in the St. Brieuc region of the Neoproterozoic Cadomian orogen, northern France. Precambrian Research, 116, 1 –17. N ANCE , R. D. & D ALLMEYER , R. D. 1994. Structural and 40 Ar/39Ar mineral age constraints for the tectonothermal evolution of the Green Head Group and Brookville gneiss, southern New Brunswick, Canada: implications for the configuration of the Avalon composite terrane. Geological Journal, 29, 293–322. N ANCE , R. D. & M URPHY , J. B. 1990. Kinematic history of the Bass River Complex, Nova Scotia: Cadomian tectonostratigraphic relations in the Avalon terrane of the Canadian Appalachians. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 395–406. N ANCE , R. D. & M URPHY , J. B. 1994. Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian–Cadomian belt. Geology, 22, 617–620. N ANCE , R. D. & M URPHY , J. B. 1996. Basement isotopic signatures and Neoproterozoic paleogeography of Avalonian–Cadomian and related terranes in the circum-North Atlantic. In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 333–346. N ANCE , R. D. & T HOMPSON , M. E. (eds) 1996. Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304. N ANCE , R. D., M URPHY , J. B., S TRACHAN , R. A., D’L EMOS , R. S. & T AYLOR , G. K. 1991. Late Proterozoic tectonostratigraphic evolution of the Avalonian
THE PERI-GONDWANAN TERRANES and Cadomian terranes. Precambrian Research, 53, 41– 78. N ANCE , R. D., M URPHY , J. B. & K EPPIE , J. D. 2002. Cordilleran model for the evolution of Avalonia. Tectonophysics, 352, 11–32. N AVARRO -S ANTILLAN , D., S OUR -T OVAR , F. & C ENTENO -G ARCIA , E. 2002. Lower Mississippian (Osagean) brachiopods from the Santiago Formation, Oaxaca, Mexico: stratigraphic and tectonic implications. Journal of South American Earth Sciences, 15, 327–336. N EUBAUER , F. 2002. Evolution of late Neoproterozoic to early Paleozoic tectonic elements in Central and Southeast European Alpine mountain belts: review and synthesis. Tectonophysics, 352, 87– 103. N OBLE , S. R., T UCKER , R. D. & P HARAOH , T. C. 1993. Lower Paleozoic 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 OBLET , CH . & L EFORT , J. P. 1990. Sedimentological evidence for a limited separation between Armorica and Gondwana during the Early Ordovician. Geology, 18, 303–306. N OEL , J., S PARIOSU , D. & D ALLMEYER , D. 1988. Paleomagnetism and 40Ar/39Ar ages from the Carolina slate belt, Albemarle, North Carolina: implications for terrane amalgamation with North America. Geology, 16, 64– 68. O’B RIEN , B., O’B RIEN , S. & D UNNING , G. 1991. Silurian cover, Late Precambrian–Early Ordovician basement, and the chronology of Silurian orogenesis in the Hermitage Flexure (Newfoundland Appalachians). American Journal of Science, 291, 760–799. O’B RIEN , S. J., W ARDLE , R. J. & K ING , A. F. 1983. The Avalon zone: a Pan-African terrane in the Appalachian orogen of Canada. Geological Journal, 18, 195–222. O’B RIEN , S. J., O’B RIEN , B. H., D UNNING , G. R. & T UCKER , R. D. 1996. Late Neoproterozoic Avalonian and related peri-Gondwanan rocks of the Newfoundland Appalachians. In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 9–28. O’B RIEN , S. J., D UNNING , G. R., D UBE´ , B., O’D RISCOLL , C. F., S PARKES , B., I SRAEL , S. & K ETCHUM , J. 2001. New insights into the Neoproterozoic geology of the central Avalon Peninsula (parts of NTS Map Areas 1N/6, 1N/7 and 1N/3), eastern Newfoundland. In: Current Research (2001). Newfoundland Department of Mines and Energy, Geological Survey, Report, 2001-1, 169– 189. O FFIELD , T., K UNK , M. & K OEPPEN , R. 1995. Style and age of deformation, Carolina slate belt, central North Carolina. Southeastern Geology, 35, 59–77. O RTEGA -G UTIE´ RREZ , F., R UIZ , J. & C ENTENO G ARCI´ A , E. 1995. Oaxaquia, a Proterozoic microcontinent accreted to North America during the late Paleozoic. Geology, 23, 1127–1130. O RTEGA -O BREGO´ N , C., K EPPIE , J. D., S OLARI , L. A. ET AL . 2003. Geochronology and geochemistry of the 917 Ma, calc-alkaline Etla granitoid pluton
379
(Oaxaca, southern Mexico): evidence of postGrenvillian subduction along the northern margin of Amazonia. International Geology Review, 45, 596– 622. P ATCHETT , P. J., G ALE , N. H., G OODWIN , R. & H UMM , M. J. 1980. Rb–Sr whole-rock isochron ages of late Precambrian to Cambrian igneous rocks from southern Britain. Journal of the Geological Society, London, 137, 649–656. P AULEY , J. C. 1990. The Longmyndian Supergroup and related Precambrian sediments of England and Wales. In: S TRACHAN , R. A. & T AYLOR , G. K. (eds) Avalonian and Cadomian Geology of the North Atlantic. Blackie, Glasgow, 5– 27. P EGRAM , W. J. 1990. Development of continental lithospheric mantle as reflected in the chemistry of the Mesozoic Appalachian tholeiites, USA. Earth and Planetary Science Letters, 97, 316–331. P E -P IPER , G. & J ANSA , L. F. 1999. Pre-Mesozoic basement rocks offshore Nova Scotia, Canada: New constraints on the accretion history of the Meguma terrane. Geological Society of America Bulletin, 111, 1773– 1791. P E -P IPER , G. & M URPHY , J. B. 1989. Geochemistry and tectonic setting of the late Precambrian Folly River Formation, Cobequid Highlands, Avalon Terrane, Nova Scotia: a continental rift within a volcanic-arc environment. Atlantic Geology, 25, 143– 151. P E -P IPER , G. & P IPER , D. J. W. 1989. The Hadrynian Jeffers Group, Cobequid Highlands, Avalon Zone of Nova Scotia: a back-arc volcanic complex. Geological Society of America Bulletin, 101, 364–376. P HARAOH , T. C. & C ARNEY , J. N. 2000. Introduction to the Precambrian rocks of England and Wales. In: C ARNEY , J. N., H ORA , J.M., P HARAOH , T. C., G IBBONS , W., W ILSON , D., B ARKLEY , W. J. & B EVINS , R. E. (eds) The Precambrian Rocks of England and Wales. Geological Conservation Review Series, 20, 1 –17. P IMENTEL , M. M. & F UCK , R. A. 1992. Neoproterozoic crustal accretion in central Brazil. Geology, 20, 375– 379. P INDELL , J. L., B ARRETT , S. F. & C ASE , J. E. 1990. Geologic evolution of the Caribbean region; a plate tectonic perspective. In: D ENGO , G. & C ASE , J. E. (eds) Caribbean Region. Decade of North American Geology, Vol. H. Geological Society of America, Boulder, CO, 405– 432. P OTREL , A., P EUCAT , 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. P OTTS , S., VAN DER P LUIJM , B. & V AN DER V OO , R. 1993. Discordant Silurian paleolatitudes for central Newfoundland: new paleomagnetic evidence from the Springdale Group. Earth and Planetary Science Letters, 120, 1 –12. P OWELL , C. Mc., L I , Z. X., M C E LHINNEY , M. W., M EERT , J. G. & P ARK , J. K. 1993. Paleomagnetic constraints on the timing of the Neoproterozoic breakup of Rodinia and the Cambrian formation of Gondwana. Geology, 21, 889– 892.
380
R. D. NANCE ET AL.
P RATT , B. R. & W ALDRON , J. W. F. 1991. A Middle Cambrian trilobite faunule from the Meguma Group of Nova Scotia. Canadian Journal of Earth Sciences, 28, 1843–1853. P RIGMORE , J. K., B UTLER , A. J. & W OODCOCK , N. H. 1997. Rifting during separation of Eastern Avalonia from Gondwana: evidence from subsidence analysis. Geology, 25, 203–206. Q UESADA , C. 1990. Precambrian terranes in the Iberian Variscan foldbelt. In: S TRACHAN , R. A. & T AYLOR , G. K. (eds) Avalonian and Cadomian Geology of the North Atlantic. Blackie, Glasgow, 109–133. Q UESADA , C. 1991. Precambrian successions in SW Iberia: their relationship to ‘Cadomian’ orogenic events. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , C. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 353– 362. Q UESADA , C. 1997. Evolucio´n geodina´mica de la zona Ossa–Morena durante el ciclo Cadomiense. In: A RUJO , A. & P EREIRA , F. (eds) Livro de Homenagem ao Prof. Francisco Goncalves: Estudio Sobre a Geologia de Zona de Ossa– Morena (Macico Ibe´rico). Instituto Geolo´gico y Minero de Espan˜a, Madrid, 205–230. Q UESADA , C. 2006. Introduction: The Ossa–Morena Zone—from Neoproterozoic arc through Early Palaeozoic rifting to late Palaeozoic orogeny. In: P EREIRA , M. F. & Q UESADA , C. (eds) Ediacaran to Vise´an Crustal Growth Processes in the Ossa–Morena Zone (SW Iberia). The International Geoscience Programme IGCP 497—The Rheic Ocean: Its Origin, Evolution and Correlatives, E´vora Meeting 2006: Conference Abstracts and Field Trip Guide. Instituto Geolo´gico y Minero de Espan˜a, Madrid, 27– 48. Q UESADA , C. & D ALLMEYER , R. D. 1994. Tectonothermal evolution of the Badajoz–Co´rdoba shear zone (SW Iberia): characteristics and 40Ar/39Ar mineral age constraints. Tectonophysics, 231, 195– 213. Q UESADA , C., B ELLIDO , F., D ALLMEYER , R. D. ET AL . 1991. Terranes within the Iberian Massif: correlations with West African sequences. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, New York, 267– 294. R AMOS , V. A. & A LEMAN , A. 2000. Tectonic evolution of the Andes. In: C ORDANI , U. G., M ILANI , E. J., T HOMAZ F ILO , A. & C AMPOS , D. A. (eds) Tectonic Evolution of South America, 31st International Geological Congress, Rio de Janeiro, 635–685. R OBARDET , M. 2003. The Armorica ‘microplate’; fact or fiction? Critical review of the concept and contradictory palaeobiogeographical data. Palaeogeography, Palaeoclimatology, Palaeoecology, 195, 125– 148. R OBERTS , D. 2003. The Scandinavian Caledonides: Event chronology, palaeogeographic settings and likely modern analogues. Tectonophysics, 365, 283–299. R OBINSON , P., T UCKER , R. D., B RADLEY , D., B ERRYL , V. H. N. & O SBERG , P. H. 1998. Paleozoic orogens in New England, USA. Geologiska Fo¨reningens Stockholm Forhandlingar, 120, 119–148. R OBISON , R. & P ANTOJA -A LOR , J. 1968. Tremadocian trilobites from Nochixtlan region, Oaxaca, Mexico. Journal of Paleontology, 42, 767–800.
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, New York, 31–61. R ODRI´ GUEZ -A LONSO , M. D., P EINADO , M., L O´ PEZ P LAZA , M., F RANCO , P., C ARNICERO , A. & G ONZALO , J.C. 2004. Neoproterozoic–Cambrian synsedimentary magmatism in the Central Iberian Zone (Spain): geology, petrology and geodynamic significance. International Journal of Earth Sciences, 93, 897–920. R OGERS , N., VAN S TAAL , C., M C N ICOLL , V., P OLLOCK , J., Z AGOREVSKI , A. & W HALEN , J. 2006. Neoproterozoic and Cambrian arc magmatism along the eastern margin of the Victoria Lake Supergroup: A remnant of Ganderian basement in central Newfoundland? Precambrian Research, 147, 320–341. R UIZ , J., T OSDAL , R. M., R ESTREPO , P. A. & M URILLO M UN˜ ETO´ N , G. 1999. Pb isotope evidence for Colombia–southern Mexico connections in the Proterozoic. In: R AMOS , V. A. & K EPPIE , J. D. (eds) Laurentia–Gondwana Connections Before Pangea. Geological Society of America, Special Papers, 336, 183–197. S ADOWSKI , G. R. & B ETTENCOURT , J. S. 1996. Mesoproterozoic tectonic correlations between eastern Laurentia and the western border of the Amazon craton. Precambrian Research, 76, 213–227. S AMSON , S., P ALMER , A., R OBISON , R. & S ECOR , D. T. 1990. Biogeographical significance of Cambrian trilobites from the Carolina slate belt. Geological Society of America Bulletin, 102, 1459–1470. S AMSON , S. D. & D’L EMOS , R. S. 1998. U– Pb geochemistry and Sm– Nd isotopic composition of Proterozoic gneisses, Channel Islands, U.K. Journal of the Geological Society, London, 155, 609 –618. S AMSON , S. D. & D’L EMOS , R. S. 1999. A precise late Neoproterozoic U–Pb zircon age for the syntectonic Perelle quartz diorite, Guernsey, Channel Islands, UK. Journal of the Geological Society, London, 156, 47–54. S AMSON , S. D., H IBBARD , J. P. & W ORTMAN , G. L. 1995. Nd isotopic evidence for juvenile crust in the Carolina terrane, southern Appalachians. Contributions to Mineralogy and Petrology, 121, 171–184. S AMSON , S. D., S ECOR , D. T. & S TERN , R. 1999. Provenance and paleogeography of Neoproterozoic circum-Atlantic arc-terranes: constraints from U– Pb ages of detrital zircons. Geological Society of America, Abstracts with Programs, 31, A-429. S AMSON , S. D., B ARR , S. M. & W HITE , C. E. 2000. Nd isotopic characteristics of terranes within the Avalon Zone, southern New Brunswick. Canadian Journal of Earth Sciences, 37, 1039–1052. S AMSON , S. D., S ECOR , D. T. & H AMILTON , M. A. 2001. Wandering Carolina: tracking exotic terrnaes with detrital zircons. Geological Society of America, Abstracts with Programs, 33, A-263. S AMSON , S. D., D’L EMOS , R. S., B LICHERT -T OFT , J. & V ERVOORT , J. D. 2003. U– Pb geochronology and Hf–Nd isotope compositions of the oldest Neoproterozoic crust within the Cadomian Orogen: new
THE PERI-GONDWANAN TERRANES evidence for a unique juvenile terrane. Earth and Planetary Science Letters, 208, 165 –180. S AMSON , S. D., D’L EMOS , R. S., M ILLER , B. V. & H AMILTON , M. A. 2005. Neoproterozoic palaeogeography of the Cadomia and Avalon terranes: constraints from detrital zircon U–Pb ages. Journal of the Geological Society, London, 162, 65–71. S A´ NCHEZ -G ARCI´ A , T., B ELLIDO , F. & Q UESADA , C. 2003. Geodynamic setting and geochemical signatures of Cambrian– Ordovician rift-related igneous rocks (Ossa– Morena Zone, SW Iberia). Tectonophysics, 365, 233– 255. S A´ NCHEZ M ARTI´ NEZ , S., J EFFRIES , T., A RENAS , R., F ERNA´ NDEZ -S UA´ REZ , J. & G ARCI´ A -S A´ NCHEZ , R. 2006. A pre-Rodinian ophiolite involved in the Variscan suture of Galicia (Cabo Ortegal Complex, NW Spain). Journal of the Geological Society, London, 163, 737–740. S ANTOS , J. O. S, H ARTMANN , L. A., G AUDETTE , H. E., G ROVES , D. I., M C N AUGHTON , N. J. & F LETCHER , I. R. 2000. A new understanding of the provinces of the Amazon Craton based on integration of field mapping and U–Pb and Sm–Nd geochronology. Gondwana Research, 3, 453–488. S CHENK , P. E. 1995. Annapolis Belt. In: W ILLIAMS , H. (ed.) Geology of the Appalachian– Caledonian Orogen in Canada and Greenland. The Geology of North America, F-1. Geological Society of America, Bouldes, co, 367 –383. S CHENK , P. E. 1997. Sequence stratigraphy and provenance on Gondwana’s margin; the Meguma Zone (Cambrian to Devonian) of Nova Scotia, Canada. Geological Society of America Bulletin, 109, 395–409. S ECOR , D., S AMSON , S. L., S NOKE , A. W. & P ALMER , A. R. 1983. Confirmation of the Carolina slate belt as an exotic terrane. Science, 221, 649– 651. S EGUIN , M., R AO , K. & D EUTSCH , E. 1987. Paleomagnetism and rock magnetism of Early Silurian Dunn Point volcanics, Avalon Zone, Nova Scotia. Physics of the Earth and Planetary Interiors, 46, 369–380. S HERGOLD , J. H. 1975. Late Cambrian and early Ordovician trilobites from the Burke River structural belt, western Queensland, Australia. Australian Geological Survey, Bureau of Mineral Resources Bulletin, 153. S HERVAIS , J. W., S HELLEY , S. A. & S ECOR , D. T., J R 1996. Geochemistry of volcanic rocks of the Carolina and Augusta terranes in central South Carolina: An exotic rifted arc? In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 219 –236. S HERVAIS , J., D ENNIS , A., M C G EE , J. & S ECOR , D. 2003. Deep in the heart of Dixie: pre-Alleghanian eclogite and HP granulite metamorphism in the Carolina terrane, South Carolina, USA. Journal of Metamorphic Geology, 21, 65–80. S INTUBIN , M., D EBACKER , T. N. & V ERNIERS , J. 2002. The tectonometamorphic history of the Brabant Massif (Belgium); the state of the art. Aardkundige Mededelingen, 12. 69– 72. S MITH , G. W. & S OCCI , A. D. 1990. Late Precambrian sedimentary geology of the Boston Basin. In: S OCCI , A. D., S KEHAN , J. W. & S MITH , G. W. (eds)
381
Geology of the Composite Avalon Terrane of Southern New England. Geological Society of America, Special Papers, 245, 75–84. S OLARI , L. A., K EPPIE , J. D., O RTEGA -G UTIE´ RREZ , F., C AMERON , K. L., L OPEZ , R. & H AMES , W. E. 2003. 990 and 1100 Ma Grenvillian tectonothermal events in the northern Oaxacan Complex, southern Mexico: roots of an orogen. Tectonophysics, 365, 257– 282. S OPER , N. J. & W OODCOCK , N. H. 1990. Silurian collision and sediment dispersal patterns in southern Britain. Geological Magazine, 127, 527– 542. S OUR -T OVAR , F., Q UIROZ -B ARROSO , S. A. & N AVARRO -S ANTILLAN , D. 1996. Carboniferous invertebrates from Oaxaca, southern Mexico: midcontinent paleogeographic extension. Geological Society of America, Abstracts with Programs, 28, A365. S TEINER , M. B. & W ALKER , J. D. 1996. Late Silurian plutons in Yucatan. Journal of Geophysical Research, 101, 17727– 17735. S TERN , R. J. 1994. Arc assembly and continental collision in the Neoproterozoic East African Orogen: Implications for the consolidation of Gondwanaland. Annual Review of Earth and Planetary Sciences, 22, 319– 351. S TEWART , D., T UCKER , R. & W EST , D. 1995. Genesis of Silurian composite terrane in northern Penobscot Bay. In: H USSEY , A. & J OHNSTON , R. (eds) Guidebook for Field Trips in Southern Maine and Adjacent New Hampshire. New England Intercollegiate Geological Conference, 87th Annual Meeting, Bowdoin College, Brunswick, Maine, A3-1–A3-21. S TEWART , D., T UCKER , R., A YUSO , R. & L UX , D. 2001. Minimum age for the Neoproterozozoic Seven Hundred Acre Island Formation and the tectonic setting of the Islesboro Formation, Islesboro block, Maine. Atlantic Geology, 37, 41–59. S TEWART , D. B. & T UCKER , R. D. 1998. Geology of the northern Penobscot Bay, Maine. US Geological Survey, Miscellaneous Investigations Series Map, I-2551. S TEWART , J. H., B LODGETT , R. B., B OUCOT , A. J., C ARTER , J. L. & L O´ PEZ , R. 1999. Exotic Paleozoic strata of Gondwanan provenance near Ciudad Victoria, Tamaulipas, Me´xico. In: R AMOS , V. A. & K EPPIE , J. D. (eds) Laurentia–Gondwana Connections Before Pangea. Geological Society of America, Special Papers, 336, 227–252. S TRACHAN , R. A. & T AYLOR , G. K. (eds) 1990. Avalonian and Cadomian Geology of the North Atlantic. Blackie, Glasgow. S TRACHAN , R. A., T RELOAR , P. J., B ROWN , M. & D’L EMOS , R. S. 1989. Cadomian terrane tectonics and magmatism in the Armorican Massif. Journal of the Geological Society, London, 146, 423– 426. S TRACHAN , R. A., R OACH , R. A. & T RELOAR , P. J. 1990. Cadomian terranes in the North Armorican Massif, France. In: S TRACHAN , R. A. & T AYLOR , G. K. (eds) Avalonian and Cadomian Geology of the North Atlantic. Blackie, Glasgow, 65–92. S TRACHAN , R. A., D’L EMOS , R. S. & D ALLMEYER , R. D. 1996a. Late Precambrian evolution of an
382
R. D. NANCE ET AL.
active plate margin: North Armorican Massif, France. In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 319–332. S TRACHAN , R. A., N ANCE , R. D., D ALLMEYER , R. D., D’L EMOS , R. S., M URPHY , J. B. & W ATTS , G. R. 1996b. Late Precambrian tectonothermal evolution of the Malverns Complex. Journal of the Geological Society, London, 153, 589–600. S TRACHAN , R. A., C OLLINS , A. S., B UCHAN , C., N ANCE , R. D., M URPHY , J. B. & D’L EMOS , R. S. 2007. Terrane analysis along a Neoproterozoic active margin of Gondwana: insights from U– Pb zircon geochronology. Journal of the Geological Society, London, 164, 57– 60. S TRONG , D. F., O’B RIEN , S. J., T AYLOR , S. W., S TRONG , P. G. & W ILTON , D. H. 1978. Aborted Proterozoic rifting in eastern Newfoundland. Canadian Journal of Earth Sciences, 15, 117– 131. S UTTER , J., M ILTON , D. & K UNK , M. 1983. 40Ar/39Ar age spectrum dating of gabbro plutons and surrounding rocks in the Charlotte belt of North Carolina. Geological Society of America, Abstracts with Programs, 15, 110. S WINDEN , H. S. & H UNT , P. A. 1991. A U–Pb zircon age from the Connaigre Bay Group, southwestern Avalon Zone, Newfoundland: Implications for regional correlations and metallogenesis. In: Radiometric Age and Isotopic Studies, Report 4. Geological Survey of Canada Paper, 90– 2, 3–10. S YMONS , D. T. A. & C HIASSON , A. D. 1991. Paleomagnetism of the Callander Complex and the Cambrian apparent polar wander path for North America. Canadian Journal of Earth Sciences, 28, 355– 363. T ANCZYK , E. I., L APOINTE , P., M ORRIS , W. A. & S CHMIDT , P. W. 1987. A paleomagnetic study of the layered mafic intrusion at Sept-Iles, Quebec. Canadian Journal of Earth Sciences, 24, 1431–1438. T ASSINARI , C. C. G. & M ACAMBIRA , M. J. B. 1999. Geochronological provinces of the Amazon Craton. Episodes, 22, 174– 182. T HEOKRITOFF , G. 1979. Early Cambrian provincialism and biogeographic boundaries in the North Atlantic region. Lethaia, 12, 281–295. T HOMPSON , M. D. & B OWRING , S. A. 2000. Age of the Squantum ‘Tillite’, Boston Basin, Massachusetts: U–Pb zircon constraints on terminal Neoproterozoic glaciation. American Journal of Science, 300, 630–655. T HOMPSON , M. D., H ERMES , O. D., B OWRING , S. A., I SACHSEN , C. E., B ESANCON , J. R. & K ELLY , K. L. 1996. Tectonostratigraphic implications of Late Proterozoic U–Pb zircon ages in the Avalon Zone of southeastern New England. In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 179– 191. T HOMPSON , M. D., G RUNOW , A. M. & R AMEZANI , J. 2006. Avalon off North Africa at 595 Ma and other implications of U –Pb geochronology and paleomagnetism in Neoproterozoic rocks around Boston, MA.
Geological Society of America, Abstracts with Programs, 38, 398–399. T HOROGOOD , E. J. 1990. Provenance of the pre-Devonian sediments of England and Wales: Sm-Nd isotopic evidence. Journal of the Geological Society, London, 147, 591–594. T OHVER , E., VAN DER P LUIJM , B. A., V AN DER V OO , R., R IZOTTO , G. & S CANDOLARA , J. E. 2002. Paleogeography of the Amazon craton at 1.2 Ga: Early Grenvillian collision with the Llano segment of Laurentia. Earth and Planetary Science Letters, 199, 185–200. T RENCH , A. & T ORSVIK , T. H. 1992. The closure of the Iapetus Ocean and Tornquist Sea: New paleomagnetic constraints. Journal of the Geological Society, London, 149, 867– 870. T UCKER , R. D. & P HAROAH , T. C. 1991. U– Pb zircon ages of late Precambrian rocks in southern Britain. Journal of the Geological Society, London, 148, 435–443. U GIDOS , J. M., V ALLADARES , M. I., B ARBA , P. & E LLAM , R. M. 2003. The Upper Neoproterozoic– Lower Cambrian of the Central Iberian Zone, Spain: Chemical and isotopic (Sm–Nd) evidence that the sedimentary succession records an inverted stratigraphy of its source. Geochimica et Cosmochimica Acta, 67, 2615– 2629. U STAO¨ MER , P. A. 1999. Pre-Early Ordovician Cadomian arc-type granitoids, the Bolu Massif, West Pontides, northern Turkey: geochemical evidence. International Journal of Earth Sciences, 88, 2– 12. V ALVERDE -V AQUERO , P. & D UNNING , G. R. 2000. New U–Pb ages for Early Ordovician magmatism in Central Spain. Journal of the Geological Society, London, 157, 15–26. V AN DER V ELDEN , A. J., VAN S TAAL , C. R. & C OOK , F. A. 2004. Crustal structure, fossil subduction and the tectonic evolution of the Newfoundland Appalachians: Evidence from a reprocessed seismic reflection survey. Geological Society of America Bulletin, 116, 1485– 1498. V AN DER V OO , R. 1988. Paleozoic paleogeography of North America, Gondwana, and intervening displaced terranes: comparisons of paleomagnetism with paleoclimatology and biogeographical patterns. Geological Society of America Bulletin, 100, 311 –324. VAN S TAAL , C. 1994. Brunswick subduction complex in the Canadian Appalachians: record of the Late Ordovician to Late Silurian collision between Laurentia and the Gander margin of Avalon. Tectonics, 13, 946– 962. VAN S TAAL , C. R., S ULLIVAN , R. W. & W HALEN , J. B. 1996. Provenance and tectonic history of the Gander Margin in the Caledonian/Appalachian Orogen: implications for the origin and assembly of Avalonia. In: N ANCE , R. D. & T HOMPSON , M. A. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 347– 367. VAN S TAAL , C. R., D EWEY , J. F., M AC N IOCAILL , C. & M C K ERROW , S. 1998. The Cambrian–Silurian tectonic evolution of the northern Appalachians: history of a complex, southwest Pacific-type segment of Iapetus. In: B LUNDELL , D. J. & S COTT , A. C. (eds)
THE PERI-GONDWANAN TERRANES Lyell: the Past is the Key to the Present. Geological Society, London, Special Publications, 143, 199– 242. V ICK , H., C HANNELL , J. & O PDYKE , N. 1987. Ordovician docking of the Carolina slate belt: paleomagnetic data. Tectonics, 6, 573– 583. V IDAL , G., P ALACIOS , T., G A´ MEZ -V INTANED , J. A., D I´ EZ B ALDA , M. A. & G RANT , S. W. F. 1994. Neoproterozoic –early Cambrian geology and palaeontology of Iberia. Geological Magazine, 131, 729 –765. VON R AUMER , J. F., S TAMPFLI , G. M., B OREL , G. & B USSY , F. 2002. Organization of pre-Variscan basement areas at the north-Gondwanan margin. International Journal of Earth Sciences, 91, 35–52. W ALDRON , J. W. F. 1992. The Goldenville– Halifax transition, Mahone Bay, Nova Scotia: relative sea-level rise in the Meguma source terrane. Canadian Journal of Earth Sciences, 29, 1091–1105. W ALDRON , J. W. F., M URPHY , J. B., M ELCHIN , M. J. & D AVIS , G. 1996. Silurian tectonics of western Avalonia: strain-corrected subsidence history of the Arisaig Group, Nova Scotia. Journal of Geology, 104, 677– 694. W ATTS , M. J. & W ILLIAMS , G. D. 1979. Fault rocks as indicators of progressive shear deformation in the Guingamp region, Brittany. Journal of Structural Geology, 1, 323–332. W EBER , B. & K O¨ HLER , H. 1999. Sm/Nd, Rb/Sr, and U–Pb geochronology of a Grenville terrane in southern Mexico: origin and geologic history of the Guichicovi complex. Precambrian Research, 96, 245–262. W EBSTER , T. L., B ARR , S. M. & M URPHY , J. B. 1998. Anatomy of a terrane boundary; an integrated structural, geographic information system, and remote sensing study of the late Paleozoic Avalon– Meguma terrane boundary, mainland Nova Scotia, Canada. Canadian Journal of Earth Sciences, 35, 787– 801. W EIL , A. B., V AN DER V OO , R., M AC N IOCAILL , C. & M EERT , J. G. 1998. The Proterozoic supercontinent Rodinia: paleomagnetically derived reconstruction for 1100 to 800 Ma. Earth and Planetary Science Letters, 154, 13–24. W EIL , A. B., V AN DER V OO , R. & VAN DER P LUIJM , B. 2001. Oroclinal bending and evidence against the Pangea megashear: the Cantabrian–Asturias arc (northern Spain). Geology, 29, 901– 904. W ENDT , J. I., K RO¨ NER , A., F IALA , J. & T ODT , W. 1993. Evidence from zircon dating for existence of approximately 2.1 Ga old crystalline basement in southern Bohemia, Czech Republic. Geologische Rundschau, 82, 42– 50. W HALEN , J. B., J ENNER , G. A., C URRIE , K. L., B ARR , S. M., L ONGSTAFFE , F. J. & H EGNER , E. 1994. Geochemical and isotopic characteristics of granitoids of the Avalon Zone, southern New Brunswick: possible
383
evidence of repeated delamination events. Journal of Geology, 102, 269–282. W HALEN , J. B., F YFFE , L. R., L ONGSTAFFE , F. J. & J ENNER , G. 1996. The position and nature of the Gander–Avalon boundary, southern New Brunswick, based on geochemical and isotopic data from granitoid rocks. Canadian Journal of Earth Sciences, 33, 129– 139. W HALEN , J. B., M C N ICOLL , V. J., VAN S TAAL , C. R., L ISSENBERG , C. J., L ONGSTAFFE , F .J., J ENNER , G. A. & VAN B REEMAN , O. 2006. Spatial, temporal and geochemical characteristics of Silurian collision-zone magmatism, Newfoundland Appalachians: An example of a rapidly evolving magmatic system related to slab break-off. Lithos, 89, 377– 404. W HITE , C., B ARR , S., B EVIER , M. & K AMO , S. 1994. A revised interpretation of Cambrian and Ordovician rocks in the Bourinot belt of central Cape Breton Island, Nova Scotia. Atlantic Geology, 30, 123– 142. W HITE , C. E., B ARR , S. M., M ILLER , B. V. & H AMILTON , M. A. 2002. Granitoid plutons of the Brookville terrane, southern New Brunswick: petrology, age, and tectonic setting. Atlantic Geology, 38, 53–74. W HITE , C. E., B ARR , S. M. & T OOLE , R. M. 2006. New insights on the origin of the Meguma Group in southwestern Nova Scotia, Canada. Geological Society of America, Abstracts with Programs, 38, 9. W INGATE , M. T. D. & G IDDINGS , J. W. 2000. Age and paleomagnetism of the Mundine Well dyke system, Western Australia: implications for Australia– Laurentia connection at 755 Ma. Precambrian Research, 100, 335–357. W ORTMAN , G. L., S AMSON , S. D. & H IBBARD , S. D. 2000. Precise U–Pb zircon constraints on the earliest magmatic history of the Carolina terrane. Journal of Geology, 108, 321–338. W RIGHT , J. & S EIDERS , V. 1980. Age of zircon from volcanic rocks of the central North Carolina Piedmont and tectonic implications for the Carolina volcanic slate belt. Geological Society of America Bulletin, 91, 287– 294. Z ULAUF , G., D O¨ RR , W., F IALA , J. & V EJNAR , Z. 1997. Late Cadomian crustal tilting and Cambrian transtension in the Tepla´ –Barrandian unit (Bohemian Massif, Central European Variscides). Geologische Rundschau, 86, 571– 584. Z ULAUF , G., S CHITTER , F., R IEGLER , G., F INGER , F., F IALA , J. & V EJNAR , Z. 1999. Age constraints on the Cadomian evolution of the Tepla´ – Barrandian unit (Bohemian Massif) through electron microprobe dating of metamorphic monazite. Zeitschrift der Deutschen Geologischen Gesellschaft, 150, 627–639.
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 M. F. PEREIRA1, M. CHICHORRO2, I. S. WILLIAMS3 & J. B. SILVA3 Departamento de Geocieˆncias, Centro de Geofı´sica de E´vora, Universidade de E´vora, Apartado 94, 7002-554 E´vora, Portugal (e-mail:
[email protected])
1
2
Research School of Earth Sciences, The Australian National University, Canberra, A.C.T. 0200, Australia
3
Departamento de Geologia, Faculdade de Cieˆncias, Universidade de Lisboa, Portugal Abstract: Sensitive high-resolution ion microprobe U– Th–Pb age determinations on detrital and inherited zircon from the E´vora Massif (SW Iberian Massif, Portugal) provide direct evidence for the provenance of the Ossa–Morena Ediacaran basins (Se´rie Negra) and a palaeogeographical link with the West African craton. Three samples of the Se´rie Negra paragneisses contain large components of Cryogenian and Ediacaran (c. 700– 540 Ma) detrital zircon, but have a marked lack of zircon of Mesoproterozoic (c. 1.8– 0.9 Ga) age. Older inherited zircons are of Palaeoproterozoic (c. 2.4– 1.8 Ga) and Archaean (c. 3.5–2.5 Ga) age. The same age pattern is also found in the Arraiolos biotite granite, which was formed by partial melting of the Se´rie Negra and overlying Cambrian rocks. These results are consistent with substantial denudation of a continental region that supplied sediments to the Ediacaran Ossa–Morena basins during the final stages of the Cadomian– Avalonian orogeny (peri-Gondwanan margin with principal zircon-forming events at c. 575 Ma and c. 615 Ma). Combined with the detrital zircon ages reported for rocks of the same age from Portugal, Spain, Germany and Algeria, our data suggest that the sediment supply to the Ediacaran– Early Palaeozoic siliciclastic sequences preserved in all these periGondwanan regions was similar. The lack of Grenvillian-aged (c. 1.1–0.9 Ga) zircon in the Ossa–Morena and Saxo-Thuringia Ediacaran sediments suggests that the sediment in these peri-Gondwanan basins was derived from the West African craton.
The sensitive high-resolution ion microprobe (SHRIMP) U –Th –Pb geochronology of detrital and inherited zircon grains in sedimentary and metasedimentary rocks is a technique that has proved very useful in constraining palaeogeographical reconstruction models for Neoproterozoic and Early Palaeozoic sedimentary basins of North Africa and Arabia (Williams et al. 2002; Avigad et al. 2003; Kolodner et al. 2006). The spectrum of ages and compositions of detrital and inherited zircons is not only an excellent indicator of sediment provenance, but also helps to identify the protoliths of high-grade rocks. These data provide direct information about igneous and metamorphic events in the continental crust of the source area, especially if the zircon grains are zoned as a result of multiple episodes of growth. A recent example of such a study in Europe is the work of Linnemann et al. (2004), who used SHRIMP data to deduce the provenance of Late Neoproterozoic and Early Palaeozoic sediments of the Saxo-Thuringia Zone (Germany). Those workers demonstrated how the
integrated analysis of SHRIMP zircon U–Pb geochronology, cathodoluminescence (CL) imaging of zircon growth textures, and statistical assessment of SHRIMP data was invaluable in defining the provenance of such peri-Gondwanan sediments. They also found that the principal cratonic source of the sediments (the West African craton where Grenvillian (c. 1.1–0.9 Ga) zircon-forming events are rare or absent) did not change from the Late Neoproterozoic to the Palaeozoic, and therefore that the SaxoThuringia Zone must have remained with North Africa for a long period of geological time. Recent studies in the SW Iberian Massif (Julivert et al. 1972) have produced new results from whole-rock geochemistry, Nd–Sr –Pb isotopic chemistry and radiometric dating of detrital zircons that help characterize the nature and provenance of the Neoproterozoic to Early Cambrian clastic sediments of the Ossa–Morena Zone (e.g. Na¨gler 1990; Beetsma 1995; Ordon˜ez-Casado 1998; Ferna´ndez-Sua´rez et al. 2002; Gutie´rrezAlonso et al. 2003; Pereira et al. 2006a). These
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 385–408. DOI: 10.1144/SP297.18 0305-8719/08/$15.00 # The Geological Society of London 2008.
386
M. F. PEREIRA ET AL.
data contributed to improving the reconstruction of the Cadomian–Avalonian orogen within the periGondwanan realm (e.g. Nance & Murphy 1996; Murphy et al. 2002, 2006; Nance et al. 2002; Gutie´rrez-Alonso et al. 2003; Linnemann et al. 2004; Pereira et al. 2006a, b). The present paper reports SHRIMP U –Th– Pb ages of complex zircon grains extracted from three SW Iberian Massif (Ossa –Morena Zone, Portugal) paragneisses and a biotite granite. The ages of detrital and inherited zircon extracted from the E´vora Massif rocks make it possible to estimate the relative abundances of different crustal components in the source area from which these rocks were ultimately derived. Comparisons with other published data from Europe, North Africa and Arabia make it possible to test hypotheses of global correlations and possible palaeogeographical links between different peri-Gondwanan basins probably located close to the West African craton during the Ediacaran.
Geological setting The Ossa –Morena Zone, located at the southwestern end of the European Variscides (Fig. 1), is part of the Iberian Massif (Julivert 1987; Matte 2001; Pereira & Quesada 2006). The local effects of the Variscan orogeny (c. 370– 280 Ma) were variable deformation and metamorphic reworking of the Neoproterozoic and Early Palaeozoic rocks. The E´vora Massif, in the westernmost part of the Ossa–Morena Zone (Figs 1 and 2) consists of a lithological succession (Fig. 3), with Ediacaran (c. 560–550 Ma) metagreywackes, micaschists, paragneisses and interbedded black metacherts, amphibolites and felsic gneisses (Se´rie Negra; Carvalhosa 1965), an Early–Mid-Cambrian (c. 530–505 Ma) igneous (felsic-dominated)–sedimentary complex with marbles, interbedded felsic and mafic metavolcanic rocks, felsic gneisses and micaschists, and finally a Mid–Late Cambrian–(?)Early Ordovician (c. 505–(?)480 Ma) igneous (mafic-dominated)– sedimentary complex including amphibolites with mica schists, quartzites, metatuffs and calc-silicate rocks. Based on structural, metamorphic and lithological criteria, the dome-like E´vora Massif (Carvalhosa 1983; Quesada & Munha´ 1990) has been divided into three main tectonic units (Pereira et al. 2003, 2007; Figs 2 and 3): the Montemor-o-Novo shear zone and the E´vora medium-grade metamorphic terranes form the hanging walls of a 25– 45 km wide and 75 km long crystalline metamorphic core, the E´vora highgrade metamorphic terranes. The transcurrent movements in this part of the Ossa –Morena Zone
involved significant crustal extension and the development of major ductile shear zones. These orogenparallel movements were responsible for the partial exhumation of a structurally complex assemblage with anatectic granitoids and migmatitic orthoand paragneisses (upper amphibolite to transitional amphibolite –granulite facies). The huge volume of melt produced by uplift and decompression at this crustal level is indicated by numerous granites and granodiorites intruded into the high-grade metamorphic rocks. Sufficient magma was accumulated to cause its emplacement along the boundaries of the E´vora high-grade metamorphic terranes boundaries, and at shallower crustal levels in the hangingwall medium- and low-grade metamorphic rocks (Pereira et al. 2007).
Sample preparation and analytical methods Four samples were collected (locations are shown in Figs 2 and 3): three samples of paragneisses from the Se´rie Negra, BSC-1 (UTM coordinates 29SNC707609), SEC-1 (UTM coordinates 29SNC742677) and CSN-26 (UTM coordinates 29SNC 786 688), and an anatectic biotite granite from Arraiolos interpreted as the product of partial melting of the Se´rie Negra sediments and the Cambrian igneous–sedimentary complexes, ARL-6 (29SNC762664). Zircon grains were extracted at the Departamento de Geocieˆncias (Universidade de Aveiro and Universidade de E´vora) using a procedure designed to minimize trace mineral contamination. Samples were crushed to 300–500 mm in a jaw crusher, then c. 500 g of the grit reduced to powder in a disc mill. The dust fraction was removed with water, then the high-density minerals were extracted in bromoform (2.89 g cm23). Strongly magnetic minerals were removed using a hand-magnet. At the Research School of Earth Sciences (Australian National University, Canberra) the zircon in the non-magnetic heavy mineral fraction was further concentrated using methylene iodide (3.33 g cm23) then a representative selection of grains was hand-picked on the basis of clarity, colour, size and morphology. Turbid grains were rejected. The selected grains were cast in epoxy resin (11 grains from sample BSC-1, 133 from sample SEC-1, 104 from sample CSN-26 and 144 from sample ARL-6), together with zircon standards SL13 (U ¼ 238) and TEMORA (206Pb*/238U ¼ 0.06683). The mount was polished to expose the grain interiors, photographed at high magnification in transmitted and reflected light, then imaged by
OSSA– MORENA EDIACARAN BASINS, IBERIA 387
Fig. 1. Schematic geological map of the European and NW African Variscan basement (modified from Pique´ et al. 1994; Matte 2001) and of the Ossa–Morena Zone with location of the main exposures of the Ediacaran and Cambrian rocks (modified from Eguiluz et al. 2000; Pereira et al. 2006a).
388
M. F. PEREIRA ET AL.
Fig. 2. Schematic geological map of the E´vora Massif (modified from Carvalhosa et al. 1969; Carvalhosa 1983, 1999; Oliveira 1992; Carvalhosa & Zbyszweski 1994; Pereira et al. 2003, 2004, 2007; Chichorro 2006).
SEM and CL to document the internal growth zoning of the grains. It was then cleaned and coated with high-purity Au in preparation for analysis. Selected areas in the zircon grains were analysed for U, Th and Pb isotopes on the ANU SHRIMP II ion microprobe using a procedure similar to that described by Williams & Claesson (1987). Detrital grains and inherited cores were chosen to have features as representative as possible of the population as a whole and the analyses were positioned to avoid textural complexity, fractures and inclusions.
A 10 kV O2 2 primary beam was focused to either c. 10 or 20 mm diameter, depending on the spatial resolution required. Positive secondary ions were extracted at 10 kV and mass analysed at c. R5000 on a single ETP multiplier by peak stepping through the Zr, Pb, U and Th species of interest. Analytical uncertainties in Tables 1–4 and Figures 7 and 8 are 1s precision estimates. Uncertainties in the mean ages discussed below are 95% confidence limits (ts, where t is the Student’s t multiplier) and, for the mean 206Pb/238U ages,
OSSA– MORENA EDIACARAN BASINS, IBERIA Fig. 3. Schematic geological transverse and stratigraphic column of the E´vora Massif (modified from Carvalhosa 1983, 1999; Oliveira et al. 1991; Carvalhosa & Zbyszweski 1994; Pereira et al. 2003, 2007; Chichorro 2006).
389
Table 1. SHRIMP zircon data from paragneiss BSC-1 Analy- Struc- Pb* U Th sis ture ppm ppm ppm
2.1 1.2 3.2 4.1 5.1 6.1 7.1
LUZO CZC BZC CZC LUZC CZ UZC
1 50 91 35 3 13 82
9 448 225 311 27 98 254
3 358 60 206 14 165 149
Th/ U
0.29 0.80 0.27 0.66 0.54 1.67 0.59
204
Pb ppb
2 4 0 1 0 0 4
204
Pb
+
3.2E-03 1.1E-04 4.7E-06 2.0E-05 2.0E-05 2.0E-05 6.4E-05
f 206Pb
206 Pb
2.1E-03 3.6E-05 3.8E-06 2.0E-05 2.0E-05 2.0E-05 4.4E-05
0.02992 0.00196 0.00007 0.00035 0.00036 0.00036 0.00096
208
Pb
+
206 Pb
– 0.2443 0.0725 0.2060 0.2002 0.5190 0.1902
208
Pb
+
232 Th
– 0.0049 0.0012 0.0044 0.0244 0.0092 0.0026
– 0.03023 0.10653 0.03203 0.03264 0.02940 0.09198
206
Pb
+
238 U
– 0.00066 0.00259 0.00080 0.00417 0.00075 0.00217
0.12976 0.09898 0.39072 0.10299 0.08843 0.09480 0.28394
207
Pb
+
206 Pb
0.00852 0.00062 0.00534 0.00114 0.00279 0.00145 0.00436
0.06846 0.05940 0.13251 0.06122 0.06417 0.06185 0.13053
Apparent age (Ma) 208 232
+
206 238
0.01329 – – 786.5 0.00103 602.0 12.9 608.5 0.00112 2046.1 47.3 2126.2 0.00104 637.3 15.6 631.9 0.00318 649.3 81.8 546.3 0.00171 585.7 14.8 583.9 0.00110 1778.5 40.2 1611.2
UZC, Unzoned core; LUZC, light unzoned core; CZC, Concentric and sector zoned core; BZC, Banded zoned core; LUZO, Light unzoned overgrowth. *, Radiogenic. f, Correction for common Pb-Fraction of total 206Pb (206PbT) that is common 206Pb (206Pbc). §, Best estimated of the age of the analysed zircon. (see text for explanation.) Uncertainties one standard error.
+
207 206
Inf Age
+
+
48.8 882.4 463.5 783.7 3.7 581.7 38.2 609.0 24.8 2131.5 14.9 2132.0 6.7 647.1 36.9 631.5 16.5 747.3 108.5 542.6 8.6 668.8 60.3 582.2 21.9 2105.1 14.9 2105.0
48.2 3.7 15.0 6.7 17.0 8.6 15.0
Table 2. SHRIMP zircon data from paragneiss SEC-1 Analy- Struc- Pb* U Th Th/ sis ture ppm ppm ppm U
3.1 7.1 10.1 12.1 13.1 10.2 14.1 15.1 16.1 16.2 17.1 18.1 19.1 20.1 21.1 21.2 4.3 4.4 7.2 28.1 22.1 23.1 24.1 25.1 26.1 9.2 12.2 27.1 3.2
UZO UZO LUZO LUZO CZC CZC LCZC UZC CZO CZC LUZC CZC BZC CZC UZC UZIO BZC UZIO BZC UZ UZC BZC UZC UZC CZC BZC CZC CZC UZC
49 129 54 253 513 34 5 22 6 4 40 16 79 221 124 208 496 386 34 44 40 290 520 417 9 72 119 49 332 172 5 45 18 28 304 68 6 55 29 95 1021 365 81 238 65 37 411 8 20 173 173 29 287 130 139 190 83 46 479 190 41 272 313 23 241 106 137 312 8 38 90 69 53 401 588 434 551 352 42 432 354 91 938 379 30 56 91
For legend see Table 1.
0.42 0.07 0.27 0.39 0.56 0.78 0.91 0.80 1.65 0.52 0.41 0.22 0.52 0.36 0.27 0.02 1.00 0.45 0.44 0.40 1.15 0.44 0.02 0.77 1.47 0.64 0.82 0.40 1.62
204
Pb ppb
7 6 5 1 1 2 2 1 3 3 1 1 1 1 4 2 0 5 3 2 4 0 3 2 3 3 4 3 1
204
Pb
+
1.9E-04 2.7E-05 1.2E-03 1.7E-03 2.4E-05 1.0E-05 8.0E-05 2.8E-06 5.3E-04 6.5E-05 6.8E-04 5.3E-05 2.7E-04 1.1E-05 6.2E-05 4.7E-05 2.0E-05 1.9E-04 3.2E-05 4.3E-05 1.4E-04 2.4E-05 2.5E-05 8.6E-05 9.0E-05 1.0E-05 1.1E-04 3.5E-05 6.4E-05
f 206Pb
206 Pb
6.0E-05 1.7E-05 4.8E-04 8.4E-04 1.1E-05 5.7E-06 6.7E-05 4.0E-06 2.8E-04 3.2E-05 4.3E-04 2.9E-05 1.9E-04 9.8E-06 1.9E-05 2.8E-05 2.0E-05 7.0E-05 1.9E-05 3.3E-05 9.2E-05 3.4E-05 1.0E-05 5.6E-05 4.6E-05 3.6E-06 8.5E-05 3.2E-05 5.8E-05
0.00282 0.00036 0.01930 0.00717 0.00037 0.00016 0.00100 0.00004 0.00949 0.00114 0.00460 0.00095 0.00471 0.00020 0.00094 0.00087 0.00036 0.00339 0.00040 0.00076 0.00245 0.00042 0.00033 0.00128 0.00160 0.00013 0.00205 0.00063 0.00095
208
Pb
+
206 Pb
0.1238 0.0158 0.0863 – 0.1632 0.2177 0.2474 0.2191 0.4890 0.1591 – 0.0703 0.1660 0.1057 0.0804 – 0.3033 0.1307 0.1174 0.1238 0.3587 0.1347 0.0039 0.2166 0.4603 0.1664 0.2811 0.1235 0.4645
208
Pb
+
232 Th
0.0032 0.0009 0.0184 – 0.0016 0.0018 0.0088 0.0019 0.0148 0.0048 – 0.0019 0.0128 0.0017 0.0016 – 0.0077 0.0039 0.0015 0.0021 0.0052 0.0029 0.0005 0.0034 0.0057 0.0007 0.0045 0.0029 0.0081
0.10446 0.10997 0.06491 – 0.09465 0.10100 0.16248 0.12697 0.02928 0.04311 – 0.03004 0.03548 0.02738 0.09639 – 0.03056 0.02860 0.16787 0.02938 0.03795 0.02780 0.06924 0.10283 0.03154 0.16592 0.02919 0.02909 0.11059
206
Pb
+
238 U
0.00419 0.00621 0.01432 – 0.00163 0.00141 0.00869 0.00145 0.00111 0.00135 – 0.00088 0.00306 0.00047 0.00238 – 0.00088 0.00091 0.00438 0.00056 0.00070 0.00073 0.00945 0.00307 0.00058 0.00201 0.00062 0.00072 0.00458
0.35193 0.46590 0.20311 0.10483 0.32410 0.36089 0.59740 0.46465 0.09859 0.14052 0.10677 0.09581 0.11213 0.09264 0.32943 0.09905 0.10078 0.09896 0.62403 0.09426 0.12181 0.09110 0.42930 0.36472 0.10065 0.63666 0.08520 0.09517 0.38680
207
Pb
+
206 Pb
0.00847 0.00439 0.00924 0.00367 0.00381 0.00325 0.02114 0.00266 0.00195 0.00102 0.00258 0.00072 0.00367 0.00042 0.00398 0.00077 0.00114 0.00087 0.01132 0.00065 0.00111 0.00118 0.00390 0.00748 0.00116 0.00587 0.00098 0.00066 0.01182
0.13151 0.21203 0.10912 0.06629 0.11391 0.12776 0.25084 0.17284 0.05608 0.06835 0.06327 0.05965 0.06303 0.05834 0.12198 0.05959 0.06168 0.05622 0.24662 0.05919 0.06469 0.06050 0.18981 0.13302 0.06069 0.26855 0.05933 0.05934 0.13373
0.00183 0.00207 0.00861 0.00863 0.00077 0.00242 0.00821 0.00080 0.00503 0.00099 0.00637 0.00135 0.00369 0.00049 0.00141 0.00082 0.00182 0.00240 0.00166 0.00108 0.00169 0.00119 0.00410 0.00249 0.00158 0.00066 0.00180 0.00106 0.00230
Apparent age (Ma) 208 232 2008.2 2108.8 1271.2 – 1827.8 1944.9 3043.1 2416.0 583.3 853.0 – 598.2 704.7 546.0 1860.0 – 608.3 570.0 3136.6 585.3 752.9 554.2 1353.1 1978.4 627.7 3102.7 581.6 579.6 2120.1
+
206 238
76.8 113.4 273.7 – 30.1 25.8 151.6 26.1 21.8 26.3 – 17.2 59.9 9.3 43.9 – 17.2 18.0 76.0 10.9 13.6 14.4 179.5 56.3 11.4 34.9 12.3 14.2 83.6
1943.8 2465.6 1192.0 642.7 1809.8 1986.4 3019.4 2460.1 606.2 847.6 653.9 589.8 685.1 571.1 1835.6 608.8 619.0 608.3 3125.9 580.7 741.0 562.1 2302.6 2004.5 618.2 3175.9 527.1 586.0 2108.0
I.d. Inf
+
207 206
+
40.5 19.3 49.7 21.5 18.6 15.4 85.9 11.7 11.5 5.8 15.0 4.3 21.3 2.5 19.3 4.5 6.7 5.1 45.1 3.8 6.4 7.0 17.6 35.5 6.8 23.2 5.8 3.9 55.2
2118.2 2921.1 1784.8 815.5 1862.8 2067.4 3190.1 2585.3 455.7 879.2 717.4 591.1 709.2 542.5 1985.4 588.6 663.2 461.1 3163.2 573.9 764.3 621.6 2740.5 2138.1 628.2 3297.6 579.3 579.7 2147.5
24.5 15.9 151.3 298.7 12.3 33.8 52.8 7.7 212.3 30.4 229.8 49.9 129.7 18.4 20.8 30.3 64.4 97.5 10.7 40.3 56.2 43.1 36.0 33.1 57.3 3.8 67.4 39.3 30.4
+
2118.0 24.0 2921.0 16.0 1785.0 150.0 638.7 21.0 1863.0 12.0 2067.0 33.0 3190.0 53.0 2585.0 8.0 609.1 11.0 846.4 5.8 652.4 15.0 589.8 4.3 684.4 21.0 571.2 2.5 1985.0 21.0 609.2 4.6 618.0 6.8 610.3 5.1 3163.0 11.0 580.8 3.9 740.3 6.3 560.9 7.0 2740.0 36.0 2138.0 33.0 618.0 6.9 3298.0 3.8 526.2 5.8 586.1 3.9 2148.0 30.0
Table 3. SHRIMP zircon data from paragneiss CSN-26 Analy- Struc- Pb* U Th Th/ sis ture ppm ppm ppm U
4.1 4.2 9.1 10.1 11.1 5.2 12.1 13.1 7.2 15.1 16.1 2.2 14.1 3.2 17.1 18.1 19.1 20.1
UZO UZC CZC CZC UZC CZC LCZC CZC UZC SZC CZC CZC BZC CZC CZC LCZC CZC LCZC
218 1455 113 395 497 221 16 132 103 53 486 348 451 1178 886 59 95 90 26 227 236 29 277 159 118 293 157 56 490 303 21 183 162 27 250 223 49 290 327 90 805 540 56 450 495 11 102 73 25 219 227 49 105 37
For legend see Table 1.
0.08 0.45 0.78 0.72 0.75 0.95 1.04 0.57 0.54 0.62 0.88 0.89 1.13 0.67 1.10 0.72 1.04 0.35
204
Pb ppb
39 3 2 3 2 6 1 2 2 2 0 3 5 4 67 2 2 0
204
Pb
+
2.0E-04 9.6E-06 1.9E-04 8.2E-05 6.3E-06 1.4E-04 6.1E-05 1.0E-04 2.0E-05 4.7E-05 1.1E-05 1.7E-04 1.6E-04 5.1E-05 1.7E-03 2.0E-04 1.3E-04 9.0E-06
f 206Pb
206 Pb
2.8E-05 1.3E-05 8.3E-05 3.0E-05 2.6E-06 3.7E-05 3.9E-05 8.6E-05 2.0E-05 2.4E-05 7.1E-05 5.5E-05 4.9E-05 1.7E-05 1.4E-04 3.2E-04 6.9E-05 1.0E-05
0.00322 0.00012 0.00336 0.00145 0.00010 0.00190 0.00109 0.00179 0.00030 0.00083 0.00020 0.00297 0.00285 0.00090 0.02984 0.00351 0.00224 0.00013
208
Pb
+
206 Pb
0.0227 0.1165 0.2425 0.2170 0.2132 0.2548 0.3254 0.1735 0.1503 0.1925 0.2718 0.2835 0.4071 0.2047 0.3655 0.2259 0.3045 0.0970
208
Pb
+
232 Th
0.0014 0.0008 0.0113 0.0039 0.0008 0.0054 0.0047 0.0061 0.0028 0.0022 0.0049 0.0043 0.0043 0.0021 0.0063 0.0134 0.0054 0.0016
0.04513 0.17274 0.03317 0.03002 0.09498 0.13346 0.03033 0.03017 0.10292 0.03300 0.03063 0.02966 0.04836 0.03132 0.03336 0.03172 0.02814 0.11880
206
Pb
+
238 U
0.00279 0.00313 0.00162 0.00068 0.00074 0.00446 0.00080 0.00114 0.00264 0.00048 0.00068 0.00063 0.00087 0.00039 0.00073 0.00203 0.00061 0.00341
0.15508 0.66006 0.10614 0.09917 0.33524 0.49740 0.09664 0.09968 0.36701 0.10607 0.09962 0.09340 0.13371 0.10254 0.10049 0.10070 0.09589 0.43161
207
Pb
+
206 Pb
0.00142 0.00853 0.00138 0.00112 0.00184 0.01045 0.00174 0.00118 0.00525 0.00079 0.00077 0.00060 0.00166 0.00057 0.00054 0.00208 0.00103 0.00794
0.09925 0.28439 0.06226 0.05993 0.11438 0.18411 0.05925 0.06130 0.12924 0.06189 0.06297 0.06069 0.07049 0.06098 0.06417 0.05998 0.05742 0.15070
0.00096 0.00119 0.00184 0.00099 0.00058 0.00229 0.00120 0.00184 0.00074 0.00075 0.00158 0.00130 0.00150 0.00090 0.00318 0.00528 0.00201 0.00129
Apparent age (Ma) 208 232 892.2 3220.6 659.5 597.8 1834.0 2532.1 603.9 600.8 1980.1 656.2 609.8 590.7 954.6 623.4 663.3 631.2 560.9 2269.0
+
206 238
+
54.0 54.0 31.7 13.4 13.6 79.8 15.7 22.3 48.4 9.3 13.4 12.3 16.8 7.7 14.3 39.8 12.1 61.8
929.4 3267.4 650.3 609.6 1863.8 2602.6 594.7 612.5 2015.3 649.9 612.2 575.6 809.0 629.3 617.3 618.5 590.3 2312.9
7.9 33.2 8.0 6.6 8.9 45.1 10.2 6.9 24.8 4.6 4.5 3.5 9.4 3.3 3.2 12.2 6.1 35.8
207 206
I.d. Inf
+
+
1610.1 18.1 1610.0 3387.2 6.6 3387.0 683.0 64.5 649.6 601.2 36.0 609.7 1870.2 9.1 1870.0 2690.2 20.7 2690.0 576.4 44.6 595.1 649.8 65.9 611.7 2087.6 10.1 2087.0 670.4 26.2 649.5 707.4 54.4 610.2 628.3 46.8 574.6 942.6 44.4 942.6 638.5 32.0 629.1 747.1 108.5 614.5 602.7 202.8 618.8 507.6 79.1 591.8 2353.9 14.6 2354.0
18.0 7.0 8.0 6.6 9.0 21.0 10.2 6.9 10.0 4.6 4.5 3.6 44.0 3.4 3.3 11.8 6.2 15.0
Table 4. SHRIMP zircon data from biotite granite ARL-6 Analy- Struc- Pb* U Th Th/ sis ture ppm ppm ppm U
6.1 8.1 9.1 10.1 10.2 11.1 12.1 14.1 15.1 16.1 17.1 18.1 19.1 20.1 21.1 24.1 26.1 27.1 28.1 29.1 30.1 30.2
UZC CZC CZC UZC CZIO BZC UZC UZC CZC UZC CZC BZC CZC BZC UZC CZC UZC CZC CZC CZC SZC SZC
29 170 167 33 346 135 114 891 496 21 58 48 304 704 10 23 49 40 48 554 354 131 1382 573 34 317 162 48 560 206 46 328 227 13 118 139 10 88 50 13 189 81 9 64 110 12 141 54 19 133 210 136 223 133 10 83 60 76 167 3 68 63 21 147 166 106
0.98 0.39 0.56 0.83 0.01 0.82 0.64 0.41 0.51 0.37 0.69 1.17 0.57 0.43 1.72 0.38 1.57 0.60 0.72 0.02 0.34 0.64
204
Pb ppb
0 2 4 2 1 1 4 9 0 1 1 2 0 4 2 1 2 4 1 1 4 0
204
Pb
+
2.0E-05 8.1E-05 4.0E-05 1.3E-04 2.0E-06 5.6E-05 1.1E-04 8.0E-05 9.3E-06 2.0E-05 2.4E-05 2.2E-04 3.7E-05 3.5E-04 3.8E-04 6.4E-05 1.4E-04 4.1E-05 1.9E-04 2.0E-05 7.7E-05 4.8E-06
f 206Pb
206 Pb
2.0E-05 4.3E-05 2.5E-05 5.0E-05 2.9E-06 3.7E-05 3.1E-05 2.5E-05 8.6E-06 1.1E-05 1.9E-05 1.9E-04 2.3E-05 8.9E-05 3.1E-04 3.0E-05 8.0E-05 1.1E-05 8.2E-05 1.4E-05 2.3E-05 4.8E-06
0.00035 0.00144 0.00071 0.00171 0.00003 0.00081 0.00204 0.00142 0.00017 0.00035 0.00042 0.00390 0.00065 0.00631 0.00675 0.00114 0.00240 0.00055 0.00330 0.00028 0.00091 0.00006
208
Pb
+
206 Pb
0.2961 0.1247 0.1751 0.2545 0.0038 0.2381 0.1740 0.1272 0.1584 0.1255 0.2087 0.3511 0.1686 0.1777 0.5055 0.1184 0.4939 0.1575 0.2335 0.0042 0.0836 0.1746
208
Pb
+
232 Th
0.0032 0.0027 0.0021 0.0148 0.0002 0.0037 0.0021 0.0023 0.0030 0.0016 0.0041 0.0099 0.0045 0.0048 0.0160 0.0031 0.0097 0.0011 0.0061 0.0005 0.0023 0.0019
0.04363 0.02987 0.03773 0.08835 0.11152 0.11339 0.02210 0.02843 0.03152 0.02893 0.03865 0.02787 0.03035 0.02618 0.02938 0.02528 0.03353 0.13777 0.03530 0.10704 0.22518 0.19045
206
+
Pb
238 U
0.00086 0.00077 0.00054 0.00661 0.00582 0.00320 0.00032 0.00054 0.00069 0.00042 0.00085 0.00093 0.00096 0.00110 0.00115 0.00076 0.00082 0.00157 0.00105 0.01342 0.00694 0.00280
0.14503 0.09355 0.12010 0.28959 0.42661 0.38849 0.08108 0.09267 0.10159 0.08471 0.12791 0.09300 0.10320 0.06352 0.09986 0.08189 0.10668 0.52140 0.10894 0.45353 0.90448 0.69874
207
Pb
+
206 Pb
0.00187 0.00109 0.00075 0.01225 0.00253 0.00724 0.00045 0.00048 0.00099 0.00048 0.00100 0.00144 0.00141 0.00195 0.00202 0.00096 0.00133 0.00407 0.00130 0.00489 0.01078 0.00547
Cores: UZC, Unzoned core; CZC, Concentric and sector zoned core; BZC, Banded zoned core; SZC, Striated zoned core. Igneous overgrowths: CZIO, Concentric zoned igneous overgrowth; UZO, Unzoned overgrowth; UZS, Unzoned striated. *, Radiogenic. f, Correction for common Pb-fraction of total 206Pb (206PbT) that is common 206Pb (206Pbc).
0.06920 0.05910 0.06472 0.18475 0.17238 0.15039 0.05716 0.05912 0.06042 0.05837 0.06461 0.05520 0.06104 0.05943 0.05895 0.05717 0.06011 0.19835 0.05984 0.15824 0.30866 0.29285
0.00134 0.00183 0.00071 0.01411 0.00292 0.00180 0.00087 0.00061 0.00081 0.00100 0.00071 0.00395 0.00152 0.00181 0.00514 0.00214 0.00180 0.00077 0.00214 0.00084 0.01176 0.00144
Apparent age (Ma) 208 232 863.1 594.9 748.5 1711.2 2136.9 2171.0 441.8 566.6 627.2 576.5 766.4 555.6 604.4 522.3 585.4 504.7 666.6 2608.9 701.1 2055.3 4104.9 3523.6
+
206 238
+
16.7 15.2 10.6 123.1 106.1 58.2 6.3 10.5 13.6 8.2 16.6 18.3 18.8 21.6 22.6 14.9 16.1 27.9 20.6 246.5 114.8 47.6
873.1 576.5 731.1 1639.5 2290.4 2115.8 502.6 571.3 623.7 524.2 775.9 573.3 633.2 397.0 613.6 507.4 653.5 2705.1 666.6 2410.9 4152.8 3415.9
10.6 6.4 4.3 61.6 11.5 33.7 2.7 2.8 5.8 2.9 5.7 8.5 8.2 11.8 11.8 5.7 7.8 17.3 7.6 21.7 36.6 20.8
207 206
Inf Age
+
+
904.6 40.4 871.8 10.6 570.6 69.0 576.6 6.5 765.3 23.3 730.1 4.3 2696.0 132.0 2696.0 125.0 2580.8 28.6 2581.0 28.0 2350.4 20.6 2350.0 20.0 497.9 33.8 502.6 2.7 571.4 22.5 571.3 2.8 618.5 29.1 623.8 5.8 543.8 37.9 523.8 2.9 761.7 23.2 776.1 5.7 420.5 168.4 575.4 8.3 640.6 54.6 632.9 8.3 582.8 67.5 583.0 67.0 565.1 202.3 614.6 11.5 498.0 84.5 507.5 5.8 607.6 65.9 654.5 7.8 2812.7 6.4 2812.7 6.4 597.9 79.6 668.0 7.7 2436.9 9.0 2437.0 9.0 3514.2 60.1 3514.0 58.0 3432.8 7.7 3433.0 7.8
394
M. F. PEREIRA ET AL.
Fig. 4. Scanning electron microscope cathodoluminescence images of selected zircon grains from the Se´rie Negra paragneisses (Biscaia, BSC-1 and Casas Novas, CSN-26), to show the growth textures of each grain analysed for U– Th–Pb isotopes by SHRIMP. (a, b) (sample BSC-1): analysis 1-2 represents a spot on a simple subeuhedral crystal with concentric oscillatory zoning (mantled by a thin and irregular convolute zoned growth); analyses 7-1 and 3-2 refer to well-defined Palaeoproterozoic cores (7-1, unzoned core; 3-2, banded zoned core
OSSA– MORENA EDIACARAN BASINS, IBERIA
include the uncertainty in the Pb/U calibration (c. 0.3–0.5%). Ages were calculated using the constants recommended by the IUGS Subcommission on Geochronology (Steiger & Ja¨ger 1977). Common Pb corrections assumed a model common Pb composition appropriate to the age of each spot (Cumming & Richards 1975). Best estimates of the individual ages (inferred age, Tables 1– 4) were calculated from the radiogenic 207 Pb/206Pb (common Pb correction based on 204 Pb) for discordant zircon and zircon older than 1.5 Ga, and from the radiogenic 206Pb/238U (common Pb correction based on 207Pb) for concordant younger zircon. In a few cases where the uncertainty on the 204Pb correction was very large, the analysis was corrected using 208Pb; this means that radiogenic 208Pb/206Pb and 208Pb/232Th were not independently determined (Tables 1 and 2). Stratigraphic ages were based on the numerical time scale of Gradstein et al. (2004).
SHRIMP U –Th – Pb results The results of the U –Th– Pb analyses of detrital and inherited grains are listed in Tables 1–4 and plotted on modified Tera–Wasserburg diagrams in Figures 7 and 8. (Analyses of metamorphic overgrowths from the paragneisses and melt-precipitated zircon from the biotite granite will be reported elsewhere.) Most of the analyses are concordant or nearly concordant within analytical uncertainty. A compilation of the main group of inferred ages is illustrated as a relative probability density distribution (Dodson et al. 1988) in Figure 9.
Se´rie Negra paragneisses The Se´rie Negra immature sediments were deformed and metamorphosed to amphibolite facies in the Mississippian (Chichorro 2006). They are enriched in light rare earth elements (LREE), have small negative Eu anomalies, and have other chemical features (La/Th ¼ 3.30–4.52; Th/Sc ¼ 0.40–0.83; La/Sc ¼ 1.83–3.21; TiO2 0.69–0.88 wt%; Ni 35–55 ppm) consistent with derivation primarily from felsic continental crust (magmatic arc) and/or recycled detritus (Pereira et al. 2006a). Sample BSC-1 (Biscaia, Fig. 2), is from a medium-grained paragneiss consisting of quartz, biotite, muscovite, plagioclase and K-feldspar aligned parallel to the regional foliation. Albite porphyroblasts include aligned and folded graphitic
395
inclusions, and have tails of new-grown oligoclase. Blades of muscovite and biotite are oriented parallel to the foliation (also marked by dynamically recrystallized quartz ribbons). The few zircon grains recovered are small (50 –100 mm diameter), colourless, prismatic to subrounded crystals. CL images (Fig. 4) show that the grains consist of a simple or complex detrital zircon core surrounded by a thin, irregular zircon overgrowth that was probably formed during postdepositional metamorphic –hydrothermal events (Chichorro 2006). The simple crystals are commonly subhedral with some edges subrounded. Ion microprobe U –Th –Pb isotopic analyses of seven zircon cores are listed in Table 1 and plotted on a concordia diagram in Figure 7a. These represent a variety of textures: concentric zoning, banded zoning, unzoned and strongly luminescent unzoned cores or overgrowths. Despite the small number of zircon grains, different age groups can be distinguished: the two oldest detrital cores (7.1, 3.2, Fig. 4b), with normal igneous Th/U (0.27 and 0.59), are Palaeoproterozoic, c. 2.1 Ga; the strongly luminescent unzoned overgrowth (2.1), with low U (9 ppm) and moderate Th/U (0.29), possibly indicating metamorphic growth in a metaluminous matrix, is Cryogenian, c. 785 Ma; three grains with simple euhedral zoning (1.2, 4.1, 6.1, Fig. 4a) and moderate Th/U (0.66 –1.67) suggesting an igneous origin (Heaman et al. 1990; Hanchar & Miller 1993) are Ediacaran (c. 630–580 Ma), and were possibly produced by Avalonian –Cadomian–Pan –African magmatism. Sample SEC-1 (S. Escoural, Fig. 2) is from a foliated paragneiss with a mosaic texture consisting of alternating millimetre-wide bands of dynamically recrystallized quartz and feldspar, ribbons of polygonal quartz, and bands with biotite, quartz, feldspar, sillimanite, zircon, apatite, rutile and opaque minerals (mainly graphite). The zircon occurs as small (30–100 mm diameter), colourless to yellowish, equant to elongate grains, most with some preserved crystal facets. CL images show that most grains consist of a core, which is sometimes embayed, surrounded by a thin (,10 mm), heterogeneous metamorphic – hydrothermal overgrowth (Chichorro 2006). Some detrital cores are composite, consisting of an inner core and overgrowth. Some cores are subhedral prisms, consistent with limited transport and recycling, whereas others are strongly rounded, possibly having survived several cycles of erosion.
Fig. 4. (Continued) with dark-CL with narrow and longitudinal bright-CL zones). It should be noted that analysis 7-1 is discordant (see Fig. 7a). (c, d) (sample CSN-26) represent two complex zircons: analysis 5-2 is a spot on a weak oscillatory zoned centre and analysis 7-2 was made on an unzoned to weakly zoned core. (e, f) (sample CSN-26): examples of two simple crystal cores with concentric oscillatory zoning.
396
M. F. PEREIRA ET AL.
Fig. 5. Scanning electron microscope cathodoluminescence images of selected zircon grains from the Se´rie Negra paragneiss (S. Escoural, SEC-1), to show the growth textures of each analysed grain for U–Th– Pb isotopes by SHRIMP. (a, b) spots on cores of highly complex zircons with several growth generations: analyses 9-2 and 7-2 represent dark longitudinal zoned zircons (banded zoned cores) similar to zircon extracted from diorites (e.g. Hoskin 2000). Analysis 7-1 is a unzoned overgrowth (very discordant). (c) A core with a concentric oscillatory pattern
OSSA– MORENA EDIACARAN BASINS, IBERIA
A variety of growth textures was recognized: cores with concentric zoning (10.2, 26.1, Fig. 5c and h), banded zoning (9.2, 19.1, Fig. 5a and g), sector zoning (4.3, Fig. 5f) and no zoning (21.1, Fig. 5d), and inner overgrowths with concentric zoning (16.1, Fig. 5e) or no zoning (4.4, 21.2, Fig. 5d and f). The cores preserve at least three, and possibly four, generations of overgrowth that predate the post-depositional overgrowths that mainly formed in the Mississippian (Chichorro 2006). The three oldest cores (7.2, 9.2, 14.1), two of them with the banded zoning common in rapidly crystallized zircon (Fig. 5a and b), have Palaeo-Archaean ages of c. 3.30 –3.16 Ga. There is a small cluster of Palaeoproterozoic ages (c. 2.15–1.79 Ga), but most of the cores (55% of the total) are Neoproterozoic to early Palaeozoic (c. 850 –530 Ma). There is a marked absence of cores with Mesoproterozoic ages. Mixture modelling of the Neoproterozoic age group (Sambridge & Compston 1994) suggests that, as well as the Cryogenian component (31%), there are possibly two Ediacaran components, one at c. 615 Ma and the other at c. 580 Ma. Examples of the various zircon ages and textures are illustrated in Figure 5. Grain 16 (Fig. 5e) consists of an Ediacaran (c. 610 Ma) zoned igneous overgrowth on a Cryogenian (c. 845 Ma) igneous core. Both the sector zoned core and unzoned inner overgrowth of grain 4 (Fig. 5f) yield Ediacaran ages. The unzoned overgrowth on grain 21 (Fig. 5d) also yields an Ediacaran age, but it has the very low Th/U (0.02) commonly found in zircon crystallized during the partial melting of peraluminous rocks, usually metasediments (Williams & Claesson 1987; Williams 2001). Grains 26 (Fig. 5h), 12, 18, 20 and 27 are examples of whole cores of Ediacaran igneous zircon with simple concentric zoning. The apparent ages measured on grain 12 are anomalous; the 206Pb/238U age of the core is 526.2 + 5.8 (s) Ma, whereas that of the surrounding low-U overgrowth – recrystallized zone is 638 + 28 (s) Ma. It is possible that the ‘age’ of the core is too low because of radiogenic Pb loss, but equally likely that the ‘age’ of the overgrowth is an overestimate because of the presence of excess radiogenic Pb, particularly given the
397
presence of excess 204Pb. In neither case was there sufficient radiogenic 207Pb present for a measurement of the 207Pb/206Pb age to be precise enough to resolve the question. Whatever might have occurred, the age of grain 12, the lowest measured on a core from this rock, is an unreliable constraint on the maximum deposition age of the protolith of the S. Escoural paragneiss. A better constraint is provided by the age of the youngest detrital zircon sub-population, 578 + 12 Ma. Sample CSN-26 (Casas Novas, Fig. 2) is from a foliated paragneiss consisting of alternating millimetre-wide bands of dynamically recrystallized quartz and feldspar, and of quartz, biotite, fibrolitic sillimanite, zircon and graphite, in which there are porphyroblasts of plagioclase, andalusite and cordierite. The zircon grains are small (30–120 mm diameter) and colourless to light rosy pink. There is a wide range of morphologies, from rounded anhedral to prismatic subhedral with preserved crystal faces. The CL images show complex internal structures, and thin, irregular and unzoned metamorphic – hydrothermal overgrowths (Chichorro 2006). Many cores consist of fragments of igneous zircon with simple concentric zoning (e.g. grains 2, 3, Fig. 4e and f). Others are composite, consisting of an inner core with one or more overgrowths (e.g. grain 5, Fig. 4c). Five main core textures were distinguished: concentric zoning, strongly luminescent concentric zoning, sector zoning, banded zoning and unzoned. The 18 core analyses (Table 3) are insufficient for a detailed characterization of the detrital component, but nevertheless give a useful indication of the age pattern. The detrital population is dominated by Neoproterozoic zircon (70% of the analysed grains). Two of the remaining grains (4, 5) yield Archaean ages (c. 2.7 and 3.4 Ga) and three (7, 11, 20) are Palaeoproterozoic (c. 2.35–1.87 Ga). A very discordant analysis of the low Th/U (0.08) overgrowth on grain 4 possibly reflects a metamorphic event at c. 1.6 Ga. An age of c. 940 Ma measured on one core (grain 14) with banded zoning and a moderately high Th/U (1.1) possibly records an episode of Tonian mafic to intermediate magmatism in the source region
Fig. 5. (Continued) (analysis 10-2, Th/U ¼ 0.78) with a highly luminescent unzoned rim (analysis 10-1, Th/U ¼ 0.27; with high 207Pb/206Pb uncertainty). (d) Analysis 21-2 on a low-luminescence unzoned overgrowth over an inherited unzoned core (analysis 21-1; with a very low Th/U ¼ 0.02 indicative of a high-grade metamorphic event at c. 609 Ma). (e) A Cryogenian rounded concentric zoned core enclosed in an igneous Ediacaran overgrowth with concentric oscillatory zoning. (f) A sector zoned core with 207Pb/206Pb age of 618 + 6.8 Ma mantled by an unzoned intermediate overgrowth with an age of 610.3 + 5.1 Ma. (g ) A planar oscillatory zoned core with elongate wide dark and bright CL zones. (h) Zircon with typical igneous concentric zoning with 207Pb/206Pb age of 618 + 6.9 Ma, growing over a very small rounded inherited core.
398
M. F. PEREIRA ET AL.
Fig. 6. Scanning electron microscope cathodoluminescence images of selected zircon grains from a biotite granite (Arraiolos, ARL-6), to show the growth textures of each inherited grain analysed for U– Th–Pb isotopes by SHRIMP. (a–c) are representative of complex crystals of zircon with several growth generations: (a ) a Palaeo-Archaean bright sector zoned core; (b) a Palaeoproterozoic core with weak oscillatory concentric zoning; (c) a Palaeoproterozoic very anhedral banded zoned core. (d–h) Neoproterozoic (Cryogenian and Ediacaran) simple crystal cores enclosed by igneous overgrowths of different thickness. These cores have oscillatory concentric zoning (analyses 17-1, 19-1 and 8-1), bright unzoned (analysis 26-1) or unzoned (analysis 14-1). (i) An anhedral igneous Cambrian unzoned core enclosed by the new igneous generation.
OSSA– MORENA EDIACARAN BASINS, IBERIA
Fig. 7. Modified Tera– Wasserburg concordia diagrams for zircon from the Se´rie Negra paragneisses: (a) sample BSC-1; (b) sample SEC-1. Inset shows a probability density diagram for detrital components.
399
400
M. F. PEREIRA ET AL.
Fig. 8. Modified Tera– Wasserburg concordia diagrams for zircon from: (a) sample CSN-26 (Se´rie Negra paragneiss) and (b) sample ARL-6 (biotitic granite from Arraiolos). Inset shows a probability density diagram for detrital and inherited components.
OSSA– MORENA EDIACARAN BASINS, IBERIA
(Hoskin 2000). There is a marked absence of Mesoproterozoic detrital zircon. The 11 analyses in the Neoproterozoic group yield a range of ages, but there are insufficient to define discrete sub-populations. Two grains (9, 15) are distinctly older than the rest (c. 650 Ma) and one (2.2) is younger (c. 575 Ma). Seven of the remaining eight analyses yield a very similar age, c. 610 Ma, which probably records a discrete thermal event in the sediment’s source region.
Biotite granite from Arraiolos Sample ARL-6 is from a medium-grained, weakly foliated granite composed of plagioclase, Kfeldspar, quartz and biotite with a weakly peraluminous calc-alkaline affinity (SiO2 = 71.3 wt%, Al2O3 = 14.45 wt%, MgO = 0.55 wt%, Fe2O3 = 2.01 wt%, TiO2 = 0.2 wt%, CaO = 2.13 wt%, Na2O = 3.43 wt%, K2O = 4.12 wt%). The rock has a negative Eu anomaly in its chondritenormalized REE pattern, a slight enrichment in large ion lithophile elements (LILE) and a negative
401
Nb anomaly (Th/Nb ¼ 1.77; La/Nb ¼ 3.11). The granite occurs as 1–10 m wide dykes cutting sheared enriched mid-ocean ridge basalt (E-MORB) amphibolites (Th/Yb ¼ 0.31; Ce/Yb ¼ 7.06; Pereira et al. 2004, 2007) and mica schists of the igneous (mafic-dominated)–sedimentary complex. The zircon population is morphologically very diverse. CL imaging shows that nearly 80% of the crystals consist of a core surrounded by an unzoned or weakly zoned overgrowth. The relative volume of the overgrowth ranges from very high (c. 95%, Fig. 6i) to high (c. 75%, Fig. 6g) to very low (c. 30%, Fig. 6d). The zircon mostly occurs as subhedral to euhedral, equant to moderately elongate prisms (aspect ratios 1 –4), some with smoothly rounded terminations, many with welldeveloped f211g crystal faces, a feature of zircon grains grown on an anhedral or subhedral nucleus (Williams 2001). The cores have a wide range of zoning textures. Many have the simple euhedral zoning commonly found in zircon precipitated from felsic to intermediate granitic magmas (e.g. Fig. 6d and f); others have banded zoning, as
Fig. 9. Probability density diagram for inherited or detrital zircon from the Se´rie Negra paragneisses (Biscaia, BSC-1; S. Escoural, SEC-1; Casas Novas, CSN-26) and the biotite granite (Arraiolos, ARL-6).
402
M. F. PEREIRA ET AL.
Fig. 10. Schematic palaeogeographical map of the Gondwanan supercontinent at 570 Ma (modified from Linnemann et al. 2004, and references therein) with zircon ages from peri-Gondwana microplates (Avalonian–Cadomian Arc: Ossa– Morena, West Avalonia, Saxo-Thuringia), the Arabian Shield, the West African craton and the Amazonian craton.
found in rapidly crystallized zircon (e.g. Fig. 6c), or no zoning at all (e.g. Fig. 6i). Some have sector zoning, as found in the zircon from more mafic igneous rocks or some metaluminous granulites (e.g. Fig. 6a), and others appear to have more than one generation of zircon growth (e.g. Fig. 6c). A minor component of the population (c. 10%) consists of very elongate crystals, mostly with weak luminescence, that are either unzoned or have longitudinal zoning and no inherited cores (e.g. the crystal on the right side of Fig. 6g). The U –Th –Pb isotopic analyses of the cores of 20 grains are listed in Table 4 and plotted on a Tera– Wasserburg concordia diagram in Figure 8d. The pattern of ages closely resembles
those of the detrital zircon from the paragneisses. The oldest grain (30), analysed twice because the first analysis showed a large excess of radiogenic Pb, is Palaeo-Archaean (c. 3.43 Ga). Four grains (10, 11, 27, 29) are Neo-Archaean to Palaeoproterozoic (c. 2.81–2.35 Ga). Two of those (10, 29) have very low Th/U (,0.1), typical of zircon precipitated during the high-grade metamorphism or partial melting of peraluminous rocks (Williams & Claesson 1987; Williams 2001). Sixty per cent of the inherited cores are Neoproterozoic, with ages ranging from c. 870 to c. 575 Ma. There is no strong clustering of the ages, but nearly half the grains in the group have ages in the narrow range 670–570 Ma, coinciding with the ages
OSSA– MORENA EDIACARAN BASINS, IBERIA
of Cryogenian –Ediacaran Pan-African–Cadomian events. Three inherited cores (grains 12, 16, 24, Fig. 6i) give Cambrian ages (c. 523 –502 Ma). Assuming that these apparent ages have not been reduced by radiogenic Pb loss, they place an upper limit of Cambrian on the deposition age of the sedimentary component of the granite magma.
Discussion The oldest rocks exposed in the Ossa – Morena Zone are a monotonous sequence of turbidite-like detrital sediments deposited in a late Neoproterozoic shallow-shelf environment (Se´rie Negra Formation; Carvalhosa 1965), probably at a passive continental margin (Quesada 1990). The basement to this sequence has not yet been identified (Eguiluz 1987; Quesada 1990; Vidal et al. 1994). The youngest detrital zircon found so far implies a maximum depositional age of c. 560 –550 Ma (Scha¨fer et al. 1993; Ordon˜ez-Casado 1998; Ferna´ndez-Sua´rez et al. 2002; Chichorro et al. 2006; Pereira et al. 2006b). This thick Ediacaran sequence of graphitic shales, quartzwackes and greywackes interbedded with black cherts and quartzites passes upward into a volcanogenic complex, a series of volcaniclastic and epiclastic sedimentary rocks and their calc-alkaline plutonic and subvolcanic equivalents, the Malcocinado and San Jeronimo Formations (Lin˜an et al. 1984; Eguiluz 1987; Lin˜an & Quesada 1990; Eguiluz & Abalos 1992; Pin et al. 2002). This complex has been interpreted as part of a magmatic arc which was active at least from c. 590 to 540 Ma (Sanchez-Carretero et al. 1989; Scha¨fer 1990; Quesada 1991; Oschner 1993; Eguiluz et al. 2000; Sanchez-Garcia et al. 2003; Pereira et al. 2006a). The Ediacaran Se´rie Negra Formation and the Malcocinado Formation are overlain unconformably by an igneous –sedimentary complex interpreted as a long-lived Cambrian to Early Ordovician (c. 530 –470 Ma) rift sequence (Lin˜an & Quesada 1990; Eguiluz et al. 2000; Sanchez-Garcia et al. 2003; Extebarria et al. 2006; Pereira & Quesada 2006). As previously noted by Scha¨fer et al. (1993), Ordon˜ez-Casado (1998) and Ferna´ndez-Sua´rez et al. (2002), the U –Pb ages measured on detrital zircon from the Se´rie Negra sediments match those of zircon from Archaean and Proterozoic crustal components from Gondwana (Gebauer et al. 1989). The ages of detrital and inherited zircon from Se´rie Negra paragneisses and biotite granite exposed in the E´vora Massif measured in the present study have a similar characteristic signature (Figs 7– 9). Most of the zircon is of Late Neoproterozoic age; the remainder is Palaeoproterozoic (2.4–1.8 Ga) and Archaean (3.5–2.5 Ga). There is a remarkable lack of Mesoproterozoic zircon
403
(1.8– 0.9 Ga), and in particular of zircon related to Grenvillian events (1.1–0.9 Ga). The high percentage of Cryogenian –Ediacaran zircon (c. 60% of grains analysed) overlaps in age (700 –550 Ma) with the Avalonian –Cadomian orogenic events at the peri-Gondwana margin (Nance & Murphy 1994; Nance et al. 2002; Linnemann et al. 2004), and coincides in part with the ages of Pan-African events in West Gondwana (Alkmim et al. 2001; Ennih & Lie´geois 2001; Neves 2003). Mixture modelling (Sambridge & Compston 1994) of the 43 Neoproterozoic zircon ages measured during this study (Fig. 9) suggests the presence of two main components with ages of c. 575 Ma (35%) and c. 615 Ma (50%). The zircon ages obtained here from the Se´rie Negra metasediments and related granite match the range of inherited zircon ages measured on other rocks from Europe and North Africa that have an affinity with the West African craton (Fig. 10), as follows. (1) Ediacaran greywacke (Se´rie Negra, Ossa– Morena Zone, SW Spain) with late Neoproterozoic (Ediacaran, 600–540 Ma), Palaeoproterozoic (2.0–1.8 Ga), Neo-Archaean (2.8–2.5 Ga) and Palaeo-Archaean (3.4 Ga) zircon ages (Ordon˜esCasado 1998). (2) Mississippian nebulite (Ossa –Morena Zone, south Spain) generated by partial melting of Se´rie Negra sediments and also Cambrian rocks with Neoproterozoic (Cryogenian–Ediacaran – Cambrian, 700–500 Ma), Palaeoproterozoic (2.0– 1.7 Ga) and Neo–Meso-Archaean (2.95–2.65 Ga) zircon ages (de la Rosa et al. 2002). (3) Ediacaran greywacke (Saxo-Thuringia Zone, eastern Germany) with Late Neoproterozoic (Cryogenian, 725–637 Ma; Ediacaran, 614–555 Ma), Palaeoproterozoic (2.0–1.8 Ga) and Neo-Archaean (2.5 Ga) zircon ages (Linnemann et al. 2004). (4) Cambrian and Ordovician sandstones (northern Sahara Platform, NE Algeria) with Late Neoproterozoic (Cryogenian–Ediacaran, 700–600 Ma), Palaeoproterozoic and a few Neoproterozoic– Mesoproterozoic zircon ages (Williams et al. 2002). (5) Ordovician volcaniclastic rocks (Central Iberian Zone, central Portugal) with Late Neoproterozoic (Cryogenian–Ediacaran, 696–577 Ma), Palaeoproterozoic (2.3 Ga) and Meso – Palaeo-Archaean (3.2– 3.1 Ga ) zircon ages (Sola´ et al. 2006). (6) Ordovician sandstone (Saxo-Thuringia Zone, eastern Germany) with Late Neoproterozoic (Cryogenian–Ediacaran, 651–546 Ma), Palaeoproterozoic (2.0–1.7 Ga) and Neo-Archaean (2.6 Ga) zircon ages (Linnemann et al. 2004). Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) zircon ages obtained from Se´rie Negra greywackes of the northern
404
M. F. PEREIRA ET AL.
central domains of the Ossa –Morena Zone also overlap with the results of our study, including the characteristic lack of Grenvillian ages (Ferna´ndezSua´rez et al. 2002; Gutierrez-Alonso et al. 2003; Pereira et al. 2006b). Zircon-forming events of Grenvillian age (1.1–0.9 Ga) are a feature of the Amazonian craton, the Arabian –Nubian shield and SE Gondwana (Keppie et al. 1998; Williams 2001; Williams et al. 2002; Avigad et al. 2003; Goodge et al. 2004; Linnemann et al. 2004). The comparison of our results with other U– Pb ages obtained from different peri-Gondwanan regions from Europe and Arabia (north and central Spain, eastern Germany, Bulgaria, NE Israel and Jordan) is useful in defining a different group of rocks containing a significant component derived from Grenvillian-aged sources (principally the southern Mozambique Belt; Williams et al. 2002), as follows. (1) Cambrian –Ordovician sandstones (Elat Area, NE Israel and Jordan) with late Neoproterozoic (Cryogenian– Ediacaran, 650 –550 Ma), Mesoproterozoic–Neoproterozoic (1.1–0.9 Ga) and Neo-Archaean– Palaeoproterozoic (2.7– 1.6 Ga) zircon ages (Avigad et al. 2003; Kolodner et al. 2006). (2) Ediacaran–Ordovician sediments (Central Iberian, Cantabrian and West Asturian –Leonese Zones, northern Spain) with Cambrian (540 Ma), Neoproterozoic (Cryogenian– Ediacaran, 800 –500 Ma), Mesoproterozoic–Neoproterozoic (1.3– 0.9 Ga), Palaeoproterozoic (2.3–1.8 Ga) and Neo-Archaean (2.7–2.5 Ga) zircon ages (Ferna´ndez-Sua´rez et al. 2002; Gutierrez-Alonso et al. 2003). (3) Mississippian orthogneisses (Central Iberian Zone, central Spain) derived from an Ediacaran protolith (c. 546 Ma) containing Neoproterozoic (Tonian, 980 Ma; Cryogenian, 830 Ma; Ediacaran, 582 Ma) zircon (Zeck et al. 2003). (4) Mississippian paragneiss (Sredna Gora Zone, western Bulgaria) derived from a Cambrian (500 Ma) protolith including Neoproterozoic– Cambrian (700–500 Ma), Mesoproterozoic –Neoproterozoic (1.1–0.9 Ga), Palaeoproterozoic (2.3–1.9 Ga) and Neo-Archaean (2.7–2.5 Ga) zircon (Carrigan et al. 2006). Sediments with West African craton affinity are considered to be derived predominantly from three possible sources that lack 1.1–0.9 Ga zirconforming events: the Avalonian –Cadomian belt, the Pan-African belt and the West African craton. The dominant source for the sediments containing Grenvillian detritus is probably the Neoproterozoic– Cambrian collision zone between East and West Gondwana, which uplifted the remnants of an earlier Grenvillian orogenic belt. Sediments with an identical detrital zircon age signature are found
in South Africa and along the palaeo-Eastern Gondwana margin of Antarctica and eastern Australia (Armstrong et al. 1998; Williams 2001; Williams et al. 2002; Goodge et al. 2004). The similarity between the detrital or inherited zircon ages from the different regions of Europe and North Africa that have an affinity to the West African craton suggests a palaeogeographical link between these regions and the continental source of the sediments. The prominent Cryogenian – Ediacaran age group (700–540 Ma) spans the period of the main episodes of tectonothermal activity related to arc construction along the periGondwanan palaeocontinent margin, the Avalonian–Cadomian events (Nance & Thompson 1996) and orogeny within Gondwana, the Pan-African events (Ennih & Lie´geois 2001; Neves 2003). The Cryogenian was a long period of glaciation (the Varanger ice age) with widely distributed glacial deposits (Knoll 2000), probably promoted by extensive cratonization (build-up of a palaeosupercontinent). The poorly diversified Ediacaran fauna on all continents (Narbonne 1998) is also consistent with the presence of a huge continental block at that time. The end of continental active margin consolidation during the Late Neoproterozoic (Avalonian– Cadomian events) and the protracted amalgamation of East and West Gondwana (Pan-African events) predate the Cambrian rifting. Several unconformities in the Cambrian and Ordovician sequences of southern Europe indicate that the basement was actively uplifted and eroded through that period. The overlying Early Palaeozoic platform sequences accumulated in shallower water in widely connected peri-Gondwanan basins. This palaeogeographical interconnection based on facies development is consistent with the distribution of Late Proterozoic to Early Palaeozoic faunas (Moczydlowska 1997; Gubanov 2002; Robardet 2002, 2003). SHRIMP U– Th–Pb age determinations on detrital and inherited zircon extracted from Se´rie Negra paragneisses and the Arraiolos biotite granite from the E´vora Massif (Ossa –Morena Zone, SW Iberian Massif, Portugal) indicate that these rocks are related products of recycled detritus from a region with a protracted history of zircon-growth events. Of the 72 dated zircon grains that had Precambrian ages, 62% were Neoproterozoic (69% Ediacarian, 27% Cryogenian). The majority of these had moderate Th/U ratios (0.3–1.7) and simple concentric oscillatory zoning, indicative of a predominantly felsic to intermediate plutonic igneous origin (Heaman et al. 1990; Hanchar & Miller 1993; Hoskin 2000). The relatively few detrital grains and inherited cores with banded zoning were probably derived from rocks in which the
OSSA– MORENA EDIACARAN BASINS, IBERIA
zircon crystallized more rapidly, either volcanic rocks or marginally Zr-saturated mafic plutonic rocks, such as diorites or gabbroic diorites (Hoskin 2000). These zircon characteristics are consistent with the compatible trace element contents of the studied samples, which indicate a felsic provenance linked to a continental magmatic arc. Only one zircon (SEC-1, grain 21) was found to have an Ediacaran overgrowth with the low Th/U ratio (0.02) common in zircon grown during the high-grade metamorphism of a peraluminous protolith (Williams & Claesson 1987; Williams 2001). The ages of the youngest detrital zircon or inherited core in each sample, 542.6 + 17 Ma (BSC-1), 526.2 + 5.8 Ma (SEC-1), 574.6 + 3.6 Ma (CSN-26) and 502.6 + 2.7 Ma (ARL-6), provided there has been negligible radiogenic Pb loss, place an upper limit on the deposition age of the Se´rie Negra sediments of mid-Cambrian. A more conservative limit, based on the youngest zircon population, is that the sediment was deposited no earlier than 575 Ma, late Ediacaran. The large contribution of Cryogenian and Ediacaran (700 –540 Ma) detrital zircon to the protolith sediment indicates strong basement denudation close to the final stages of the Neoproterozoic continental consolidation at the peri-Gondwana margin (Cadomian–Avalonian orogeny). The insignificant amounts of Mesoproterozoic (1.8– 0.9 Ga) zircon suggest a continental source that lacked both Grenvillian zircon-forming events and detritus derived from areas affected by such events. Such a possible source is Central and North Africa. The presence of a small component of Palaeoproterozoic (2.4–1.8 Ga) and Archaean (3.5–2.5 Ga) zircon is consistent with such a source. This conclusion is in agreement with the detrital zircon ages reported from other regions of Europe (Portugal, Spain and Germany) and North Africa (Algeria). The geodynamic regime that decisively contributed to the deposition of these major siliciclastic sedimentary sequences, associated with igneous activity, on an extensive peri-Gondwanan margin was most probably located close to the West African craton (Linnemann et al. 2004). The ideas presented here resulted from stimulating discussions with U. Linnemann and C. Quesada. This work is a contribution to the projects: CGL 2004-06808-C04-02-BTE, ‘Estudo geoquimico, tectonico y experimental de los processos de reciclage cortical y interaccion manto-corteza’ IGCP 497, ‘The Rheic Ocean: Its origin, evolution and correlatives’; and IGCP 485, ‘Cratons, metacratons and mobile belts: keys from the West African craton boundaries’. The quality of this paper was significantly improved following constructive reviews by F. Bussy, C. Carrigan and J.-P. Lie´geois.
405
References A LKMIM , F. F., M ARSHAK , S. & F ONSECA , M. A. 2001. Assembling West Gondwana in the Neoproterozoic: Clues from the Sa˜o Francisco craton region, Brazil. Geology, 29, 319– 322. A RMSTRONG , R. A., DE W IT , M. J., R EID , D., Y ORK , D. & Z ARTMAN , R. 1998. Cape Town’s Table Mountain reveals rapid Pan African uplift of its basement rocks. Journal of African Earth Sciences, 27, 10–11. A VIGAD , D., K OLODNER , K., M C W ILLIAMS , M., P ERSING , H. & W EISSBROD , T. 2003. Origin of northern Gondwana Cambrian sandstone revealed by detrital zircon SHRIMP dating. Geology, 331, 227– 230. B EETSMA , J. J. 1995. The Late Proterozoic/Palaeozoic and Hercynian crustal evolution of the Iberian Massif, N Portugal, as traced by geochemistry and Sr–Nd– Pb isotope systematics of pre-Hercynian terrigenous sediments and Hercynian granotoids. PhD thesis, University of Amsterdam. C ARRIGAN , C. W., M UKASA , S. B., H AYDOUTOV , I. & K OLCHEVA , K. 2006. Neoproterozoic magmatism and Carboniferous high-grade metamorphism in the Sredna Gora Zone, Bulgaria: An extension of the Gondwana-derived Avalonian– Cadomian belt? Precambrian Research, 141, 404–416. C ARVALHOSA , A. 1965. Contribuic¸a˜o para o conhecimento geolo´gico da regia˜o entre Portel e Ficalho (Alentejo). Memo´ria dos Servic¸os Geolo´gicos de Portugal, 11. C ARVALHOSA , A. 1983. Esquema geolo´gico do Macic¸o de E´vora. Comunicac¸o˜es dos Servic¸os Geolo´gicos de Portugal, 69, 201–208. C ARVALHOSA , A. 1999. Carta Geolo´gica de Portugal, Noticia Explicativa da Folha 36-C (Arraiolos), Instituto Geolo´gico e Mineiro, scale 1:50 000. C ARVALHOSA , A. & Z BYSZEWSKI , G. 1994. Carta Geolo´gica de Portugal, Noticia Explicativa da Folha 35-D (Montemor-o-Novo), Instituto Geolo´gico e Mineiro, scale 1:50 000. C ARVALHOSA , A., G ALOPIM DE C ARVALHO , A. M., M ATOS A LVES , C. A. & P INA , H. L. 1969. Carta Geolo´gica de Portugal, Noticia Explicativa da Folha 40-A (E´vora). Servic¸os Geolo´gicos de Portugal, scale 1:50 000. C HICHORRO , M. 2006. Estrutura do Sudoeste da Zona de Ossa–Morena: A´rea de Santiago de Escoural– Cabrela (Zona de Cisalhamento de Montemor-o-Novo, Macic¸o de E´vora). PhD dissertation, Universidade de E´vora. C HICHORRO , M., P EREIRA , M. F., W ILLIAMS , I. & S ILVA , J. B. 2006. Clues for Cadomian orogenic events in SW Iberian Massif: U/Pb– SHRIMP zircon evidence from the Se´rie Negra sediments (Escoural Formation, Ossa– Morena Zone, Portugal). In: M IRAO , J. & B ALBINO , A. (eds) Proceedings, VII Congresso Nacional de Geologia, Universidade de Evora, Estremoz, Portugal, Vol. I, 25– 28. C UMMING , G. L. & R ICHARDS , J. R. 1975. Ore lead isotope ratios in a continuously changing Earth. Earth and Planetary Science Letters, 28, 155– 171.
406
M. F. PEREIRA ET AL.
R OSA , J. D., J ENNER , G. A. & C ASTRO , A. 2002. A study of inherited zircons in granitoid rocks from the South Portuguese and Ossa– Morena Zones, Iberian Massif: support for the exotic origin of the South Portuguese Zone. Tectonophysics, 352, 245– 256. D ODSON , M. H., C OMPSTON , W., W ILLIAMS , I. S. & W ILSON , J. F. 1988. A search for ancient detrital zircons in Zimbabwean sediments. Journal of the Geological Society, London, 145, 977– 983. E GUILUZ , L. 1987. Petrogenesis de rocas igneas y metamorficas en el antiforme Burgillos–Monesterio. Macizo Iberico Meridional. PhD thesis, Universidad del Pais Vasco, Bilbao. E GUILUZ , L. & A BALOS , B. 1992. Tectonic setting of Cadomian low-pressure metamorphism in the central Ossa–Morena Zone (Iberian Massif). Precambrian Research, 56, 113 –137. E GUILUZ , L., G IL I BARGUCHI , J. I., A BALOS , B. & A PRAIZ , A. 2000. Superposed Hercynian and Cadomian orogenic cycles in the Ossa– Morena Zone and related areas of the Iberian Massif. Geological Society of America Bulletin, 112, 1398– 1413. E NNIH , N. & L IE´ GEOIS , J.-P. 2001. The Moroccan Anti-Atlas: 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 XTEBARRIA , M., C HALOT -P RAT , F., A PRAIZ , A. & E GUILUZ , L. 2006. Birth of a volcanic passive margin in Cambrian time: Rift palaeography of the Ossa– Morena Zone, SW Spain. Precambrian Research, 147, 366–386. F ERNA´ NDEZ -S UA´ REZ , J., G UTIE´ RREZ -A LONSO , G. & J EFFRIES , T. E. 2002. The importance of alongmargin terrane transport in northern Gondwana: insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany. Earth and Planetary Science Letters, 204, 75–88. G EBAUER , D., W ILLIAMS , I. S., C OMPSTON , W. & G RU¨ NENFELDER , M. 1989. The development of the central European continental crust since the Early Archaean based on conventional and ion-microprobe dating of up to 4.84 b.y. old detrital zircons. Tectonophysics, 157, 81–96. G OODGE , J. W., W ILLIAMS , I. S. & M YROW , P. 2004. Provenance of Neoproterozoic and lower Palaeozoic siliciclastic rocks of the central Ross orogen, Antarctica: detrital record of rift-, passive-, and activemargin sedimentation. Geological Society of America, Bulletin, 116, 1253– 1279. G RADSTEIN , F. M., O GG , J. G., S MITH , A. G., B LEEKER , W. & L OURENS , L. J. 2004. A new geological time scale with special reference to Precambrian and Neogene. Episodes, 27, 83–100. G UBANOV , A. 2002. Early Cambrian palaeogeography and the probable Iberia– Siberia connection. Tectonophysics, 352, 153–168. G UTIERREZ -A LONSO , G., F ERNANDEZ -S UAREZ , J., J EFFRIES , T. E., J ENNER , G. A., T UBRETT , M. N., C OX , R. & J ACKSON , S. E. 2003. Terrane accretion and dispersal in the northern Gondwana margin. An Early Palaeozoic analogue of a long-lived active margin. Tectonophysics, 365, 221–232. DE LA
H ANCHAR , J. M. & M ILLER , C. F. 1993. Zircon zonation patterns as revealed by cathodoluminescence and backscattered electron images: Implications for interpretation of complex crustal histories. Chemical Geology, 110, 1 –13. H EAMAN , L. M., B OWINS , R. & C ROCKET , J. 1990. The chemical composition of igneous zircon studies: implications for geochemical tracer studies. Geochimica et Cosmochimica Acta, 54, 1597– 1607. H OSKIN , P. W. O. 2000. Patterns of chaos: Fractal statistics and the oscillatory chemistry of zircon. Geochimica et Cosmochimica Acta, 64, 1905– 1923. J ULIVERT , M. 1987. The structure and evolution of the Hercynian Fold Belt in the Iberian Peninsula. In: S CHAER , J.-P. & R ODGERS , J. (eds) The Anatomy of Mountain Belts. Princeton University Press, Princeton, NJ, 65– 103. J ULIVERT , M., F ONTBOTE , J. M., R IBEIRO , A. & C ONDE , L. 1972. Mapa tecto´nico de la Penı´nsula Ibe´rica y Baleares, scale 1:1 000 000. Instituto Geolo´gico y Minero de Espan˜a. K EPPIE , J. D., D AVIS , D. W. & K ROGH , T. E. 1998. U–Pb geochronological constraints on Precambrian stratified units in the Avalon Composite Terrane of Nova Scotia, Canada: tectonic implications. Canadian Journal of Earth Sciences, 35, 222–236. K NOLL , A. H. 2000. Learning to tell Neoproterozoic time. Precambrian Research, 100, 3–10. K OLODNER , K., A VIGAD , D., M C W ILLIAMS , M. & I RELAND , T. R. 2006. U–Pb SHRIMP dating of detrital zircons from Palaeozoic and Mesozoic sandstone in Israel and Jordan. Palaeozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia. Geophysical Research Abstracts, 8, 00387. L IN˜ AN , E. & Q UESADA , C. 1990. Ossa– Morena Zone: 2. Stratigraphy. In: D ALLMEYER , R. D. & M ARTINEZ G ARCIA , E. (eds) Pre-Mesozoic Geology of Iberia. Springer, Berlin, 229–266. L IN˜ AN , E., P ALACIOS , T. & P EREJON , A. 1984. Precambrian– Cambrian boundary and correlation from southwestern and central parts of Spain. Geological Magazine, 121, 122– 228. L INNEMANN , U., M C N AUGHTON , N. J., R OMER , R. L., G EHMLICH , M., D ROST , K. & T ONK , C. 2004. West African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana?: U/Pb-SHRIMP zircon evidence and the Nd-isotopic record. Geologische Rundschau, 93, 683–705. M ATTE , P. 2001. The Variscan collage and orogeny (480–290 Ma) and the tectonic definition of the Armorica microplate: a review. Terra Nova, 13, 122–128. M OCZYDLOWSKA , M. 1997. Proterozoic and Cambrian successions in Upper Silesia: an Avalonian terrane in southern Poland. Geological Magazine, 134, 679–689. M URPHY , J. B., E GUILUZ , L. & Z ULAUF , G. 2002. Cadomian orogens, peri-Gondwanan correlatives and Laurentia–Baltica connections. Tectonophysics, 352, 1 –9.
OSSA– MORENA EDIACARAN BASINS, IBERIA M URPHY , J. B., G UTIERREZ -A LONSO , G., N ANCE , R. D. ET AL . 2006. Origin of the Rheic Ocean: Rifting along a Neoproterozoic suture? Geology, 34, 325–328. N A¨ GLER , T. 1990. Sm– Nd, Rb– Sr and common lead isotope geochemistry on fine-grained sediments of the Iberian Massif. PhD thesis, ETH, Zurich. N ANCE , R. D. & M URPHY , J. B. 1994. Contrasting basement signatures and the palinspastic restoration of peripheral orogens: an example from the Neoproterozoic Avalonian– Cadomian belt. Geology, 22, 617–620. N ANCE , R. D. & M URPHY , J. B. 1996. Basement isotopic signatures and Neoproterozoic palaeogeography of Avalonian– Cadomian and related terranes in the circum-North Atlantic. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 333–346. N ANCE , R. D. & T HOMPSON , M. D. (eds) 1996. Avalonian and Related Peri-Gondwanan Terranes of the circum-North Atlantic. Geological Society of America-Special Papers, 304. N ANCE , R. D., M URPHY , J. B. & K EPPIE , J. D. 2002. A cordilleran model for the evolution of Avalonia. Tectonophysics, 352, 11–31. N ARBONNE , G. M. 1998. The Ediacara Biota: a terminal Neoproterozoic experiment in the evolution of life. GSA Today, 8, 1– 6. N EVES , S. P. 2003. Proterozoic history of the Borborena province (NE Brazil): Correlations with neighboring cratons and Pan-African belts and implications for the evolution of western Gondwana. Tectonics, 22, 1031–1045. O LIVEIRA , J. T. (coord.) 1992. Carta Geolo´gica de Portugal: Folha Sul. Instituto Geolo´gico e Mineiro de Portugal, scale 1:500 000. O LIVEIRA , J. T., O LIVEIRA , V. & P IC¸ ARRA , J. M. 1991. Trac¸os gerais da evoluc¸a˜o tectonoestratigrafica da Zona de Ossa Morena em Portugal. Cuadernos Laboratorio Xeologico Laxe, Corun˜a, 16, 221– 250. O RDON˜ EZ -C ASADO , B. 1998. Geochronological studies of the Pre-Mesozoic basement of the Iberian Massif: the Ossa–Morena Zone and the Allochthonous Complexes within the Central Iberian Zone. PhD dissertation, ETH, Zurich. O SCHNER , A. 1993. U–Pb geochronology of the Upper Proterozoic –Lower Palaeozoic geodynamic evolution in the Ossa–Morena Zone (SW Iberia): Constraints on the timing of the Cadomian Orogeny. PhD thesis, ETH, Zurich. P EREIRA , M. F. & Q UESADA , C. (eds) 2006. Ediacaran to Vise´an crustal growth processes in the Ossa–Morena Zone (SW Iberia). IGCP 497 Evora Meeting. Publicaciones Instituto Geologico y Minero de Espana. P EREIRA , M. F., S ILVA , J. B. & C HICHORRO , M. 2003. Internal structure of the E´vora high-grade terrains and the Montemor-o-Novo shear zone (Ossa– Morena Zone, Portugal). Geogaceta, 33, 79– 82. P EREIRA , M. F., C HICHORRO , M., S ANTOS , J. F., M OITA , P. & S ILVA , J. B. 2004, Geochemistry of lower Palaeozoic anorogenic basic rocks from the E´vora Massif (Western Ossa–Morena Zone, Portugal). Geogaceta, 35, 87–91.
407
P EREIRA , M. F., C HICHORRO , M., L INNEMANN , U., E GUILUZ , L. & S ILVA , J. B. 2006a. Inherited arc signature in Ediacaran and Early Cambrian basins of the Ossa–Morena Zone (Iberian Massif, Portugal): palaeogeographic link with European and North African Cadomian correlatives. Precambrian Research, 144, 297–315. P EREIRA , M. F., L INNEMANN , U., D ORST , K., C HICHORRO , M. & J EFFRIES , T. E. 2006b. Provenances of Ediacaran and Early Cambrian siliciclastic sediments from the Ossa– Morena Zone: Laser ablation U–Pb zircon data. In: M IRAO , J. & B ALBINO , A. (eds) Proceedings VII Congresso Nacional de Geologia, Estremoz, Universidade de Evora, Portugal, Vol. I, 21–24. P EREIRA , M. F., S ILVA , J. B., C HICHORRO , M., M OITA , P., S ANTOS , J. F., A PRAIZ , A. & R IBEIRO , C. 2007. Crustal growth and deformational processes in the Northern Gondwana margin: Constraints from the E´vora Massif (Ossa–Morena Zone, SW Iberia, Portugal). In: L INNEMANN , U., N ANCE , R. D., K RAFT , P. & Z ULAUF , G. (eds) The evolution of the Rheic Ocean: From Avalonian –Cadomian Active Margin to Alleghenian –Variscan Collision. Geological Society of America, Special Papers, 423, 333– 358. P IN , C., L IN˜ AN , E., P ASCUAL , E., D ONAIRE , T. & V ALENZUELA , A. 2002. Late Neoproterozoic crustal growth in the European Variscides: Nd isotope and geochemical evidence from the Sierra de Co´rdoba Andesites (Ossa–Morena Zone, Southern Spain). Tectonophysics, 352, 133– 151. P IQUE´ , A., C ORNE´ E , J.-J., M ULLER , J. & R OUSSEL , J. 1994. The Moroccan Hercynides. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 229–263. Q UESADA , C. 1990. Precambrian successions in SW Iberia: their relationship to ‘Cadomian’ orogenic events. In: D’L EMOS , R. S., S TRACHAN , R. A. & T OPLEY , G. G. (eds) The Cadomian Orogeny. Geological Society, London, Special Publications, 51, 353– 362. Q UESADA , C. 1991. Geological constraints on the Palaeozoic tectonic evolution of tectonostratigraphic terranes in the Iberian Massif. Tectonophysics, 185, 225–245. Q UESADA , C. & M UNHA´ , J. M. 1990. Ossa– Morena Zone: Metamorphism. In: D ALLMEYER , R. D. & M ARTINEZ G ARCIA , E. (eds) Pre-Mesozoic Geology of Iberia. Springer, Berlin, 314– 320. R OBARDET , M. 2002. Alternative approach to the Variscan Belt in southwestern Europe: Preorogenic palaeobiogeographical constraints. In: M ARTINEZ C ATALAN , J. R., H ATCHER , R. D., J R , A RENAS , R. & D IAZ G ARCIA , F. (eds) Variscan–Appalachian Dynamics: The Building of the Late Palaeozoic Basement. Geological Society of America, Special Papers, 364, 1–15. R OBARDET , M. 2003. The Armorica ‘microplate’: fact or fiction? Critical review of the concept and contradictory palaeobiogeographical data. Palaeogeography, Palaeoclimatology, Palaeoecology, 195, 125–148. S AMBRIDGE , M. S. & C OMPSTON , W. 1994. Mixture modelling of multi-component data sets with
408
M. F. PEREIRA ET AL.
application to ion-probe zircon ages. Earth and Planetary Science Letters, 128, 373–390. S ANCHEZ -C ARRETERO , R., C ARRACEDO , M., E GUILUZ , L., G ARROTE , A. & A PALATEGUI , O. 1989. El magmatismo calcoalcalino del Precambrico terminal en la Zona de Ossa–Morena (Macizo Ibe´rico). Revista de la Sociedad Geolo´gica Espana, 2, 7 –21. S ANCHEZ -G ARCIA , T., B ELLINDO , F. & Q UESADA , C. 2003. Geodynamic setting and geochemical signatures of Cambrian –Ordovician rift-related igneous rocks (Ossa–Morena Zone, SW Iberia). Tectonophysics, 365, 233–255. S CHA¨ FER , H.-J. 1990. Geochronological investigations in the Ossa–Morena Zone, SW Spain. PhD dissertation, ETH, Zurich. S CHA¨ FER , H. J., G EBAUER , D., N A¨ GLER , T. F. & E GUILUZ , L. 1993, Conventional and ion-microprobe U– Pb dating of detrital zircons of the Tentudia Group (Se´rie Negra, SW Spain): implications for zircon systematics, stratigraphy, tectonics and the Precambrian/ Cambrian boundary. Contributions to Mineralogy and Petrology, 113, 289– 299. S OLA´ , A. R., P EREIRA , M. F., R IBEIRO , M. L. ET AL . 2006. The ‘Urra Formation’: Age and Precambrian inherited record. In: M IRAO , J. & B ALBINO , A. (eds). Proceedings, VII Congresso Nacional de Geologia, Universidade de Evora, Estremoz, Portugal, Vol. I, 29–32. 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 Science Letters, 36, 359– 362. V IDAL , G., P ALACIOS , T., G AMEZ -V INTANED , J. A., D IEZ B ALDA , M. A. & G RANT , S. W. 1994. Neoproterozoic–early Cambrian geology and palaeontology of Iberia. Geological Magazine, 131, 729– 765. W ILLIAMS , I. S. 2001. Response of detrital zircon and monazite, and their U –Pb isotopic sytems, to regional metamorphism and host-rock partial melting, Cooma Complex, southeastern Australia. Australian Journal of Earth Sciences, 48, 557– 580. W ILLIAMS , I. S. & C LAESSON , S. 1987. Isotopic evidence for the Precambrian provenance and Caledonian metamorphism of high grade paragneisses from the Seve Nappes, Scandinavian Caledonides, II Ion microprobe zircon U– Th–Pb. Contributions to Mineralogy and Petrology, 97, 205– 217. W ILLIAMS , I., G OODGE , J., M YROW , P., B URKE , K. & K RAUS , J. 2002. Large scale sediment dispersal associated with the Late Neoproterozoic assembly of Gondwana. In: P REISS , V. P. (ed.) Geoscience 2002: Expanding Horizons. Abstracts of the 16th Australian Geological Convention, Adelaide, SA, 1– 5 July 2002, 67, 457. Z ECK , H. P., W INGATE , M. T. D., P OOLEY , G. D. & U GIDOS , J. M. 2003. A sequence of Pan-African and Hercynian events recorded in zircons from an orthogneiss from the Hercynian belt of western Central Iberia—an ion microprobe U–Pb study. Journal of Petrology, 45, 1613–1629.
Petrogenesis and geodynamic evolution of the Late Neoproterozoic post-collisional felsic magmatism in NE Afyon area, western central Turkey ¨ RSU1 & M. C. GONCUOGLU2 S. GU 1
Natural History Museum, Mineralogy– Petrography Division, MTA, 06520, Ankara, Turkey (e-mail:
[email protected]) 2
Department of Geological Engineering, Middle East Technical University, 06531, Ankara, Turkey
Abstract: In western Turkey, Late Neoproterozoic basement rocks are represented by variably deformed metasedimentary and meta-igneous rocks within different tectonostratigraphical units that make up the Alpine Tauride– Anatolide Platform. In the Ku¨tahya– Bolkar Dagı unit to the NE of Afyon this basement mainly includes garnet-bearing mica schists intruded by metamorphic granitic rocks with relict porphyritic textures. The youngest zircon ages obtained from the granitic rocks by the single zircon evaporation method are 542 + 5.0 Ma on average, which correlate with the Late Pan-African– Cadomian granitic magmatism. The granitic rocks are rhyodacitic or dacitic and peraluminous in composition, and display geochemical characteristics of I-type (tonalite–trondhjemite–granodiorite (TTG) source) felsic intrusive rocks. Trace and rare earth element patterns with distinct depletion in Rb, K, Nb, Sr, P and Ti relative to the other trace elements correlate very well with a Proterozoic TTG source. The petrogenetic modelling also implies that they were developed by partial melting of a TTG source by 20% fractional melting plus 20% Rayleigh fractional crystallization. The emplacement temperatures estimated by using zircon (790–820 8C), apatite and monazite saturation thermometry are about 827– 1035 8C; these are in accordance with I-type rather than S-type granite melts. A geochemical comparison of the NE Afyon granitic rocks with the coeval quartz-porphyries in the Sandikli area of the Geyik Dag tectonic unit suggests that the latter may represent the more evolved felsic part of the Cadomian magmatism. Hence, both basement complexes are parts of the same Gondwanan terrane and represent the eastern continuation of the North African–Southern European terrane assemblage.
A number of Precambrian terranes or microplates, named the Peri-Gondwanan terranes, are widespread in Southern Europe and the Eastern Mediterranean along the northern continental margin of West Gondwana (Amazonia and West Africa). They were separated from North Africa by the opening of several oceanic seaways known as Iapetus and Proto-, Palaeo- or Neotethys. The Tauride– Anatolide Platform in Turkey is one of the Cadomian-type Peri-Gondwanan terranes that formed along the West African margin by recycling of ancient West African crust (e.g. Murphy et al. 2002, 2004; Neubauer 2002) and had drifted from North Africa by the Early Mesozoic opening of the Bitlis– Zagros Neotethyan oceanic branch. The Alpine closure of this ocean resulted in a redistribution of the Late Neoproterozoic –Palaeozoic basement rocks and their Mesozoic cover within numerous allochthonous tectonostratigraphical units in eastern, central and southern Turkey. The Late Neoproterozoic basement rocks in these units are represented by diverse lithologies, including
high-grade metamorphic complexes (e.g. basement of Menderes Massif, Dora et al. 1995; Candan et al. 2001), very low-grade metavolcanic assemblages (Gu¨rsu & Goncuoglu 2001; Gu¨rsu et al. 2004a, b) and sedimentary–volcanic successions (Kozlu & Goncuoglu 1997; Goncuoglu et al. 1997). The radiometric age data from the felsic igneous rocks range around c. 541– 545 Ma, which correlates well with the Late Pan-African–Cadomian granitoids in North Africa and Gondwana-derived terranes in Southern and Central Europe (Balle`vre et al. 2001; Chantraine et al. 2001; El-Nisr et al. 2001; Pin et al. 2002; Bandres et al. 2002; Do¨rr et al. 2002; Genna et al. 2002; Mushkin et al. 2003) in terms of age and geochemistry. In the western central Taurides these basement rocks mainly occur within two main tectonic units (Fig. 1, distribution of Late Neoproterozoic basement rocks in western Anatolia). These are the Ku¨tahya–Bolkar Dagı (NE of Afyon) and Geyik Dag (SW of Afyon) units, as reported by Goncuoglu & Kozlu (2000).
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 409–431. DOI: 10.1144/SP297.19 0305-8719/08/$15.00 # The Geological Society of London 2008.
410
¨ RSU & M. C. GONCUOGLU S. GU
Fig. 1. (a) The main tectonic units of the Tauride–Anatolide Belt (modified after Goncuoglu et al. 1997); ¨ zcan et al. 1989) showing exposures of Neoproterozoic to Upper (b) geological map of NE Afyon (adapted from O Miocene rocks and location of samples that were collected for chemical and Pb– Pb analyses.
The Late Neoproterozoic low-grade metamorphic sediments and felsic igneous rocks in the Geyik Dag tectonic unit, slightly metamorphic rhyolites and granites, were described recently by Gu¨rsu et al. (2004a) and Gu¨rsu & Goncuoglu (2006). The Late Neoproterozoic basement rocks
in the Ku¨tahya–Bolkar Dagı tectonic unit, named the Afyon Basement Complex (ABC) (Gu¨rsu et al. 2003), differ from the former by being mainly composed of quartz–albite– garnet–biotite schists intruded by porphyritic granites. The geochemical, tectonic and petrogenetic significance of
FELSIC MAGMATISM, AFYON, TURKEY
these granitic rocks is not well known and has not been previously studied. Hence, the nature of this Late Neoproterozoic basement and the tectonomagmatic framework of the felsic rocks within the Gondwanan evolution are unclear. This paper presents new mineralogical, geochemical and geochronological data for Late Neoproterozoic porphyritic granitic rocks in the Ku¨tahya– Bolkar Dagı unit, with the aim of interpreting their petrology, petrogenesis and tectonic significance. We then address possible correlations with the geochemical characteristics of the better-known felsic intrusions in the Geyik Dag tectonic unit, which is established as a representative of the Cadomian granitic magmatism in Peri-Gondwana.
Geological setting and petrography The Afyon Basement Complex (ABC) as well as its Palaeozoic– Mesozoic cover in the Ku¨tahya – Bolkar Dagı tectonic unit to the NE of Afyon has different stratigraphical and metamorphic features from that in the Sandikli area of the Geyik Dag unit. In the area between Iscehisar –Bayat and Ihsaniye (Fig. 1) the ABC is mainly composed of quartz – albite –garnet –biotite schists, together with metamorphic porphyritic granitoids and micro-gabbroic stocks (Fig. 2). It is affected by extensive polyphase deformation and metamorphism and disconformably overlain by metaconglomerates and chlorite– quartz-schists alternating with laminated quartzites (Calıslar Formation of Gu¨rsu et al. 2003, 2004b). The formation resembles the Devonian successions in the Taurides and is unconformably overlain by the Permian Eldes Formation (Goncuoglu et al. 1992, 2003). The metaconglomerates of the Eldes formation include pebbles of quartz-schist, mica-schists and metagranitic rocks, and the sequence continues upwards with quartzites, calc-schists and fossiliferous recrystallized limestones (Fig. 2). Here, the almost 2 km thick Early Cambrian –Late Ordovician units of the Geyik Dag unit are missing. The Mesozoic cover starts with Lower Triassic clastic deposits, followed by a continuous platform-type carbonate deposition that lasted without a depositional break until the Cretaceous. The entire succession has been affected by HP –LT Alpine metamorphism ¨ zcan et al. 1989; Goncuoglu et al. 1992; (O Candan et al. 2003, 2005), which has partly obliterated the metamorphic mineral parageneses of the earlier metamorphisms. The country rocks of the ABC consist of metasedimentary rocks that are mainly composed of quartzo-feldspathic schists and garnet –mica
411
schists (Gu¨rsu et al. 2003) containing the assemblage quartz þ albite þ chlorite þ white mica + garnet + biotite. In the northern part of Afyon, the granitic dykes are more abundant than elsewhere. The vertical to subvertical dykes vary in length (from 50 m to 10 km) and in width (from 50 cm to 25 m). 207Pb – 206Pb zircon ages around 542 Ma from the intruding granitic rocks (in this study) indicate that the metasedimentary rocks are Precambrian in age. Based on the field observations and crosscutting relations (Fig. 3) the metabasic dykes represent younger intrusive events (Fig. 3a and b). They are micro-gabbroic dykes with the paragenesis albite + sodic amphibole + zoisite/epidote + garnet (Candan et al. 2005) (Fig. 3a and b). On outcrops near NE Afyon, the metapelites, granitic rocks and metabasic dykes were affected by low-grade metamorphism prior to the Alpine HP–LT event (Goncuoglu et al. 2003; Candan et al. 2003, 2005). They show welldeveloped metamorphic foliations and multiphase deformations, ascribed to the Cadomian, Late Variscan and Alpine events (e.g. Gu¨rsu et al. 2003, 2004b; Candan et al. 2005). The felsic rocks in the study area are characterized by metamorphosed porphyritic granitoids and show evidence of multi-phase deformation. They mainly include relict igneous minerals such as alkali feldspar (disordered orthoclase) and sodic plagioclase (oligoclase –andesine in composition) (Fig. 4a). Biotite is the predominant ferromagnesian mineral in the granitoids. Highly fractured orthoclase and plagioclase crystals are observed in thin sections (Fig. 4a). Syntectonic albite porphyroblasts with linear muscovite inclusions (Fig. 4b) that grew synchronously with the early stages of crenulation cleavage are sharply discordant with external schistosity formed during a later deformational phase (Fig. 4b). Recrystallized quartz þ fine-grained white micas (Fig. 4) grew synchronously with deformation under the low-grade metamorphic conditions during the Cadomian orogeny (Gu¨rsu et al. 2003). As a result of metamorphic overprinting, the original microstructures and textures in the granitic rocks were ultimately lost and fine- to coarse-grained metamorphic quartz and white mica neo-formations developed. Mafic minerals such as biotite are rarely preserved, and were converted to chlorite during retrograde metamorphism. Accessory phases are mainly composed of titanite, zircon, monazite, apatite and rarely opaque minerals. A discrete crenulation cleavage is well developed in granitic rocks, with an S1 fabric parallel to the main regional NNW–SSE foliation trend, overprinted by an S2 cleavage at nearly 908 to S1. S3 has overprinted and folded the S2 fabric diagonally at a low angle (308) and may have developed during the Alpine event (Fig. 4c).
412
¨ RSU & M. C. GONCUOGLU S. GU
Fig. 2. Generalized columnar section of the NE Afyon area (after Gu¨rsu et al. 2003, 2004b).
FELSIC MAGMATISM, AFYON, TURKEY
Fig. 3. Field occurrences of the intrusive rocks in the ABC: (a) and (b) metabasic dykes (MB) crosscut both garnet– mica schists (GMS) and granitic rocks (MG) of the study area; (c) the metamorphosed granitic rocks intrude the garnet–mica schists and show the development of metamorphic foliations and multiphase deformations.
Analytical methods Eight representative samples were analysed for major, trace and rare earth element (REE) abundances in Acme Laboratories, Vancouver, Canada. Determination of major element concentrations in rock samples expressed as common oxides for each element (SiO2, TiO2, Al2O3, Fe2O3, MnO,
413
Fig. 4. (a) Microphotograph of the granitic rocks of the ABC. The highly fractured orthoclase (Or) and plagioclase (Pl) porphyroclasts are surrounded by very fine-grained recrystallized quartz (Qtz) and neoformation of sericite (Ser) parallel to mylonitic fabric. (b) Syntectonic albite (Ab) porphyroblasts with linear inclusion fabrics that are mainly composed of fine-grained white mica occurrence (Ser) and are sharply discordant with external schistosity. (c) Deformation and development of S1, S2 and S3 planes in the granitic rocks of the ABC.
MgO, CaO, Na2O, K2O, P2O5 and Cr2O3), loss on ignition (LOI) and Sc were analysed by inductively coupled plasma atomic emission spectrometry
¨ RSU & M. C. GONCUOGLU S. GU
414
(ICP-AES) after fusion with lithium metaborate. The LOI was determined by weight difference after ignition at 1000 8C. Trace elements and REE were determined by inductively coupled plasma mass spectrometry (ICP-MS) following a lithium metaborate fusion and nitric acid digestion of 0.2 g samples. Analytical errors, as calculated from replicate analyses, are 0.5–1.0% for major elements and 0.5–3.5% for trace elements and REE. Radiometric age dating is based on the singlegrain zircon evaporation technique. The laboratory procedures were as described by Kober (1986). Zircons were separated by using a Wilfley table, magnetic separator, heavy liquids and finally handpicking. Isotopic measurements on the typical magmatic igneous zircons were analysed by the 207 Pb– 206Pb evaporation technique on a Finnigan MAT 262 mass spectrometer at the University of Tu¨bingen.
Single zircon Pb– Pb ages Our new data for the Ku¨tahya–Bolkar Dagı unit provide zircon ages for the granitic rocks. The calculated ages with the mean errors (2s) for a sample from the Late Neoproterozoic basement analysed using the single zircon evaporation method are presented in Table 1. The 207 Pb– 206Pb ages determined by this method are considered to be minimum ages. Sample 3 is a light grey granitic rock from Doganlar near Iscehisar (see Fig. 1). The zircons are clear, transparent, euhedral and long-prismatic with some inherited cores. Based on Pupin’s (1980) classifications, S17, S18 and S19 pyramidal terminations were the predominant zircon crystal types in the granitic rocks. Five grains were evaporated and all
produced different 207Pb – 206Pb ages ranging from 541+4 Ma to 2086+4 Ma (Table 1; Fig. 5a and b). The cathodoluminescence (CL) images of zircon crystals in the granitic rocks commonly have magmatic domains representing zircons from the protolith (Fig. 5c and d). Additionally, some inherited cores, and some alteration features such as flow domains, recrystallization and the zigzag-shaped sector zoning on small rounded cores are also observed (Fig. 5d). Although this type of sector zoning is typical for metapelitic rocks, it has reported been also from metaigneous rocks (e.g. Vavra et al. 1996). Two grains (Grains 1 and 2) yielded 207Pb – 206Pb ages of 541+4 Ma and 542+5 Ma with a mean of 542+5 Ma (Fig. 5a). We are inclined to consider the two youngest grains as reflecting the time of emplacement of the granitic protoliths, and the ages are similar to those of the quartz porphyry rocks (543+7 Ma, Kro¨ner & Sengo¨r 1990; 541.3+10.9 Ma, Gu¨rsu & Goncuoglu 2006) in the Sandikli area of the Geyik Dag unit. Additionally, similar ages have been obtained on the augen gneisses from the Menderes Massif (207Pb – 206Pb single zircon, 540–572 Ma, e.g. Hetzel & Reischmann 1996; Loos & Reischmann 1999; Koralay et al. 2004). The remaining three grains (Grains 3, 4 and 5) having brown, reddish, semi-transparent and well-rounded terminations provided xenocryst ages of 867+5 Ma, 2062+3 Ma and 2086 +4 Ma, respectively (Table 1; Fig. 5b). Three samples fall in the interval between 541 and 867 Ma with a peak at 542 Ma, whereas the oldest zircons show Palaeoproterozoic ages of 2.06– 2.08 Ga. The peak zircon age (542 Ma) correlates with that of the post-tectonic intrusions that followed the Cadomian arc-related magmatism in northern Anatolia, which is dated at 565–590 Ma in NW Anatolia as well as along
Table 1. Zircon morphology and Pb isotopic data from single grain evaporation Zircon morphology
Grain
Mass scans*
Evaporation temperature (8C)
C, E, LP, T C, E, LP, T
1 2 1, 2‡ 3 4 5
190 184 377 183 190 106
1380 1400
C, E, LP, T B, ST, WR R, ST, WR
1380 1400 1420
207
Pb – 206Pb ratio and 2s error†
0.058301 + 19 0.058319 + 24 0.05831 + 21.5 0.067953 + 26 0.127384 + 11.8 0.129081 + 20.5
207
Pb– 206Pb age and 2s error 541 + 4 542 + 5 542 + 5 867 + 5 2062 + 3 2086 + 4
(Sample 3, porphyritic metagranite) B, brown; C, colourless; E, euhedral; LP, long prismatic; R, reddish; SR, sub-rounded; ST, semi-transparent; T, transparent; WR, well-rounded. *Number of 207Pb – 206Pb ratios evaluated for age assessment. † Observed mean ratio corrected for non-radiogenic Pb where necessary. Standard errors based on uncertainties in counting statistics. ‡ Mean ages calculated for two grains.
FELSIC MAGMATISM, AFYON, TURKEY
415
Fig. 5. Histograms showing distribution of radiogenic lead isotope ratios derived from evaporation of zircon populations obtained from the granitic rocks of the ABC. (a) Graph for two grains, integrated from 374 ratios and interpreted to reflect age of protolith emplacement. (b) Graph for xenocrystic grains of granitic rocks. (c) CL images show single-phase magmatic zoning with small resorbed and recrystallized areas; arrows indicate regions of oscillatory zoning. (d) CL images showing typical internal structures. ic, rounded inherited cores; fl, flow domains; ft, sector zoning.
the northern margin of the Peri-Gondwana (e.g. Chen et al. 2002; Gu¨rsu & Goncuoglu 2005; Ustao¨mer et al. 2005). Arc-related magmatism and post-tectonic granitoid emplacement have also been observed in many Cadomian basements in Armorica and NW France (540 –746 Ma, Egal et al. 1996; Strachan et al. 1996; Balle`vre et al. 2001; Chantraine et al. 2001), Saxo-Thuringia, northern Germany (540– 660 Ma; Linnemann et al. 2000; Linnemann & Romer 2002; Mingram et al. 2004), the Iberian massif, southern Spain (540 to 575 –600 Ma, e.g. Bandres et al. 2002; Pin et al. 2002), Tepla –Barrandia (Bohemia), Czech Republic (550– 670 Ma, Do¨rr et al. 2002; Linnemann & Romer 2002; Patocˇka & Sˇtorch 2004) and the Carpathians, Romania
(567 –777 Ma, Lie´geois et al. 1996). The older ages obtained in this study are generally from brown –reddish, sub-rounded to well-rounded zircon grains that are interpreted as xenocrysts from the chronologically heterogeneous Palaeoproterozoic basement. These older zircon ages (2.06 –2.08 Ga) have been documented from the Eburnean basement of the West African craton (Alle`gre & Ben Othman 1980; Egal et al. 2002; Hirdes & Davis 2002; Linnemann et al. 2004; Murphy et al. 2004). The data presented here also support the conclusions of Goncuoglu et al. (1997) and Gu¨rsu & Goncuoglu (2006) that Tauride –Anatolite Platform was derived from the NW margin of Gondwana during the Late Neoproterozoic –Early Palaeozoic.
416
¨ RSU & M. C. GONCUOGLU S. GU
Geochemistry Major and trace element and REE element concentration of granitic rocks in the ABC are listed in Table 2. The rocks span a wide range of SiO2 contents (66.31 –73.68 wt%), are of subalkaline affinity and plot in the rhyodacite –dacite field on the Zr/TiO2 –Nb/Y discrimination diagram of Winchester & Floyd (1977) (Fig. 6a). High-silica rocks are subdivided into peralkaline, peraluminous and meta-aluminous groups on the basis of the alkali index (Hildreth 1981; Leat et al. 1986; Davies & MacDonald 1987; Kirstein et al. 2000). The granitic rocks show evidence of alkali mobility, which may affect the classification based on alkali elements. To avoid any misinterpretations for alkali indices, Nb–Zr variation diagrams (Leat et al. 1986) were used to determine the geochemical signatures of the granitic rocks. The log–log Nb v. Zr diagram shows a very scattered positive trend with ,326 ppm Zr and ,14 ppm Nb and the data plot close to the subalkaline fiew, characteristic of peraluminous magmas (Fig. 6b). Major element variations for the granitic rocks in the NE Afyon area are shown on selected Harker diagrams (Fig. 7). Al2O3 (11.57–15.56 wt%), Fe2O3 (3.71– 5.58 wt%), K2O (1.70 –3.99 wt%), MgO (1.79– 2.51 wt%) and TiO2 (0.64– 0.82 wt%) concentrations decrease with increasing SiO2 contents defining a nearly linear trend, and suggest that these rocks are probably products of magmatic differentiation from a parental magma (Fig. 7). K2O (1.70 –3.99 wt%) decreases along a hyperbolic curve, reaching near-one concentrations with increasing SiO2, which shows that the major part of the granitic rocks is high- to medium-K calc-alkaline. CaO (0.22–1.23 wt%) and Na2O (1.12–3.10 wt%) contents correlated very poorly with SiO2 as a result of the element mobility during the metamorphism (Fig. 7). The total alkali concentrations (Na2O þ K2O) in granitic rocks are uniform but have low values (4.13–6.06 wt%). The higher LOI (2.9–3.22 wt%), lower to high K2O/Na2O ratios (0.60–3.56), Na2O– SiO2, CaO–SiO2 and Sr–SiO2 variations indicate that the felsic rocks have been affected by alkali mobility related to the low-grade metamorphism and albitization. Consideration of major elements such as Na and K, which are commonly influenced by albitization, substantiates the outcome of the Na-metasomatism in the granitic rocks. This is in support of petrographical observations that potassium feldspar is affected by albitization, but relatively low Na2O/K2O ratios also indicate that some of original potassium feldspar is still preserved in the granitic rocks. Rb depletion also may have been a likely consequence of chloritization and argillitic alteration (breakdown of feldspar and mica). Sample 22 has
abnormally high silica content (73.68 wt%) and may reflect the locally silicification related to the secondary process (Fig. 7). Hence, SiO2, alkalis, Rb and Sr are not used in further geochemical diagrams to avoid any misinterpretation resulting from secondary processes. Variations of trace elements against SiO2 are plotted in Figure 7. Ba (316 –823 ppm), Rb (68 –140.9 ppm), Nb (9.3– 14.3 ppm), Th (9.3–16 ppm), V (69–105 ppm), Zr (181.8–326.5 ppm), Ce (46.7 –91.1 ppm), Y (19.9–30.7 ppm) and Lu (0.3–0.49 ppm) show well-defined negative trends with increasing SiO2 (with the exception of Sr (29.4– 84.3 ppm)), indicating the fractionation of plagioclase, alkali feldspar, biotite, titanite, zircon and iron oxides, respectively (Fig. 7). The relatively high values for large ion lithophile elements (LILE) and high field strength elements (HFSE) such as Nb (average 11.8 ppm), Rb (average 95 ppm), Y (average 26), Ce (average 67 ppm), Zr (average 251 ppm), Ta (0.8 ppm), and Yb (average 2.5 ppm), and the ratios Ti/Zr (15.5– 22.5 ppm), Ti/Y (135.5– 216.8 ppm), Nb/ Y (0.33–0.55 ppm) and Zr/Nb (16.5–28.3 ppm) in the granitic rocks are relatively consistent with average chemical composition of the Proterozoic tonalite–trondhjemite –granodiorite (TTG: 7.1, 63, 17.3, 45, 152, 0.72, 1.33, 18.5, 162.8, 0.41 and 21.4, respectively) (average data taken from Condie 2005). Chondrite-normalized (Sun & McDonough 1989) trace elements patterns of the granitic rocks are characterized by slightly enrichment in Ba, Th, La, Ce, Nd and Zr with depletion of Rb, K, Nb, Sr, P and Ti (Fig. 8a). The distinct Rb and Sr anomaly may also be ascribed to mobilization during the metamorphism. Chondrite-normalized REE patterns of granitic rocks show light REE (LREE) enrichment with respect to heavy REE (HREE), and negative Eu anomalies are moderately pronounced, with Eu/Eu* between 0.61 and 0.71 (Fig. 8a). The negative Eu anomaly in the chondrite-normalized REE plot suggests that plagioclase and alkali feldspar fractionation within the stability field of feldspar has played a significant role in formation of these granitic rocks and indicates intracrustal fractionation. The granitic rocks have moderate to high REE contents P ( REE 99.8 –196.8 ppm) with fractionated (La/Yb)N ¼ 4.6 –10 and (La/Sm)N ¼ 2.6–3.7, and relatively flat (Gd/Yb)N ¼ 1.02 –1.85. The LILE and REE chondrite-normalized patterns display very similar patterns to the average chemical composition of Proterozoic TTG (Fig. 8b). The continuous negative trends of granitic rocks in Nb/U –U and Ce/Pb –Pb variations indicate that U and Pb may have been added as a result of intracrustal melting (Fig. 8c and d), and the low Nb/U (,10) and variable Ce/Pb (,65) ratios
FELSIC MAGMATISM, AFYON, TURKEY
417
Table 2. Representative chemical analyses of granitic rocks in the ABC to the NE of Afyon in the Ku¨tahya–Bolkar Dagı unit Sample no. SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 Cr2O3 LOI Total Ba Sc Co Pb Zn Ni Cs Ga Hf Nb Rb Sr Ta Th U V W Zr Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Ti Zr/Y Th/Nb Th/Y Y/Nb La/Nb Ti/Y (La/Yb)N (La/Sm)N (Gd/Yb)N Eu/Eu*
Afyon3
Afyon6
Afyon9
Afyon11 Afyon22 Afyon177 Afyon179 Afyon187 Average
68.52 66.87 70.29 68.61 73.68 67.90 69.11 66.31 68.91 0.69 0.65 0.72 0.82 0.64 0.69 0.77 0.81 0.72 14.49 15.55 13.71 13.74 11.57 14.48 13.67 15.56 14.09 4.99 5.58 4.56 5.56 3.71 4.46 5.35 5.57 4.97 0.05 0.05 0.05 0.08 0.04 0.04 0.09 0.08 0.06 2.21 1.80 1.79 2.51 1.81 2.15 2.0 2.47 2.09 0.53 0.22 0.44 0.60 0.36 1.23 0.39 0.41 0.52 1.12 2.68 2.97 2.10 2.38 3.09 2.81 3.10 2.28 3.99 3.38 2.01 2.70 1.75 2.54 1.70 2.45 2.56 0.21 0.18 0.23 0.20 0.24 0.24 0.21 0.22 0.21 0.019 0.009 0.015 0.014 0.014 0.011 0.012 0.013 0.013 3.1 3.2 3.0 2.9 3.8 3.1 3.6 3.1 3.2 100.02 100.25 99.84 99.93 100.04 100.01 99.77 100.17 100.00 823 688 420 728 316 658 483 622 592 13 19 12 12 9 12 12 14 13 13 8.4 9.4 13.4 9.8 11.4 14.9 12.6 11.6 10 14.8 4 1.4 13.8 4.5 2.2 22.3 9.1 70 67 61 80 85 56 75 86 72 52.5 18.5 31.2 43.2 32.4 27.6 36.9 38.3 35.1 4 2.4 2.4 4 2.7 3.1 2.7 2.6 3.0 20.7 21.4 16.8 18.2 14.4 18 17.4 20.1 18.4 6.2 7.4 6.7 9.7 7.9 7.4 7.6 6.2 7.4 11.6 12.2 11 14.3 9.3 11.1 11.9 13.1 11.8 140.7 120.3 81.7 99.6 68 87.8 68.7 89.7 94.6 80.9 33.2 84.3 58.7 41.7 29.4 63.6 49.9 55.2 0.80 0.80 0.80 0.90 0.7 0.8 0.8 1.0 0.8 12.4 12.1 11.7 16 9.3 12.2 10.4 11.2 11.9 2.4 1.4 1.9 3.0 2.1 2.3 1.7 1.8 2.1 98 105 93 94 69 89 93 97 92 1.7 1.2 0.6 1.2 1.0 0.8 1.7 1.1 1.2 191.8 251.6 246 326.5 263 231 279.4 216.1 250.6 21.2 27.4 19.9 25 28.3 28.4 30.2 30.7 26.4 15.9 18.3 20.9 38.6 32.3 31.1 29.7 37.1 28.0 46.7 62.1 56.6 91.1 61.0 76.9 61.8 75.9 66.5 3.72 4.05 4.75 7.87 6.60 6.74 6.22 7.72 5.95 16 17.8 19.1 33.8 27.2 28 25.8 31.8 24.9 3.8 3.8 3.9 6.6 5.7 5.8 5.3 6.6 5.2 0.78 0.76 0.80 1.20 1.39 1.30 1.17 1.40 1.10 3.26 3.37 3.22 5.38 5.63 5.71 5.31 5.57 4.68 0.64 0.63 0.53 0.86 0.92 0.89 0.82 0.86 0.76 3.30 4.18 2.81 4.53 4.74 4.91 4.76 5.0 4.27 0.66 0.87 0.56 0.83 0.89 0.86 0.92 0.89 0.81 2.24 2.76 1.94 2.67 2.86 2.95 2.97 2.95 2.66 0.28 0.35 0.22 0.35 0.31 0.41 0.40 0.38 0.33 2.21 2.69 1.84 2.61 2.47 2.71 2.68 2.89 2.51 0.36 0.44 0.30 0.44 0.38 0.43 0.46 0.49 0.41 4135 3896 4315 4915 3836 4135 4615 4855 4338 9.04 9.18 12.36 13.06 9.29 8.13 9.25 7.04 9.66 1.07 0.99 1.06 1.12 1.00 1.09 0.87 0.85 1.00 0.58 0.44 0.58 0.64 0.33 0.43 0.34 0.36 0.46 1.82 2.24 1.81 1.74 3.04 2.56 2.53 2.34 2.26 1.37 1.50 1.90 2.69 3.47 2.80 2.49 2.83 2.38 195.08 142.2 216.8 196.6 135.55 145.6 152.8 158.1 167.8 4.86 4.60 7.68 9.99 8.84 7.75 7.49 8.67 7.48 2.63 3.03 3.37 3.68 3.57 3.38 3.53 3.54 3.34 1.20 1.02 1.42 1.67 1.85 1.71 1.61 1.56 1.50 0.68 0.65 0.71 0.61 0.75 0.69 0.67 0.70 0.68
Subscript N indicates normalization data from Sun & McDonough (1989).
418
¨ RSU & M. C. GONCUOGLU S. GU
saturation thermometry (Table 3). Petrographic and geochemical signatures of the granitic rocks from the study area show that the felsic magmas must have been saturated in zircon, monazite and apatite minerals. Zircon, monazite and apatite saturation temperatures can be calculated from wholerock geochemical data to estimate temperatures and composition effects of crustal magma types by using the experimental models of Watson & Harrison (1983), Harrison & Watson (1984), Montel (1993) and Piccoli et al. (1999). Hydrothermal experiments in the temperature range of 750–1020 8C show the saturation behaviour of zircon in crustal anatectic melts as a function of both temperature and the zircon solubility range in peraluminous granites changing from c. 100 ppm dissolved at 750 8C c. 1330 ppm at 1020 8C (Watson & Harrison 1983). The zircon solubility model of Watson & Harrison (1983) is given by equation ln DZr ¼ f3:80 ½0:85 ðM 1Þg þ 12900=T
Fig. 6. (a) Zr/TiO2 v. Nb/Y diagram (Winchester & Floyd 1977) and (b) Nb– Zr diagram (Leat et al. 1986) for granitic rocks of the ABC; O, granitic rocks.
suggest magmatic differentiations of intracrustal melting of a TTG source that is in the garnet stability field under the mantle wedge (Martin et al. 2005; Condie 2005). The low to very low unfractionated HREE and Y abundances suggest that the granitic rocks were evolved from second-stage melts of a TTG source during intracrustal melting without residual garnet. To summarize, the geochemical variations confirm that the granitic rocks are sub-alkaline (peraluminous) in composition. They have geochemical characteristics of I-type (TTG source) felsic intrusive rocks and magmatic differentiations such as fractional crystallization trends are observed in geochemical diagrams. Trace element and REE patterns show similar patterns, with distinctive depletion in Rb, K, Nb, Sr, P and Ti relative to the other trace elements. Trace element and REE patterns correlate very well with the Proterozoic TTG data.
Magmatic temperatures, crystallization and source rock conditions In this study, emplacement temperatures were estimated by using zircon, monazite and apatite
where T is zircon saturation temperature (in 8C), DZr is the bulk concentration of Zr and M is the cationic ratio [(Na þ K þ 2 Ca)/(Al Si)] of the whole-rock concentration of SiO2, Al2O3, NaO, K2O, CaO. The Zr concentration and cationic ratio (M) show similar scatter in granitic rocks (192 –327 ppm; 0.70 –1.05). Zircon saturation temperatures of granitic rocks have a distinct peak at 790–809 8C with only one higher value at 820 8C (average 803 8C) (Table 3). Harrison & Watson (1984) showed that the whole-rock concentrations of SiO2 and P2O5 of granitic rocks were equivalent to the initial melt composition of apatite crystallized from the melt. Apatite precipitation in peralkaline, meta-aluminous, slightly peraluminous and highly peraluminous silicate melts (excluding low-Ca melts) can be approximated as a function of P2O5 and SiO2 (Harrison & Watson 1984; Pichavant et al. 1992). Apatite saturation temperatures at the temperature at which apatite began to crystallize from the magmas were calculated by using SiO2 and P2O5 in the following equation of Piccoli et al. (1999), which is adapted from Harrison & Watson (1984): T ¼ ð26400 C1 SiO2 4800Þ= ð12:4 C1 SiO2 lnðC1 P2 O5 Þ 3:97Þ 273:15 where T is the apatite saturation temperature (in 8C), and C1SiO2 and C1P2O5 are the concentration of silica and phosphorus (expressed as weight fractions in the melt at the apatite
FELSIC MAGMATISM, AFYON, TURKEY
419
Fig. 7. Major and selected trace element variation diagrams v. SiO2 (field boundaries for K2O, I- and A-type granites are from Peccerillo & Taylor (1976) and Collins et al. (1982), respectively); W, granitic rocks; R, quartz porphyry rocks of Geyik Dag Unit in Sandikli, SW Afyon (data taken from Gu¨rsu & Goncuoglu 2006).
¨ RSU & M. C. GONCUOGLU S. GU
420
Fig. 8. (a) Chondrite-normalized trace multi-element diagram; (b) chondrite-normalized REE patterns (all normalizing values from Sun & McDonough 1989). Shaded area for quartz porphyry rocks is from Gu¨rsu & Goncuoglu (2006). (c) and (d) Nb/U– U and Ce/Pb–Pb diagrams. O, granitic rocks; A, upper crust; N, lower crust; O, N-MORB (mid-ocean ridge basalt); half-filled triangles, enriched mid-ocean ridge basalt (E-MORB); V, ocean-island basalt (OIB); B, TTG. Upper and lower continental crust, N-MORB, E-MORB, OIB and TTG data are from Taylor & McLennnan (1995), Sun & McDonough (1989) and Condie (2005), respectively.
crystallization temperature). Piccoli & Candela (1994) and Piccoli et al. (1999) showed that the relationship between the SiO2 and P2O5 bulk concentrations in the initial melt can be used for the apatite crystallization temperature at crustal
pressures. This equation is based on the kinetic studies on the solubility of apatite in felsic melts by Harrison & Watson (1984), and is valid at crustal pressures for rocks with 45 –75 wt% SiO2 and ,10% H2O (Harrison & Watson 1984). The
Table 3. The zircon, monazite and apatite saturation temperatures of granitic rocks in the ABC Sample no.
Zircon saturation temperature (8C)
Monazite saturation temperature (8C)
Apatite saturation temperature (8C)
Afyon3 Afyon6 Afyon9 Afyon11 Afyon22 Afyon177 Afyon179 Afyon187 Average
820 802 803 790 804 797 799 809 803
833 836 830 886 860 827 862 873 851
970 935 998 965 1035 979 976 952 976
FELSIC MAGMATISM, AFYON, TURKEY
average SiO2 and P2O5 weight concentrations of granitic rocks are distinguished by being lower 64.5– 73.4% (average 68.9%) and higher 0.18– 0.24% (average 0.24%), respectively. The estimates of apatite saturation temperatures of granitic rocks have a distinct peak at 935 –979 8C with only two higher values at 998– 1035 8C (average 976 8C) (Table 3). As shown by petrography, monazite is the main LREE-bearing mineral in less evolved granitic rocks. In contrast to zircon, the monazite solubility is strongly dependent on the H2O contents in Ca-poor felsic magmas (Montel 1993) and is described in the following equation of Montel (1993): p lnðREEt Þ ¼ 9:50 þ 2:34D þ 0:3879 H2 O 13318=T 273:15 P where REEt ¼ [REE (ppm)/atomic weight (g mol21)], D is the cationic % (Na þ K þ Li þ 2Ca)/Al(Al þ Si), H2O is in wt% and T is the monazite saturation temperature (in 8C). The REE considered extend from La to Gd, excluding Eu (Montel 1993). We calculated temperatures for 3.5 wt% H2O, which is a typical value expected in crustal magmas (Thomson 1996). Estimated monazite saturation temperatures of granitic rocks have a distinct peak at 827 –836 8C with higher values at 860 –886 8C (average 851 8C); (Table 3). The monazite saturation temperatures are compatible with zircon data rather than apatite data for the studied granitic rocks from the ABC. The calculated apatite saturation temperatures are higher than those obtained from the zircon and monazite saturation thermometry by 100 –180 8C. To justify this discrepancy, Pichavant et al. (1992) noted that the calibration of Harrison & Watson (1984) fails to yield reliable results in peraluminous rocks, and all the granitic rock samples from the ABC contain low CaO (about ,1.0 wt%). Additionally, Winchester & Floyd (1976) found that P could easily show mobility as a result of progressive alteration and metamorphism, and this may affect the apatite saturation thermometry. In a source rock without monazite, apatite dissolution will release both LREE and phosphorus into the melt and if REE in melt become oversaturated, monazite crystallization occur (Zeng et al. 2005). The mineralogical–petrographical analyses of the granitic rocks show that zircon and monazite crystallized early. They occur as small inclusions within the mica minerals with grain sizes of about 0.024 –0.045 mm and 0.032– 0.152 mm, respectively. The Pb –Pb systematics of the zircon from the granitic rocks in the ABC indicates the presence
421
of an inherited older crustal component, and inherited zircons were partly dissolved in the melt. The zircon temperatures may be suggestive of a high proportion of assimilated crustal material in the magmas. Therefore zircon saturation temperatures calculated from total Zr abundances may be higher than the real temperatures of magmas. We obtained slightly higher temperatures (790 –820 8C) for the granitic rocks of the Neoproterozoic basement by using zircon thermometry. Our average estimates of zircon, monazite and apatite saturation temperatures in the felsic magmas range from 803 to 976 8C. The calculations provide a reasonable estimate of the temperatures over which zircon and monazite crystallized. On the other hand, based on Pupin’s (1980) classifications, S17, S18 and S19 pyramidal terminations (see Table 3) in the studied granitic rocks were the predominant zircon crystal types, indicating temperatures of 800 + 50 8C. These results suggest that the felsic magma temperatures were between c. 800 8C and c. 850 8C in the studied granitic rocks of the ABC and could be interpreted as temperatures of LREE and Zr saturation at the onset of monazite and zircon crystallization. The generation of I-type granite melts requires substantially higher temperatures (800 –900 8C) than that of S-type granitoid melts (700 8C), as shown by experimental work (e.g. Liew & McCulloch 1985). Estimated magma temperatures of the granitic rocks in this study of about 850 8C are in agreement with I-type rather than S-type granite melts. Nb –SiO2 and Y –SiO2 variation diagrams (see Fig. 7) also indicate an I-type source (Collins et al. 1982) rather than S- and A-type granitic rocks. In addition to these data, Al2O3 –SiO2 variations also identify the compositional differences of experimental melts produced by partial melting of various source rocks (Helz 1976; Spulber & Rutherford 1983; Beard & Lofgren 1989, 1991; Winther & Newton 1991; Wolf & Wyllie 1994). The protoliths of granitic rocks of the ABC are probably homogeneous and may have originated from crust-derived (probably TTG-like compositions with negative Eu and flat HREE) felsic magmas and may have been generated from a chemically similar magma source.
Magma modelling and the generation of felsic magmas The decrease in Zr abundance with increasing SiO2 suggests that zircon was a fractionating phase (e.g. Watson & Harrison 1983). The multielement behaviour is characterized by lower
422
¨ RSU & M. C. GONCUOGLU S. GU
abundances of Rb, K, Nb, Sr, P and Ti and higher abundances of Ba, Th, La, Ce, Nd and Zr suggesting a source similar to TTG. The flat HREE patterns in the metagranitic rocks show that garnet was not restite phase, which indicates that the felsic magmas were formed at pressures ,10 kbar (Rutter & Wyllie 1988; Vielzeuf & Montel 1992); this implies crustal depths of melting. The granitic rocks have low silica (66.3– 73.6 wt%) and relatively low to high alumina (11.5–15.5 wt%) contents consistent with partial melting of an igneous source (average 67.3 wt% SiO2 and average 15.8 wt% Al2O3 in TTG; average 73.16 wt% SiO2 and 13.90 wt% Al2O3 in granites; 73.12 wt% SiO2 and 13.82 wt% Al2O3 in felsic volcanic rocks, respectively) rather than a sedimentary protolith (65.83 wt% SiO2 and 15.33 wt% Al2O3 in greywackes; 63.50 wt% SiO2 and 17.72 wt% Al2O3 in cratonic shales (e.g. Condie 1993, 2005; Chappell 1999). The trace element and REE pattern and variation diagrams (see Fig. 8) indicate that the source was igneous with Ti/Zr (average 20.97), Nb/Y (average 0.45), Zr/Nb (average 21.42), Th/Y (average 0.46) and La/Nb (average 2.38) ratios similar to those of TTG (18.5, 0.41, 21.4, 0.35 and 3.66, respectively; data from Condie (2005)). All these data support the idea that the granitic rocks in the ABC may be derived from dehydration melting of a TTG source rather than from sedimentary protolith. Experimental data also support that dehydration of tonalites may produce granitic melts within the temperature range 750– 900 8C at 2–10 kbar pressure (Sing & Johannes 1996a, b), and peraluminous tonalites can contain up to about 30% biotite þ muscovite and quartz þ plagioclase concentrations, which suggests remelting to yield mobile granitic magmas (Patin˜o & Patin˜o-Douce 1987). The TTG source composition (Condie 2005) was used as the possible source rock (C0), and fractional melting and Rayleigh fractional crystallization process were modelled by using Rb, Ba, K, Sr, Nb, Th, U, Zr, Ti, Y, La, Ce, Nd, Sm, Eu, Gd, Tb, Dy, Er, Yb and Lu to constrain the generation of granitic rocks. The phase proportions were estimated from the normative mineralogy of possible source rocks and partition coefficient values (D0) were calculated assuming 51% plagioclase, 21.5% quartz, 11.5% alkali feldspar, 11% amphibole and 5% biotite in the source residue with mineral –melt coefficient for dacitic –rhyolitic magma compositions (Arth 1976; Pearce & Norry 1979; Henderson 1982; Watson & Harrison 1983; Nash & Crecraft 1985; Table 4). The fractional melting and Rayleigh crystallization were used to determine the approximate minor and trace element, and REE compositions. Bulk D10 values
for the Rayleigh crystallization modelling were calculated from average normative mineralogical contents of granitic rocks considered as an assemblage of 43.7% quartz, 21.4% plagioclase, 15.7% alkali feldspar, 7.3% corundum, 5.4% orthopyroxene, 5.2% magnetite and 0.48% apatite. Calculated 20% fractional melting plus 20% fractional crystallization of TTG source rocks gives similar trends to those of granitic rocks (Table 5, Fig. 9a and b). Experimental data (e.g. Clemens & Vielzeuf 1987; Rutter & Wyllie 1988; Sing & Johannes 1996a, b) also show that melt proportions obtained by dehydration melting of tonalites may increase to nearly 15 vol% at 850 8C and 22.5 vol% at 900 8C; these values are similar to the 20% partial melting of a TTG source in our modelling study. Chondrite-normalized calculated compatible/ incompatible element patterns, shown in Figure 9a, are very similar to those of the studied granitic rocks and display enrichment in Ba, La, Ce, Nd, Zr and Y and depletion in Rb, K, Nb, Sr and Ti. Sr and Ti negative anomalies are consistent with significant fractional crystallization. The REE pattern of the calculated TTG melts also shows a pattern concordant with that of the granitic rocks and displays an enrichment of LREE rather than HREE and a pronounced negative Eu anomaly (Fig. 9b). A 20% fractional melting and 20% fractional crystallization of a Proterozoic TTG source provides a significant control on the change in Eu/Eu* with increasing Th abundances for the studied granitic rocks (Fig. 10). In conclusion, we suggest that the felsic magmas in western Anatolia were generated by dehydration melting of a TTG source at ,10 kbar pressure. Our petrogenetic modelling also implies that the granitic rocks were developed from partial melting of a TTG source by 20% fractional melting plus 20% Rayleigh fractional crystallization.
Comparison of the Neoproterozoic granitic magmatism in the basements of Ku¨tahya– Bolkar Dagı and Geyik Dag tectonic units The metaquartz porphyry rocks in the Geyik Dag unit in the Sandikli area display petrographical and chemical differences from the granitic rocks of the Ku¨tahya–Bolkar Dagı unit in the NE Afyon area. The zircon and apatite thermometers of metaquartz porphyry rocks in Sandikli yielded two distinct peaks at 783–811 8C and 821– 845 8C (average 816 8C) and three distinct peaks at 785–823, 881–925 and 947–998 8C (average 912 8C), respectively, and include well-developed perthitic textures (Table 6). The phenocryst phase
FELSIC MAGMATISM, AFYON, TURKEY
423
Table 4. Mineral–melt partition coefficients for dacitic –rhyolitic melts used in the modelling of fractional melting and Rayleigh fractional crystallization Distribution coefficients
Quartz K-feldspar Plagioclase Biotite Hornblende Orthopyroxene Apatite Magnetite melt*† melt† melt† melt* melt* melt* melt* melt†
Rb Ba Th U K Nb Sr Zr Ti Y La Ce Nd Sm Eu Gd Tb Dy Er Yb Lu
0.041 0.022 0.009 0.025 0.013
1.75 6.12 0.023 0.048 1.490 3.87 0.03
0.038 0.015 0.014 0.016 0.014 0.056
0.08 0.037 0.035 0.025 4.45
0.017 0.015
0.025 0.055 0.006 0.03 0.033
0.017 0.014
0.105 1.515 0.048 0.093 0.1 0.06 4.4 0.135 0.05 0.13 0.38 0.267 0.203 0.165 1.214* 0.125 0.112 0.055 0.09 0.092
3.20 6.36 1.227 0.167 5.63 6.367 0.447 1.197 1.233 5.713 4.357 2.56 2.117 0.87* 0.067 1.957 1.72 0.35 1.473 1.617
0.014 0.044 0.081 4.0 0.022 4.0 7.0 6.0 1.52 4.26 7.77 5.14 10.0 13.0 12.0 8.39 5.5
0.003 0.003 0.13 0.145 0.02 0.8 0.009 0.2 0.4 1.0 0.78 0.93 1.25 1.6 0.825 0.34 1.85 1.8 0.65 2.2 2.25
0.1 0.1 0.1 40 14.5 34.7 57.1 62.8 30.4 56.3
0.8 12.5 2.0
50.7 37.2 23.9 20.2
*Nash & Crecraft (1985). † Arth (1976) and Pearce & Norry (1979). ‡ Henderson (1982).
Table 5. Parameters used in the modelling of 20% fractional melting plus 20% Rayleigh fractional crystallization of a TTG source to produce granitic rocks of the ABC in NE Afyon Trace element Rb Ba Th U K Nb Sr Zr Ti Y La Ce Nd Sm Eu Gd Tb Dy Er Yb Lu
D0
C0
20% fractional melting
D10
C10
20% fractional crystallization
Average of metagranitic rocks
0.425 1.804 0.090 0.066 0.515 0.789 9.195 0.572 0.804 0.788 0.492 0.528 0.707 1.050 1.752 1.177 0.104 0.835 0.699 0.613 0.569
63 717 6.1 2.1 19093 7.1 473 152 2817 17.3 26 45 18 3.50 0.95 3.0 0.49 3.50 1.90 1.33 0.23
110 439 7.14 1.38 30029 8.47 62.8 225 3319 20.7 42.0 70.0 23.2 3.37 0.59 2.63 0.69 4.01 2.46 1.88 0.34
0.315 1.295 0.025 0.046 0.261 0.200 1.549 0.089 0.700 0.378 0.212 0.243 0.341 0.361 1.338 0.315 0.111 0.307 0.156 0.151 0.219
110 439 7.14 1.38 30029 8.47 62.8 225 3319 20.7 42.0 70.0 23.2 3.37 0.59 2.63 0.69 4.01 2.46 1.88 0.34
128 411 8.87 1.70 35411 10.12 55.5 275 3549 27.7 50.0 82.6 26.9 3.88 0.58 3.06 0.84 4.68 2.96 2.27 0.40
94.6 592 11.91 2.07 21252 11.81 55.2 251 4338 26.4 28.0 66.5 24.9 5.18 1.10 4.68 0.76 4.27 2.66 2.51 0.41
424
¨ RSU & M. C. GONCUOGLU S. GU
Fig. 9. (a) and (b) trace element and REE diagrams for the granitic rocks and their comparison with the calculated composition for 20% fractional melting plus 20% Rayleigh fractional crystallization of a TTG source. W, granitic rocks; V, granitic rocks obtained from 20% fractional melting plus 20% Rayleigh fractional crystallization of the TTG source of Condie (2005).
includes quartz, microperthitic K-feldspar and microcline, with accessory minerals such as titanite, allanite, apatite and zircon. The granitic rocks in the latter locality contain alkali feldspar (disordered orthoclase), plagioclase (oligoclase –andesine in composition), quartz and biotite as relict phases. The accessory minerals are titanite, zircon, monazite, apatite and, rarely, opaque minerals. Geochemically, the quartz porphyry rocks of Sandikli (Gu¨rsu & Goncuoglu 2006) have higher SiO2 (73.56 –77.87 wt%), K2O (4.73–7.81 wt%) and TiO2 (0.16–0.48 wt%) and lower Al2O3 (11.68–13.68 wt%), Fe2O3 (1.05 –3.38 wt%), P2O5 (0.02 –0.13 wt%) and MgO (0.21 –1.1 wt%)
than the granitic rocks of NE Afyon (see Fig. 7). They can be clearly distinguished from the latter also by lower Y/Nb (2–5 ppm)Pand Ti/Zr (,11), higher REE abundances ( REE 220 ppm), lower Eu/Eu* ¼ 0.16 –0.33, highly fractionated (La/Yb)N ¼ 2.6 –13.8 and (La/Sm)N ¼ 2.2–4.2, and relatively flat (Gd/Yb)N ¼ 0.9–1.9 (for a brief review see Gu¨rsu & Goncuoglu 2006). The granitic rocks, on the other hand, show relatively less enriched LREE and more depleted HREE patterns (see Fig. 8a and b). The correlative histogram
Table 6. Zircon and apatite saturation temperatures of the quartz porphyry rocks of the Geyikdag˘ unit in SW Afyon Sample no.
Fig. 10. Eu/Eu*– Th diagram (Kirstein et al. 2000) with calculated model curves for closed-system fractional crystallization. Modal curve was generated from 20% fractional partial melting of starting composition of TTG source of Condie (2005), where amounts of fractional crystallization (0.1– 75%) are indicated.
7 176 192 198 199 301 336 337A 344 664 692 740 744 745 944 1058 Average
Zircon saturation temperature (8C)
Apatite saturation temperature (8C)
781 792 833 845 838 821 791 783 804 844 811 807 832 839 807 829 816
919 899 823 905 907 998 873 925 881 947 970 959 952 975 886 785 912
The whole-rock analyses of the samples are from Gu¨rsu & Goncuoglu (2006).
FELSIC MAGMATISM, AFYON, TURKEY
diagrams also indicate that on average the granitic rocks in the ABC of NE Afyon display Zr, Hf, Eu and Ti enrichments and Th, Nb, La, Ce, Nd, Sm, Gd, Tb, Dy, Y, Ho, Er, Tm, Ty and Lu depletions P (Fig. 11a). Moreover, the Ti/Zr– Nb/Y, REE– (La/Yb)N and Eu/Eu*–(La/Yb)N variation diagrams clearly reflect the presence of two chemically different felsic rock groups in the Late Neoproterozoic basements of the Geyik Dag and Ku¨tahyaBolkar Dagı units (Fig. 11b –d). Petrographical and geochemical characteristics suggest that the metaquartz porphyry rocks in Sandikli may represent the more felsic part of the coeval granitic magmatism that produced the granitic rocks in the NE Afyon area. Both rock types were formed from a TTG-type source by 20% fractional melting and 20% Rayleigh fractional crystallization (for a brief review see Gu¨rsu & Goncuoglu 2006) in the upper continental crust. A comparison of the Neoproterozoic felsic magmatism in the Ku¨tahya–Bolkar Dagı (NE Afyon) and Geyik Dag (Sandikli) units in the Western Taurides is shown in Table 7.
425
Geodynamic implications and conclusions The granitic rocks within the Late Neoproterozoic basement of the Taurides were ascribed to a postcollisional extension in the North Gondwanan margin (Goncuoglu 1996; Goncuoglu & Kozlu 2000; Gu¨rsu et al. 2004a; Gu¨rsu & Goncuoglu 2006). This event very probably followed the southward subduction of the oceanic lithosphere beneath the Gondwanan continental crust (e.g. see Gu¨rsu & Goncuoglu 2005, fig. 13) that formed the 590 and 570 Ma arc-type granitoids intruding an older (890 –710 Ma, Chen et al. 2002) sector of continental crust in the Safranbolu area of northern Turkey. The early Late Neoproterozoic rocks in northern Turkey were attributed to North Gondwana, are of tonalitic and granitic composition, and have low Nb/Y ratios and Ti contents, consistent with those of arc rocks (Chen et al. 2002). This tonalitic basement may be the TTG-type source of the Late Neoproterozoic granitic rocks in the Afyon and Sandikli areas. An alternative source for the studied granitoids may be the Mid-Proterozoic
P Fig. 11. (a) Correlative, (b) Ti/Zr –Nb/Y, (c) REE– (La/Yb)N (from Zhao et al. 2004), (d) Eu/Eu*–(La/Yb)N variation diagrams indicating the differences between the granitic rocks in the Ku¨tahya–Bolkar Dagı unit of NE Afyon (ABC) and quartz porphyry rocks of the Geyik Dag Unit of Sandikli, SW Afyon.
426
Table 7. A comparison of the Neoproterozoic felsic magmatism in the Ku¨tahya– Bolkar Dagı (NE Afyon) and Geyik Dag˘ (Sandıklı) units in the Western Taurides Feature Stratigraphy Type of metamorphism
Mineral contents Geochemical fingerprints
Melting process Source P – T conditions during melting Timing of emplacement (zircon ages) Associated deformation Geodynamic model
Geyik Dag˘ (Sandıklı) Unit
The Late Neoproterozoic basement is unconformably overlain by Late Palaeozoic units and the Early Cambrian – Late Ordovician units are missing Multiple phases of tectonothermal events were observed as greenschist-facies (first), HP/LT blueschist-facies (350 8C and 6–9 kbar pressure, 30 km burial depths) (second) and greenschist-facies (third) overprints (Go¨ncu¨oglu et al. 2001; Candan et al. 2005)
The Late Neoproterozoic basement is unconformably overlain by Early Cambrian – Late Ordovician units. The Late Palaeozoic units are missing Multiple phases of tectonothermal events were observed as pre-Early Cambrian low-grade dynamothermal metamorphism (c. 300 8C, 4.2 kbar pressure, c. 15 km burial depths) and Late Palaeozoic very low grade metamorphism (3.2 kbar pressure, c. 10 km burial depth)
Metamorphosed leucogranitic rocks and sills or dykes of porphyry rocks Alkali feldspar, plagioclase (oligoclase – andesine in composition), biotite, titanite, zircon, monazite, apatite and, rarely, opaque minerals Sub-alkaline (peraluminous). Magmatic differentiation by fractional crystallization is the main factor in formation of the leucogranitic rocks. Trace element and REE patterns show distinctive depletion in Rb, K, Nb, Sr, P and Ti, and slight enrichment in Ba, Th, La, Ce, Nd, Zr Low H2O activity, TTG dehydration I-type (Proterozoic TTG) Lower than 10 kbar, 800 –850 8C
Meta-rhyolites and irregularly distributed sills or dykes of meta-quartz porphyries Quartz, microperthitic orthoclase, microcline with minor amount of biotite, titanite, allanite, apatite, zircon and opaque minerals
541.1 + 3.6 Ma (207Pb – 206Pb) in leucogranitic rocks (this study)
543+7 Ma (207Pb– 206Pb) in meta-quartz porphyry rocks (Kro¨ner & S¸engo¨r 1990), 541.3 + 10.9 Ma (207Pb– 206Pb) in meta-rhyolites (Gu¨rsu & Go¨ncu¨oglu 2006) Synkinematic S1, S2 and S3 discrete crenulation cleavage is well developed Post-collisional extension, crustal thinning
Synkinematic S1, S2 and S3 discrete crenulation cleavage is well developed Post-collisional extension, crustal thinning
Sub-alkaline and cogenetic. The role of the fractional crystallization is consistent with the fractionation of an assemblage of plagioclase, alkali feldspar and iron oxides. The trace element patterns show depletion in Nb, Sr, P and Ti and enrichment in Th, U, La, Ce, Nd, Sm and Zr Low H2O activity, granite and felsic rock dehydration I-type (Proterozoic granites or felsic rocks) Lower than 10 kbar, 780 – 845 8C
¨ RSU & M. C. GONCUOGLU S. GU
Felsic magmatism Nature
Ku¨tahya–Bolkar Dag˘ı (NE Afyon) Unit
FELSIC MAGMATISM, AFYON, TURKEY
(U –Pb single-zircon and SHRIMP ages; Anders et al. 2006) arc-related rocks recently reported from the orthogneisses in the Pelagonian basement in Greece, which were also attributed to a Gondwanan terrane (the Florina Terrane of Anders et al. 2006). As indicated by many workers (e.g. Condie 1981, 2005; Martin 1988, 1999; Martin et al. 2005), TTG magmas were derived from subducted basaltic crust in most Archaean cratonic terranes. The TTG-type melting of such a crustal source by post-collisional extension or lithospheric delamination after subduction of the oceanic lithosphere may have resulted in the formation of the Late Neoproterozoic post-collisional granitoids in the basement of the Taurides. This geodynamic scenario is in good agreement with models for various PeriGondwanan terranes (e.g. Murphy et al. 2002; Nance et al. 2002). Based on these similarities, Gu¨rsu & Goncuoglu (2001, 2006) suggested that this magmatic activity in the Taurides should be considered as the eastern counterpart of the Cadomian magmatism. Despite the differences in their post-Cadomian history (Variscan(?) and Alpine events) the geochemical comparison of the granitic rocks in the basements of two tectonic units (Geyik Dag and Ku¨tahya-Bolkar Dagı units) in the Taurides has revealed the presence of two chemically different but genetically linked groups. In very general terms, it can be assumed that the rocks represent the products of more evolved end-members of the same parent magma. However, the present data are not sufficient to allow us speculate on details of the magmatic evolution; detailed radiometric age dating coupled with further geochemical evaluation of associated metaigneous rocks may in obtaining a better understanding of the relations. To conclude, the recent petrological work on the granitic rocks in the Ku¨tahya–Bolkar Dagı tectonic unit of the Taurides has revealed the following results. (1) The Cadomian felsic magmas of the Ku¨tahya– Bolkar Dagı unit were mostly emplaced between 545 and 540 Ma and they are believed to have been intruded during post-collisional stages of Cadomian orogeny. The mineralogical, petrographic and geochemical features of the porphyritic granitoids in the ABC indicate a sub-alkaline (peraluminous) nature. They have geochemical characteristics of I-type felsic intrusive rocks. It is suggested that fractional crystallization was the main differentiation process during the formation of these rocks. (2) The low to very low unfractionated HREE and Y abundances in the granitic rocks suggest that the felsic magmas were evolved through the second-stage melting of a TTG source during intracrustal melting without residual garnet. Their trace
427
element and REE patterns show distinctive depletion in Rb, K, Nb, Sr, P and Ti and slightly enrichment in Ba, Th, La, Ce, Nd and Zr, with fractionated (La/Yb)N and (La/Sm)N, and relatively flat (Gd/Yb)N patterns. They display very similar patterns to a TTG source. The petrogenetic modelling and magmatic temperature studies also imply that the protoliths of the granitic rocks of the ABC were derived by partial melting of a TTG source by 20% fractional melting plus 20% Rayleigh fractional crystallization. This was achieved under ,10 kbar pressure and water-undersaturated conditions. The felsic magma temperatures determined (c. 800–850 8C) are in general agreement with values from experimental studies (e.g. Clemens et al. 1986; Skjerlie & Johnston 1992; Patin˜o-Douce 1997) and support the model that I-type granites may be produced by partial melting of a tonalitic source. (3) We propose that the felsic magmas were generated by crustal extension related to the Cadomian post-collisional magmatism. Therefore they represent the eastern equivalents of the Late Pan-African –Cadomian granitoids in North Africa and Gondwana-derived terranes in Southern and Central Europe (e.g. Chantraine et al. 2001; El-Nisr et al. 2001; Pin et al. 2002; Do¨rr et al. 2002; Mushkin et al. 2003). (4) The ABC-type basement of the Ku¨tahya– Bolkar Dagı unit and the Sandikli-type basement of the Geyik Dag unit are part of the same Cadomian terrane and the present differences in stratigraphy and metamorphism are due to subsequent (Variscan and Alpine) geological events. This research was financially supported by General Directorate of Mineral Research and Exploration (MTA) and State Planning Organisation of the Republic of Turkey (MTA/DPT Project 2004-16B37). The authors gratefully acknowledge the comments of H. Kozlu and N. Turhan. We would like to thank M. Satır for help with Pb–Pb geochronological analyses at the Tu¨bingen University, Laboratory of Geochronology. Special thanks go to S. Ko¨ksal for comments on the zircon typology and morphology. N. Ilbeyli, R. Kerrich and J. P. Lie´geois are thanked for careful reviews and constructive editing.
References A LLE` GRE , C. J. & B EN O THMAN , D. 1980. Nd–Sr isotopic relationship in granitoid rocks and continental crust development: A chemical approach to orogenesis. Nature, 286, 335–342. A NDERS , B., R EISCHMANN , T., K OSTOPOULOS , D. & P OLLER , U. 2006. The oldest rocks of Greece: first evidence for a Precambrian terrane within the Pelagonian Zone. Geological Magazine, 143, 41–58. A RTH , J. G. 1976. Behaviour of trace elements during magmatic process—a summary of theoretical models
428
¨ RSU & M. C. GONCUOGLU S. GU
and their applications. Journal of Research of the US Geological Survey, 4, 41–47. B ALLE` VRE , M., L E G OFF , E. & H E´ BERT , R. 2001. The tectonothermal evolution of the Cadomian belt of northern Brittany, France: a Proterozoic volcanic arc. Tectonophysics, 331, 19–43. B ANDRES , A., E GUILUZ , L., G IL I BARGUCHI , I. J. & P ALACIOS , T. 2002. Geodynamic evolution of a Cadomian arc region: the northern Ossa–Morena zone, Iberian massif. Tectonophysics, 352, 105– 120. B EARD , J. S. & L OFGREN , G. E. 1989. Effect of water on the composition of partial melts of greenstone and amphibolites at 1, 3 and 6.9 kb. Science, 244, 195– 197. B EARD , J. S. & L OFGREN , G. E. 1991. Dehydration melting and water-saturated melting of basaltic and andesitic greenstone and amphibolites at 1, 3 and 6.9 kb. Journal of Petrology, 32, 365–401. ¨ ., O BERHANSLI , R., C ANDAN , O., D ORA , O. O C ETINKAPLAN , M., P ARTZCH , J. H., W ARKUS , F. C. & D U¨ RR , S. 2001. Pan-African high-pressure metamorphism in the Precambrian basement of the Menderes Massif, western Anatolia, Turkey. International Journal of Earth Sciences, 89, 793– 811. ¸ ETINKAPLAN , M., C ANDAN , O., O BERHANSLI , R., C R IMMELE´ , G. & A KAL , C. 2003. Regional occurrence of Fe–Mg carpholite in Triassic metasediments of Afyon Zone, Turkey: implications for metamorphic ¨ . (ed.) Proceedings of 56th evolution. In: D ORA , O. O National Geological Congress of Turkey, Chamber of Turkish Geological Engineers, Ankara, 57–59. ¸ ETINKAPLAN , M., O BERHANSLI , R., C ANDAN , O., C R IMMELE´ , G. & A KAL , C. 2005. Alpine high-P/ low-T metamorphism of the Afyon Zone and implication for the metamorphic evolution of Western Anatolia, Turkey. Lithos, 84, 102 –124. C HANTRAINE , J., E GAL , E., T HIEBLEMONT , D., L E G OFF , E., G UERROT , C., B ALLE` VRE , M. & G UENNOC , P. 2001. The Cadomian active margin (North Armorican Massif, France): a segment of the North Atlantic Pan-African belt. Tectonophysics, 331, 1–18. C HAPPELL , B. W. 1999. Aluminium saturation in I- and S-type granites and the characterization of fractionated haplogranites. Lithos, 46, 535–551. C HEN , F., S IEBEL , W., S ATIR , M. & T ERZIOGLU , M. N. 2002. Geochronology of Karadere basement (NW Turkey) and implications for the geological evolution of the I˙stanbul zone. International Journal of Earth Sciences, 91, 469 –481. C LEMENS , J. D. & V IELZEUF , D. 1987. Constraints on melting and magma production in the crust. Earth and Planetary Science Letters, 86, 287– 306. C LEMENS , J. D., H OLLOWAY , J. R. & W HITE , A. J. R. 1986. Origin of an A-type granite: experimental constraints. American Mineralogist, 71, 317– 324. C OLLINS , W. J., B EAMS , S. D., W HITE , A. J. R. & C HAPPELL , B. W. 1982. Nature and origin of A-type granites, with particular reference to south-eastern Australian. Contributions to Mineralogy and Petrology, 80, 189–200. C ONDIE , C. K. 1981. Archaean Greenstone Belts. Elsevier, Amsterdam. C ONDIE , C. K. 1993. Chemical composition and evolution of the upper continental crust: contrasting results
from surface samples and shales. Chemical Geology, 104, 1 –37. C ONDIE , C. K. 2005. TTGs and adakites: are they both slab melts? Lithos, 80, 33–44. D AVIES , G. R. & M AC D ONALD , R. 1987. Crustal influence in the petrogenesis of the Naivasha basalt– comendite complex: combined trace elements and Sr– Nd– Pb isotope constraints. Journal of Petrology, 28, 1009– 1031. ¨ ., C ANDAN , O., D U¨ RR , S. & O BERHANSLI , R. D ORA , O. O 1995. New evidence on the geotectonic evolution of ¨ ., AKGUN , M., the Menderes Massif. In: PISKIN , O SAVASCIN , M. Y. & TARCAN , G. (eds) International Earth Science Colloquium on the Aegean Region, Proceedings, 9 Eylu¨l University, I˙zmir, Turkey, V-I, 53–72. D O¨ RR , W., Z ULAUF , G., F IALA , J., F RANKE , W. & V EJNAR , Z. 2002. Neoproterozoic to Early Cambrian history of an active plate margin in the Tepla– Barrandian unit—a correlation of U– Pb isotopicdilution– TIMS ages (Bohemia, Czech Republic). Tectonophysics, 352, 65– 85. E GAL , E., G UERROT , C., L E G OFF , D., T HIE´ BLEMONT , D. & C HANTRAINE , J. 1996. The Cadomian orogeny revisited in northern France. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic. Geological Society of America, Special Papers, 304, 281–318. E GAL , E., T HIE´ BLEMONT , D., L AHONDE´ RE , D. ET AL . 2002. Late Eburnean granitization and tectonics along the western and northwestern margin of the Archean Ke´ne´ma– Man domain (Guinea, West African Craton). Precambrian Research, 117, 57– 84. E L -N ISR , S. A., E L -S AYED , M. M. & S ALEH , G. M. 2001. Geochemistry and petrogenesis of Pan-African late to post orogenic younger granitoids at Shalatin– Halaib, South Eastern Desert, Egypt. Journal of African Earth Sciences, 33, 261–282. G ENNA , A., N EHLIG , P., L E G OFF , E., G UERROT , C. & S HANTI , M. 2002. Proterozoic tectonism of the Arabian Shield. Precambrian Research, 117, 21–40. G ONCUOGLU , M. C. 1996. Terranes with contrasting pre-Lower Palaeozoic basement rocks in NW Gondwana: geodynamic implications. Revista Yacimiento Petroliferos Fiscales Bolivianos, 17, 475–476. G ONCUOGLU , M. C. & K OZLU , H. 2000. Early Palaeozoic evolution of the NW Gondwanaland: data from southern Turkey and surrounding regions. Gondwana Research, 3, 315–323. ¨ ZCAN , A., T URHAN , N. & I S¸ IK , A. G ONCUOGLU , M. C., O 1992. Stratigraphy of the Ku¨tahya region. A geotraverse across suture zones in NW Anatolia. General Directorate of Mineral Research and Exploration, Special Publications, 1, 3– 8. G ONCUOGLU , M. C., D IRIK , K. & K OZLU , H. 1997. General characteristics of pre-Alpine and Alpine Terranes in Turkey: explanatory notes to the terrane map of Turkey. Annales Ge´ologiques des Pays Helle´nique, 37, 515– 536. G ONCUOGLU , M. C., T URHAN , N. & T EKIN , U. K. 2003. Evidence for the Triassic rifting and opening of the Neotethyan Izmir– Ankara Ocean and discussion on
FELSIC MAGMATISM, AFYON, TURKEY the presence of Cimmerian events at the northern edge of the Tauride– Anatolide Platform, Turkey. Bollettino Societa` Geologica Italiana, Volume Speciale, 2, 203–212. G U¨ RSU , S. & G ONCUOGLU , M. C. 2001. Characteristic features of the Late Precambrian felsic magmatism in Western Anatolia: implications for the Pan-African evolution in NW PeriGondwana. Gondwana Research, 4, 169– 170. G U¨ RSU , S. & G ONCUOGLU , M. C. 2005. Early Cambrian back-arc volcanism in the western Taurides, Turkey: implications for rifting along the northern Gondwanan margin. Geological Magazine, 142, 617–631. G U¨ RSU , S. & G ONCUOGLU , M. C. 2006. Petrogenesis and tectonic setting of Cadomian felsic igneous rocks, Sandıklı area of the western Taurides, Turkey. International Journal of Earth Sciences, 95, 741–757. G U¨ RSU , S., K OZLU , H., G ONCUOGLU , M. C. & T URHAN , N. 2003. Correlation of the basement rocks and lower Palaeozoic covers of the western part of the Central Taurides. Turkish Association of Petroleum Geologist Bulletin, 15, 129–153 [in Turkish]. G U¨ RSU , S., G ONCUOGLU , M. C. & B AYHAN , H. 2004a. Geology and geochemistry of the pre-Early Cambrian rocks in Sandıklı area: implications for the PanAfrican evolution in NW Gondwanaland. Gondwana Research, 7, 923 –935. G U¨ RSU , S., G ONCUOGLU , M. C., T URHAN , N. & K OZLU , H. 2004b. Characteristic features of Precambrian, Palaeozoic and Lower Mesozoic successions of two different tectono-stratigraphic units of the Taurides in Afyon area, western Central Turkey. In: C HATZIPETROS , A. & P AVLIDES , S. B. (eds) 5th International Symposium on Eastern Mediterranean Geology Proceedings, Aristotle University, Thessaloniki, Greece, 80– 83. H ARRISON , M. T. & W ATSON , B. E. 1984. The behaviour of apatite during crustal anatexis: Equilibrium and kinetic considerations. Geochimica et Cosmochimica Acta, 48, 1467–1477. H ELZ , R. T. 1976. Phase relations of basalts in their melting ranges at PH2O ¼ 5 kb. Part II. Melt composition. Journal of Petrology, 17, 139–193. H ETZEL , R. & R EISCHMANN , T. 1996. Intrusion age of Pan-African augen-gneisses in the southern Menderes Massif and the age of cooling after Alpine ductile extensional deformation. Geological Magazine, 33, 565–572. H ENDERSON , P. 1982. Inorganic Geochemistry. Pergamon, Oxford. H ILDRETH , W. 1981. Gradients in silicic magma chambers: implications for lithospheric magmatism. Journal of Geophysical Research, 86, 10153– 10192. H IRDES , W. & D AVIS , D. W. 2002. U –Pb geochronology of Palaeoproterozoic rocks in the southern part of the Kedougou –Ke´nie´ba inlier, Senegal, West Africa: Evidence for diachronous accretionary development of the Eburnean province. Precambrian Research, 118, 83–99. K IRSTEIN , A. L., P EATE , W. P., H AWKESWORTH , J. C., T URNER , P. S., H ARRIS , C. & M ANTOVANI , M. S. M. 2000. Early Cretaceous basaltic and rhyolitic magmatism in southern Uruguay associated with the
429
opening of the South Atlantic. Journal of Petrology, 41, 1423–1438. K OBER , B. 1986. Whole-grain evaporations for 207 Pb– 206Pb age investigations on single zircons using a double-filament thermal ion source. Contributions to Mineralogy and Petrology, 93, 482–490. ¨ ., C HEN , F., S ATIR , M. & K ORALAY , E., D ORA , O. O C ANDAN , O. 2004. Geochemistry and geochronology of orthogneisses in the Derbent (Alas¸ehir) area, eastern ¨ demis¸ –Kiraz submassif, Menderes Massif: part of O Pan-African magmatic activity. Turkish Journal of Earth Science, 13, 37–61. K OZLU , H. & G ONCUOGLU , M. C. 1997. Stratigraphy of the Infracambrian rock-units in the Eastern Taurides and their correlation with similar units in Southern Turkey. In: G ONCUOGLU , M. C. & D ERMAN , A. S. (eds) Early Palaeozoic in NW Gondwana. Turkish Association Petroleum Geologists, Special Publications, 3, 50– 61. K RO¨ NER , A. & S ENGO¨ R , A. M. C. 1990. Archean and Proterozoic ancestry in the Late Precambrian to Early Palaeozoic crustal elements of southern Turkey as revealed by single zircon dating. Geology, 18, 1186– 1190. L EAT , P. T., J ACKSON , S. E., T HORPE , R. S. & S TILLMAN , C. T. 1986. Geochemistry of bimodal basalt–subalkaline/peralkaline rhyolite provinces within the Southern British Caledonides. Journal of the Geological Society, London, 143, 259– 273. L IE´ GEOIS , J. P., B ERZA , T., T ATU , M. & D UCHESNE , J. C. 1996. The Neoproterozoic Pan-African basement from the Alpine Lower Danubian nappe system (South Carpathians, Romania). Precambrian Research, 80, 281– 301. L IEW , T. C. & M C C ULLOCH , M. T. 1985. Genesis of granitoid batholiths of Peninsular Malaysia and implications for models of crustal evolution: evidence from a Nd–Sr isotopic and U–Pb zircon study. Geochimica et Cosmochimica Acta, 49, 587– 600. L INNEMANN , U. & R OMER , R. L. 2002. The Cadomian orogeny in Saxo-Thuringia, Germany. Geochemical and Nd–Sr– Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance. Tectonophysics, 352, 33–64. L INNEMANN , U., G EHMLICH , M., T ICHOMIROWA , M. ET AL . 2000. From Cadomian subduction to early Palaeozoic rifting: the evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (Central European Variscides, Germany). In: F RANKE , W., H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Orogenic Processes: Quantification and Modelling in the Variscan Belt of Central Europe. Geological Society, London, Special Publications, 179, 131–153. L INNEMANN , U., M C N AUGHTON , N. J., R OMER , R. L., G EHMLICH , M., D ROST , K. & T ONK , C. 2004. West African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana?—U/Pb-SHRIMP zircon evidence and the Nd-isotopic record. International Journal of Earth Sciences, 93, 683 –705. L OOS , S. & R EISCHMANN , T. 1999. The evolution of the southern Menderes Massif in SW Turkey as revealed
430
¨ RSU & M. C. GONCUOGLU S. GU
by zircon dating. Journal of Geological Society, London, 156, 1021–1030. M ARTIN , H. 1988. Archaean and modern granitoids as indicators of changes in geodynamic processes. Revista Brasileira de Geociencias, 17, 360– 365. M ARTIN , H. 1999. The adakitic magmas: modern analogues of Archaean granitoids. Lithos, 46, 411– 429. M ARTIN , H., S MITHIES , R. H., R APP , R., M OYEN , J. F. & C HAMPION , D. 2005. An overview of adakite, tonalite– trondhjemite– granodiorite (TTG), and sanukitoid: relationships and some implications for crustal evolution. Lithos, 79, 1– 24. M INGRAM , B., K RO¨ NER , A., H EGNER , E. & K RENTZ , O. 2004. Zircon ages, geochemistry, and Nd isotopic systamatics of pre-Variscan orthogneisses from the Erzgebirge, Saxony (Germany), and geodynamic interpretation. International Journal of Earth Sciences, 93, 706 –727. M ONTEL , J. M. 1993. A model for monazite/melt equilibrium and the applications to the generation of granitic magmas. Chemical Geology, 110, 127–146. M URPHY , J. B., E GUILUZ , L. & Z ULAUF , G. 2002. Cadomian Orogens, peri-Gondwanan correlatives and Laurentia– Baltica connections. Tectonophysics, 352, 1– 9. M URPHY , J. B., P ISAREVSKY , S. A., N ANCE , R. D. & K EPPIE , J. D. 2004. Neoproterozoic– Early Palaeozoic evolution of peri-Gondwanan terranes: implications for Laurentia–Gondwana connections. International Journal of Earth Sciences, 93, 659– 682. M USHKIN , A., N AVON , O., H ALICZ , L., H ARTMANN , G. & S TEIN , M. 2003. The petrogenesis of A-type magmas from the Amram massif, Southern Israel. Journal of Petrology, 44, 815 –832. N ANCE , R. D., M URPHY , J. B., K EPPIE , J. D. & O’ BRIEN , S. J. 2002. A Cordilleran model for the evolution of Avalonia. Tectonophysics, 352, 11–31. N ASH , W. P. & C RECRAFT , H. R. 1985. Partition coefficient for trace elements in silicic magmas. Geochimica et Cosmochimica Acta, 49, 2309–2322. N EUBAUER , F. 2002. Evolution of late Neoproterozoic to early Palaeozoic tectonic elements in Central and Southeast European Alpine mountain belts: review and synthesis. Tectonophysics, 352, 87– 103. ¨ ZCAN , A., G ONCUOGLU , M. C. & T URHAN , N. 1989. O The geology of Ku¨tahya, Cifteler, Bayat and Ihsaniye area. Mineral Research and Exploration General Directorate Report, 8974. P ATIN˜ O -D OUCE , A. E. 1997. Generation of metaaluminous A-type granites by low-pressure melting of calc-alkaline granitoids. Geology, 25, 743– 746. P ATIN˜ O , M. L. G. & P ATIN˜ O -D OUCE , A. E. 1987. Petrologia y petroge´nesis del Batolito de Achala, provincial de Co´rdoba, a la luz de la evidencia de campo. Revista de la Asociacio´n Geolo´gica Argentina, 42, 201– 205. P ATOCˇ KA , F. & Sˇ TORCH , P. 2004. Evolution of geochemistry and depositional settings of Early Palaeozoic siliciclastics of the Barrandian (Tepla´ – Barrandian Unit, Bohemian Massif, Czech Republic). International Journal of Earth Sciences, 93, 728–741. P EARCE , J. A. & N ORRY , M. J. 1979. Petrogenetic implication of Ti, Zr, Y and Nb variations in volcanic
rocks. Contributions to Mineralogy and Petrology, 69, 33–47. P ECCERILLO , A. & T AYLOR , S. R. 1976. Geochemistry of Eocene calcalkaline volcanic rocks from the Kastamonu area, northern Turkey. Contributions to Mineralogy and Petrology, 58, 63–81. P ICCOLI , P. M. & C ANDELA , P. A. 1994. Apatite in felsic rocks, a model for the estimation of initial halogen concentrations in the Bishop Tuff (Long Walley) and Tuolumne intrusive suite (Sierra Nevada Batholith) magmas. American Journal of Science, 294, 92–135. P ICCOLI , P. M., C ANDELA , P. A. & W ILLIAMS , T. J. 1999. Estimation of aqueous HCl2 and Cl2 concentrations in felsic systems. Lithos, 46, 591–604. P ICHAVANT , M., M ONTEL , J. M. & R ICHARD , L. R., 1992. Apatite solubility in peraluminous liquids: experimental data and an extension of the Harrison– Watson model. Geochimica et Cosmochimica Acta, 56, 3855– 3861. P IN , C., L INA´ N , E., P ASCUAL , E., D ONAIRE , T. & V ALENZUELA , A. 2002. Late Proterozoic crustal growth in the European Variscides: Nd isotope and geochemical evidence from the Sierra de Cordoba andesites (Ossa– Morena Zone, southern Spain). Tectonophysics, 352, 133–151. P UPIN , J. P. 1980. Zircon and granite petrology. Contributions to Mineralogy and Petrology, 73, 207–220. R UTTER , M. J. & W YLLIE , P. J. 1988. Melting of vapour-absent tonalite at 10 kbar to simulate dehydration melting in the deep crust. Nature, 331, 159–160. S ING , J. & J OHANNES , W. 1996a. Dehydration melting of tonalities. Part I. Beginning of melting. Contributions to Mineralogy and Petrology, 125, 16– 25. S ING , J. & J OHANNES , W. 1996b. Dehydration melting of tonalities. Part II. Composition of melting and solids. Contributions to Mineralogy and Petrology, 125, 26–44. S KJERLIE , K. P. & J OHNSTON , A. D. 1992. Fluid-absent melting behaviour of and F-rich tonalitic gneiss at mid-crustal pressures: implications for the generation of anorogenic granites. Journal of Petrology, 34, 785–815. S PULBER , S. D. & R UTHERFORD , M. J. 1983. The origin of rhyolite and plagiogranite in oceanic crust: an example study. Journal of Petrology, 24, 1– 25. S TRACHAN , R. A., D’L EMOS , R. S. & D ALLMEYER , R. D., 1996. Late Precambrian evolution of an active plate margin: North Armorican Massif, France. In: N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian and Related Peri-Gondwanan Terranes of the CircumNorth Atlantic. Geological Society of America, Special Papers, 304, 319– 332. 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. 1995. The geochemical evolution of the continental crust. Reviews of Geophysics, 33, 241–265.
FELSIC MAGMATISM, AFYON, TURKEY T HOMSON , A. B. 1996. Fertility of crustal rocks during anatexis. Transactions of the Royal Society of Edinburgh, Earth Sciences, 87, 1– 10. U STAO¨ MER , P. A., M UNDIL , R. & R ENNE , P. R. 2005. U–Pb and Pb– Pb zircon ages for arc-related intrusions of the Bolu Massif (W Pontides, NW Turkey): evidence for Late Precambrian (Cadomian) age. Terra Nova, 17, 215– 223. V AVRA , G., G EBAUER , D., S CHMID , R. & C OMPSTON , W. 1996. Multiple zircon growth and recrystallization during polyphase late Carbonifereous to Triassic metamorphism in granulites of the Ivrea Zone (Southern Alps): an ion microprobe (SHRIMP) study. Contributions to Mineralogy and Petrology, 122, 337– 358. V IELZEUF , D. & M ONTEL , J. M. 1992. Experimental determination of fluid-absent melting of a natural quartz-rich metagreywacke, 1. phase relations. Terra Abstracts, 3, 30. W ATSON , B. E. & H ARRISON , M. T. 1983. Zircon saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth and Planetary Science Letters, 64, 295– 304. W INCHESTER , J. A. & F LOYD , P. A. 1976. Geochemical magma type discrimination: application to altered
431
and metamorphosed basic igneous rocks. Earth and Planetary Science Letters, 28, 459–469. 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. W INTHER , K. T. & N EWTON , R. C. 1991. Experimental melting of hydrous low-K tholeiite: evidence on the origin of Archaean craton. Bulletin of the Geological Society of Denmark, 39, 213– 228. W OLF , M. B. & W YLLIE , P. J. 1994. Dehydration melting of amphibolites at 10 kbar: the effects of temperature and time. Contributions to Mineralogy and Petrology, 115, 369–383. Z ENG , L., A SIMOW , D. P. & S ALEEBY , J. B. 2005. Coupling of anatectic reactions and dissolution of accessory phases and the Sr and Nd isotope systematics of anatectic melts from a metasedimentary source. Geochimica et Cosmochimica Acta, 64, 3671– 3682. Z HAO , J. H., H U , R. & L IU , S. 2004. Geochemistry, petrogenesis and tectonic significance of Mesozoic mafic dykes, Fujian Province, Southeastern China. International Geology Review, 46, 542– 557.
The Anti-Atlas chain (Morocco): the southern margin of the Variscan belt along the edge of the West African craton ABDERRAHMANE SOULAIMANI1,2 & MARTIN BURKHARD3,4 1
Universite´ Cadi Ayyad, Faculte´ des Sciences Semlalia, av. Moulay Abdellah, BP 2390, Marrakech, Morocco 2
Present address: 18 rue Stephenson, 59000 Lille, France (e-mail:
[email protected]) 3 Institut de Ge´ologie, Universite´ de Neuchaˆtel, rue E´mile-Argand 11, CP 2, 2007 Neuchaˆtel, Switzerland 4
Deceased August 2006
Abstract: Broadly synchronous circum-Atlantic Variscan –Alleghanian orogenic belts developed during the Late Palaeozoic Gondwana –Laurentia collision. In the northern part of the West African craton (WAC), the Variscan orogeny produced basement-controlled structures in the Anti-Atlas, which represents the pericratonic foreland, now located south of the Variscan domains of Morocco and north of the Mauritanides belt. New structural field observations document the strong involvement of the basement and the inversion and folding of the Palaeozoic sedimentary basins at the edge the WAC. Two contrasting domains differently responding to regional NW– SE shortening are recognized: (1) a narrow belt along the Atlantic coast characterized by thin-skinned folding and ESE-vergent thrusting (para-autochthonous Anti-Atlas); (2) a large area between the WAC sensu stricto and the South Atlas front showing huge basement uplifts amidst a folded Palaeozoic cover with upright polyharmonic folds (autochthonous Anti-Atlas). The structural trend of the basement inliers is inherited at least in some case from previous Proterozoic fractures. Compressional reactivations led to basement uplift and concomitant folding of the Palaeozoic cover. Cover series are horizontally shortened by mostly upright symmetrical buckle folds of various wavelengths in response to thickness variations between abundant incompetent silt and shale horizons and rare competent carbonate and quartzite beds. Deformation is greatest near the borders of and between closely spaced basement uplifts. Regionally, deformation intensity decreases, either abruptly or progressively, towards the SE and it vanishes within the undeformed Tindouf basin.
The West African craton (WAC), stable since 2 Ga, constitutes the basement of northwestern Africa (Rocci et al. 1991). Along the northern margin of this craton, especially in Morocco, the Reguibate shield is rimmed by a series of mobile belts with decreasing ages toward the north (Fig. 1a). The southernmost of these zones, the Anti-Atlas belt, is located between the Alpine Atlas chain and the northern rim of the Palaeozoic Tindouf basin. Gentil (1918), Lecointre (1926) and many other workers showed that a Late Palaeozoic orogeny affected the northern Moroccan Meseta as well as the Palaeozoic inliers of the Atlas and Rif. The Anti-Atlas belt exposes thick, unmetamorphosed and often mildly deformed Palaeozoic rocks, which unconformably overlie a Precambrian basement included in several NE –SW inliers (Choubert 1952). Pique´ & Michard (1989) proposed that these Moroccan Palaeozoic structures were contemporaneous with the Variscan orogeny in Europe.
The oldest Precambrian rocks exposed in the Anti-Atlas belt represent splinters derived from the Eburnean WAC (Aı¨t Malek et al. 1998; Thomas et al. 2002; Walsh et al. 2002; Gasquet et al. 2004). However, although comparable rocks located within the shield remained undisturbed since about 2 Ga, those of the Anti-Atlas region display evidence of two superimposed pre-Variscan orogenic events: (1) a Neoproterozoic compressive event, related to the Pan-African orogeny (Leblanc 1975; Saquaque et al. 1989); (2) a subsequent extensional event, which affected the northern margin of the craton during the Early Palaeozoic, or as early as the late Neoproterozoic (Pique´ et al. 1999; Doblas et al. 2002; Soulaimani et al. 2003). During this extension, the Anti-Atlas domain experienced rifting, tilting of basement blocks and the creation of sedimentary rift basins. This event is documented by the volcanoclastic series of the Latest Proterozoic ‘Ouarzazate Supergroup’ to
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 433–452. DOI: 10.1144/SP297.20 0305-8719/08/$15.00 # The Geological Society of London 2008.
434 A. SOULAIMANI & M. BURKHARD
Fig. 1. (a) Location of the Anti-Atlas within the main structural domains of NW Africa. (b) Generalized geological map of the Anti-Atlas, simplified from the 1:1 000 000 geological map of Morocco (Maroc Service Ge´ologique 1985); showing the location of the areas studied: 1, western Atlantic Bas Draˆa; 2, eastern Bas Draˆa and western Bani; 3, Lakhssas Plateau; 4, Irherm–Tata; 5, Bou Azzer–El Graara; 6, Saghro –Ougnate. Inliers: BD, Bas Draˆa; If, Ifni; Kr, Kerdous; Ir, Igherm; TA, Tagragra d’Akka; TT, Tagragra Tata; AM, Agadir Melloul; Ze, Zenaga; Sr, Sirwa; Bz, Bou Azzer– El Graara; Sg, Saghro; Og, Ougnate.
THE ANTI-ATLAS CHAIN, MOROCCO
Earliest Cambrian (terminology after Thomas et al. 2004). Synsedimentary extensional deformations are also recorded within the Early Cambrian marine ‘Taroudant Group’ (Buggisch & Siegert 1988; Algouti et al. 2002; Benssaou & Hamoumi 2003). Later, up to 10 km of shallow-marine sediments overlying the synrift successions were deposited during the rest of the Palaeozoic in an intracratonic setting (Burkhard et al. 2006), at least in the western –central Anti-Atlas. The architecture of the Precambrian basement inherited from the Late Proterozoic– Early Cambrian extensional episode plays a significant role in the geometry of subsequent geodynamic events. In particular, it has a strong influence on the Variscan structures that developed in the course of the Late Carboniferous compressive deformations. The role of the Precambrian basement configuration controlling the mode of Variscan shortening has been documented in many studies (Michard 1976; Jeannette & Pique´ 1981; Soulaimani et al. 1997). Different trends of the Variscan regional folds have been interpreted as the result of reactivation of basement faults, in particular by the basement uplifts following the inherited zones of crustal weakness (Leblanc 1972, 1975; Donzeau 1974; Jeannette & Pique´ 1981; Hassenforder 1987; Soulaimani 1998; Belfoul et al. 2001). Similarly, variations in the thickness of the Palaeozoic cover series between and above the Precambrian rigid basement uplifts induced the development of disharmonic folds with different wavelengths, separated by numerous de´collement levels (Soulaimani et al. 1997; Burkhard et al. 2001; Caritg et al. 2004; Helg et al. 2004). This paper provides a comprehensive description of the geometry of the Anti-Atlas Variscan belt, including a kinematic study that has been hitherto missing from the literature as well as inferences for the geodynamics of this southern branch of the Variscan belt. The present synopsis is based on the geometry and finite strain observations of the Variscan deformation in several key areas of the Anti-Atlas.
Regional geology of the Anti-Atlas General structural pattern The Anti-Atlas chain is a broad anticlinorium some 800 km long and 200 km wide, trending ENE– WSW (Fig. 1), parallel to the Alpine High Atlas chain. This situation suggests that Alpine –Atlas rejuvenation might have contributed significantly to the present-day elevation (up to 2500 m) of the Variscan Anti-Atlas belt. Indeed, this recent uplift is well established in the eastern part of the
435
Anti-Atlas (Ougnat and Tafilalt areas), where the Cretaceous and Neogene cover series are clearly tilted northward along the northern border of the belt (Robert-Charrue 2006). In comparison with the structural observations in the adjacent High Atlas, the Alpine uplift of the Anti-Atlas belt occurred during the Pliocene –Pleistocene (Frizon de Lamotte et al. 2000), partly owing to a Neogene thermal uplift (Teixell et al. 2005; Missenard et al. 2006). Along the main axis of the Anti-Atlas anticlinorium, the most extensively exposed rocks are carbonates of Early to Middle Cambrian age, wrapping around a series of Precambrian basement inliers (Bas Draˆa, Kerdous, Irherm, Zenaga, Bou Azzer –El Graara, Saghro, etc.). Because Cambrian carbonates are more resistant to erosion than the Proterozoic crystalline basement rocks, these basement ‘uplifts’ often correspond to topographic depressions. This explains the term ‘boutonnie`re’ (literally ‘buttonhole’) of the French workers (Choubert 1952, 1963). Along the southern border of the Anti-Atlas anticlinorium, Palaeozoic sedimentary series dip generally to the SSE, thus revealing successively younger strata to the south (Fig. 1). The total thickness of the Palaeozoic stratigraphical column varies regionally from more than 10 km in the west (Fig. 2) to less than 5 km in the east. Palaeozoic cover sequences are disharmonically folded with deformation intensity decreasing towards the SE on a given transect. Along strike, folding style also changes and a general decrease in folding intensity is apparent from west to east. In the north, the Alpine orogeny affected the northern part of the Anti-Atlas, creating thrust faults and folds, contemporaneous with the High Atlas structures. These Alpine structures are unconformably overlain by Neogene clastic sediments of the south-Atlas Souss and Ouarzazate basins.
Tectonostratigraphic units The Proterozoic crystalline basement of the AntiAtlas and its thick (.10 km) sedimentary cover have been extensively described (e.g. Choubert 1963; Destombes et al. 1985; Thomas et al. 2004; Gasquet et al. 2005) and therefore only a synthetic column will be briefly described below (Fig. 2). The Proterozoic basement. The Proterozoic rocks of Morocco are classically subdivided into three major unconformity-bounded assemblages (Choubert 1963; Leblanc 1975; Charlot 1978; Hassenforder 1987). The lowermost Palaeoproterozoic units consist of low- to medium-grade schists and intrusive granitoids, some of them now in the form of orthogneisses, attributed to the Eburnean orogeny (c. 2 Ga) (Aı¨t Malek et al. 1998; Thomas et al.
436
A. SOULAIMANI & M. BURKHARD
2002; Walsh et al. 2002; Gasquet et al. 2004). Eburnean basement rocks occur exclusively SW of the ESE–WNW-trending Anti-Atlas Major Fault (AAMF) (Choubert 1947). However, they probably constitute the basement of the Pan-African rocks to the NE (Ennih & Lie´geois 2001; Gasquet et al. 2005). Consistently with observations elsewhere in the West African craton, no Mesoproterozoic event or rocks are recorded in the Anti-Atlas. Palaeoproterozoic basement rocks are, therefore, unconformably overlain directly by Neoproterozoic metasedimentary and magmatic units, which were subsequently affected by the Pan-African orogeny (Clauer 1974; Leblanc 1975; Hassenforder 1987). The Pan-African event was related to the oblique suturing of the northern rifted margin of the Eburnean continental mass and the Neoproterozoic Saghro magmatic arc (Saquaque et al. 1989). An ophiolitic complex has been described in the central AntiAtlas along the AAMF from the southern side of the Sirwa Massif (Admou & Juteau 2000; Thomas et al. 2002) to the Bou Azzer– El Graara inlier (Leblanc 1975; Saquaque et al. 1992). The polarity of the Neoproterozoic subduction is still a matter of debate (Ennih & Lie´geois 2001; Hefferan et al. 2000; Gasquet et al. 2005), but the present-day deep structure supports the hypothesis of a worthdipping subduction zone, at least during the late Pan-African stage (Soulaimani et al. 2006). The unconformably overlying syntectonic volcanoclastic molasse deposits, attributed to the top of the ‘Saghro Group’, have been affected by the ultimate phases of the Neoproterozoic Pan-African deformation (Thomas et al. 2002). In the Bou Azzer –El Graara inlier, the Tiddiline series are regarded as late Pan-African syntectonic molasse deposits, folded and cut across by sinistral oblique north-dipping thrust faults in late Pan-African syntectonic molasse basins (Leblanc & Lancelot 1980; Hefferan et al. 1992).
Fig. 2. Synthetic tectonostratigraphic columns for the Proterozoic basement and Palaeozoic cover of the Anti-Atlas.
The Late Precambrian and Palaeozoic cover. The Latest Neoproterozoic rocks are the sedimentary, mainly clastic, and volcanic rocks of the ‘Ouarzazate Supergroup’, separated from the crystalline basement by a major unconformity. They grade upward after a slight unconformity into the ‘Adoudounian’ carbonates, marls and siltstones, which record the Early Cambrian marine transgression. These latest Neoproterozoic–Early Cambrian cover sequences unconformably overlie older Precambrian basement structures. The ‘Ouarzazate Supergroup’ sequences are characterized by dramatic thickness changes across fault-bounded basement blocks (Pique´ et al. 1999). They were deposited in faulted basins developed in an
THE ANTI-ATLAS CHAIN, MOROCCO
intracontinental rifting context during a late to post-Pan-African extensional event (Azizi Samir et al. 1990; Thomas et al. 2002; Soulaimani et al. 2003). This event is associated with a major calc-alkaline late to post-orogenic magmatism (Boyer et al. 1978; Youbi 1998), which evolved toward tholeitic and alkaline lavas by the end of the Proterozoic and during the Early Cambrian (Jbel (J.) Boho seynite) (Ducrot & Lancelot 1977; ´ lvaro et al. 2006). Soulaimani et al. 2004; A The geometry and geodynamic significance of this rifting episode are still open to discussion. This event could correspond to a late orogenic extensional collapse of the Pan-African chain. In the most recent interpretation (Soulaimani & Pique´ 2004; Oudra et al. 2006), the Late Proterozoic extensional event is associated with an important tectonothermal reworking of the Precambrian basement, which would have formed metamorphic domes, similar to the related structures developed along the rim of the West African craton during Late Precambrian–Early Cambrian times (Doblas et al. 2002). The Adoudounian rocks are conformably overlain by a post-rift sedimentary sequence starting towards the end of the Early Cambrian and continuing through most of the Palaeozoic. From the Middle Cambrian to the Middle Devonian, the AntiAtlas domain was a shallow-marine shelf, where mostly fine-grained detrital sediments eroded from the West African craton accumulated. A longlasting subsidence resulted in a progressive southward migration of the shoreline over the Reguibate shield. Transgressive sediments of the Anti-Atlas domain thus became progressively finer grained upward, forming a megasequence consisting of Ordovician sands and silts, Silurian shales and, except in the eastern Anti-Atlas, Early and Middle Devonian platform carbonates. The presence of very fine-grained clastic detritus within the Late Devonian carbonates might indicate the end of the passive transgressive regime and the beginning of the central Anti-Atlas uplift (Hassenforder 1987). The Early Carboniferous strata of the eastern Anti-Atlas, which include olistostromes, were deposited in an intramontane trough. Along the southern border of the Anti-Atlas, the Late Vise´an, Namurian and Westphalian sequences of the Ouarkziz, Betana and Bechar remained undisturbed, although they were deposited in a regressive mega-sequence. The Late Mesozoic-Cenozoic cover. Triassic and Jurassic sediments are limited to the north, along the South Atlas Fault in the Ouarzazate and Souss basins, and in the SW, in the Atlantic Tarfaya basins onshore and offshore (e.g. Le Roy & Pique´ 2001). They are related to the Pangaea break-up and opening of the central Atlantic Ocean and Atlas rift basins in Late Triassic– Early Jurassic
437
times. This extension affects both the crystalline basement and the folded Palaeozoic cover of the inner Anti-Atlas itself, being recorded there by the intrusion of NE– SW-trending dykes and associated sills of gabbro and dolerite (Sebai et al. 1991). During the Late Cretaceous, a shallow sea invaded the internal domains of the Anti-Atlas, resulting in the deposition of vast flat-lying strata (hamada). The deposition of these nearly horizontal plateaux lasted until the Oligocene, and sealed the eroded Palaeozoic fold belt.
The Variscan deformation The western Atlantic Bas Draˆa area In the westernmost part of the Anti-Atlas, south of the Precambrian Ifni inlier, the western Atlantic Bas Draˆa domain is a narrow, elongate terrane located between the Bas Draˆa inlier to the east and the Atlantic coast to the west (Fig. 1). Deformed Cambrian rocks constitute NNE– SSW-trending ridges (Choubert 1963; Soulaimani 1998; Belfoul et al. 2001). The folds, with wavelengths decreasing eastward, are cut by repeated east-vergent thrust faults (Fig. 3b). These thrust units bounding minor units form recurrent parallel ridges responsible for the ‘valley and ridge’ morphology of this region. At the outcrop scale, the main structure shown in the Cambrian rocks is a west-dipping penetrative cleavage. Along the Atlantic coast, Lower Cambrian strata are involved in tight recumbent folds, trending NNE –SSW and clearly overturned eastward as shown by their wellpreserved inverted limbs. Their axial planes are parallel to a pervasive foliation, which is slightly inclined to the WNW to subhorizontal; for example, at the so-called ‘Plage Blanche’ (Fig. 3d). The WNW-plunging stretching lineation is parallel to the dip of the cleavage. Shear bands related to non-coaxial flow along the cleavage planes represent the latest Variscan structures. Regionally the deformation intensity decreases rapidly eastward, and in the Bou-Jerif area the Cambrian strata are deformed only by decametrescale folds with a westward steeply dipping rough cleavage, mainly visible in the pelitic layers. Farther to the east, in the Goulmime area some 20 km from the Atlantic shoreline, the Middle Cambrian strata are only slightly tilted from a nearhorizontal orientation and lack any cleavage. The metamorphic conditions associated with the deformation nowhere exceed the lower greenschist facies. The typical east-vergent thrusts and folds described in this area, which considerably differs from the other Anti-Atlas areas, can be seen as typical of the external zone of an orogenic belt.
438
A. SOULAIMANI & M. BURKHARD
Fig. 3. (a) Location of the Atlantic Bas Draˆa area in the Anti-Atlas belt. (b) Structural map of the western Atlantic Bas Draˆa domain. (c) Stereograms (Wulff stereonet, lower hemisphere) representing bedding data (S0) and foliation data (S1). (d) Geological cross-section (section location A– A0 in the map).
The eastern Bas Draˆa and western Bani area East of the Goulmine–Tan Tan meridian line, the Bas Draˆa area (Fig. 4b) exposes Precambrian basement (Bourcart 1937; Choubert & Faure-Muret 1969) and its Palaeozoic cover affected by the Variscan contractional deformation (Maze´as & Pouit 1968; Soulaimani et al. 1997). The crystalline basement remained relatively undeformed throughout the Variscan orogeny, showing only reverse faults along its margins. In contrast, intense deformation is observed at the base of the sedimentary cover, especially within the Late Precambrian volcanoclastic rocks of the ‘Ouarzazate Supergroup’. A locally pervasive ENE –WSW-striking, moderately to steeply dipping foliation occurs NW of the basement massif. To the SE of the massif, where way-up criteria can be ascertained, kinematic indicators point to a clear thrust component toward the SE. A strong stretching lineation is observed within the deformed ‘Ouarzazate Supergroup’ conglomerates, associated with a pervasive axial planar cleavage and tight to isoclinal SE-vergent folds. Detailed mapping of these regional structures
around the inlier (Fig. 4b) demonstrates that the cleavage intensity increases towards the margins of the inlier, where cleavage is rotated into parallelism with the boundary of the inlier. This cleavage refraction suggests that the basement acted as a rigid buttress during the contractional deformation. The basement was pushed upon the cover units along its southern border (Fig. 4c; Soulaimani et al. 1997). To the SE of the Bas Draˆa inlier, disharmonic folds in the Lower Cambrian limestones suggest strain heterogeneity across the ductile deformed strata. Overlying Cambrian –Ordovician Bani sandstones are deformed by kilometre-scale open and cylindrical folds, slightly overturned southeastward and associated with an incipient cleavage in their hinges. To the SE, the J. Rich folds developed within the Devonian sandstones and carbonates are characterized by a decametre- to kilometre-scale wavelength with subhorizontal fold axes. Folds generally appear asymmetrical with their southern limbs in a subvertical to overturned position. Oblique folds, with deviations as much as 20 –308E from the regional N708E fold axis trend of the J. Rich, display an en-echelon orientation.
THE ANTI-ATLAS CHAIN, MOROCCO
439
Fig. 4. (a) Location of the eastern Bas Draˆa and western Bani area in the Anti-Atlas belt. The stereograms (Wulff stereonet, lower hemisphere) represent bedding data (S0) and foliation data (S1). (b) Structural map of the eastern Bas Draˆa and western Bani area and (c) geological cross-section (section location B–B0 in the map).
These en-echelon folds are attributed to ENE– WSW-striking dextral fractures. The asymmetry of these folds has been interpreted as the result of a right-lateral N708E-striking Variscan shear zone along the southern border of the Anti-Atlas, parallel to the South Atlas Fault (Michard 1976; Jeannette & Pique´ 1981). However, the geometry of the J. Rich folds, and the gentle lateral plunge of their axes, is regarded as incompatible with such a mega-shear zone interpretation, let alone with the ‘Alpine wrench folding’ with vertical axes such as proposed by Weijermars (1993). South of the Oued Draˆa, the Upper Devonian shales are the southernmost formations of the AntiAtlas affected by metre-size folds. The J. Tazout and J. Ouarkziz Carboniferous strata are slightly tilted toward the south, at the northern border of the Tindouf basin. The J. Ouarkziz limestones are imprinted only by tectonic stylolites, indicating a very small amount of horizontal shortening (Helg et al. 2004).
The Lakhsass Plateau area North of the Bas Draˆa domain, about 100 km south of Agadir, the Lakhsass Plateau (Fig. 1) is located
between the Ifni and Kerdous inliers. The plateau area corresponds to a large synclinorium of Lower Cambrian carbonate rocks showing an anticlinal structure (J. Inter horst) in its core (Fig. 5b). In the central part of the plateau, gravimetric and magnetic data suggest the presence of an uplifted basement block (J. Inter) beneath the deformed limestones. This basement high is interpreted as a horst produced by the Late Proterozoic extensional episode, and subsequently inverted during the Variscan compression (Soulaimani 1998). The Variscan compressive deformation in the Lakhsass Plateau area is very heterogeneous, with the greatest intensity towards the centre (Fig. 5c). In the cover rocks above the basement horst, and particularly along its flanks, the prominent fabric comprises a pervasive north–south-striking axial planar cleavage associated with tight to isoclinal upright folds. A subvertical stretching lineation is observed along ductile shear planes east of the J. Inter, suggesting a significant component of dip-slip motion. Deformation decreases laterally from the central horst area towards the margins of the synclinorium, where the Upper Proterozoic and Lower Palaeozoic strata are gently tilted towards the axis of the plateau. Everywhere,
440
A. SOULAIMANI & M. BURKHARD
Fig. 5. (a) Location of the Lakhssas Plateau area in the Anti-Atlas belt. (b) Structural map of the Lakhssas Plateau area and (c) geological cross-section (section location C– C0 in the map). The Stereograms (Wulff stereonet, lower hemisphere) represent bedding data (S0) and foliation data (S1).
bedding strikes north– south and dips either to the east or west, defining folds with hinges plunging gently to the north or south. We conclude that Lakhssas Plateau has been affected by a major Variscan deformation with a nearly horizontal regional compressive maximum stress. During this event, pre-existing basement
fractures were reactivated, leading mainly to the uplift of the J. Inter horst, and of the Kerdous and Ifni inliers. These reactivations and the induced compressional structures in the overlying cover probably operated in a transpressional zone, as suggested by the NW– SE direction of shortening at the regional scale.
THE ANTI-ATLAS CHAIN, MOROCCO
The Irherm – Tata area Northeast of the Kerdous inlier (Fig. 1), the Irherm –Tata area is a zone with wide synclines cored by Middle Cambrian rocks (Fouanou, Issafene, Talat N’Issi and Tagmout), and separated from each other by narrow anticlines cored by Precambrian rocks (Aı¨t Abdellah, Alma, Igherm and Tata) (Fig. 6b). Within the northern and southern parts of the area, the structural pattern of the Variscan folds seems erratic and, according to one of us (M.B.), at least two successive folding episodes are recognizable (Caritg et al. 2004; Helg et al. 2004; Burkhard et al. 2006). However, this is not supported by field observations of refolded axes or intersecting cleavage planes. Moreover, there is a regional pattern of fold trends that would instead suggest a mosaic of basement blocks, as follows. (1) To the south, in the Tata region, the east –west J. Bani Quartzite ridge (‘Tata Fault’: Hassenforder 1987; Faik et al. 2001; Caritg et al. 2004; Helg et al. 2004) separates two areas: a southern area where the NE–SW-trending J. Rich folds are abruptly juxtaposed against the J. Bani quartzite ridge, and a northern area where south-vergent folds strike east– west to NW– SE. The abrupt change of the general strike from NE –SW to east –west was interpreted as a response to dextral shear along the Tata Fault (Hassenforder 1987; Faik et al. 2001), although transcurrent displacement along this zone is not apparent at the outcrop scale. The north– south to N208E directions of the Variscan foliations and folds observed at the western prolongation of the Tata Fault, close to the Agouliz inlier, suggest that the latter massif was less reactivated, with transcurrent motion during the Variscan compression. It is important to note that the Tata Fault, like many other east –west faults within the Anti-Atlas, still controls the recent Atlas uplift, which was estimated here to reach some 600 m (Choubert 1952). (2) North and NW of the Tata region, the kilometre-scale Variscan folds yield a roughly north–south-striking trend (N160 –308E) with symmetrical (Talat n’Issy and Fouanou) (Fig. 6c) or asymmetrical (Issafene) synclines. In the lowermost Cambrian pelitic strata, north–south-trending, decamete-scale folds containing a vertical axialplane cleavage are well developed. These folds progressively vanish upward, indicating a de´collement level between the Cambrian limestones and the overlying, mainly sandy strata. (3) Between the Irherm inlier and the Wawfengha inlier, N70 –1108E anticlinal axes, bedding planes and decametre-scale fold axes in the Lower Cambrian rocks suggest a sinistral shear controlled by regional east –west sinistral faults (Fig. 6b).
441
(4) North of the Wawfengha inlier, the Lower Cambrian limestones rocks are not folded, but they are cut by east– west subvertical faults and tilted to the north, and finally concealed below the Quaternary alluvium of the Souss plain. Most probably, part of these faults operated during Atlas deformation.
The Bou Azzer – El Graara area In the central part of the Anti-Atlas (Fig. 1), the Bou Azzer –El Graara inlier is an eroded NW–SE Variscan basement high with a box-fold shape that borders the Anti-Atlas Major Fault (Choubert 1947) (Fig. 7). The main part of the Proterozoic basement consists of a Pan-African dismembered ophiolitic sequence and arc fragments (Leblanc 1975; Saquaque et al. 1989). The Pan-African structures are unconformably overlain by a thick Neoproterozoic to Cambrian volcano-sedimentary cover. At its base, the folded Tiddiline Formation attributed to the ‘Saghro Group’ (Thomas et al. 2004) is confined to NE –SW-trending late Pan-African basins (Hefferan et al. 1992). The overlying ‘Ouarzazate Supergroup’ volcanoclastic sequence, which ranges in thickness from 0 to 700 m, is the highest exposed stratigraphical unit. The Ouarzazate sequences form impressive cliffs around the Proterozoic basement high. These latest Neoproterozoic sequences are disconformably overlain by the Lower Cambrian carbonates interlayered with the alkaline syenitic Alougoum ´ lvaro et al. 2006), dated at volcanic flows (A 534 + 10 Ma (Ducrot & Lancelot 1977) and 529 + 3 Ma (Gasquet et al. 2005). The Late Palaeozoic compressional event reactivated the basement structures along the borders of the inlier (Fig. 7c). The basement rocks were not deformed, except in the vicinity of the previous fractures, where kink bands and a rough cleavage developed. In contrast, folds of varying wavelengths are very common in the cover rocks. The regional anticlines are characterized by box-fold shapes throughout the Bou Azzer– El Graara area, and by large open synclines of Cambrian rocks. Metre- to decametre-scale, upright detachment folds are common in the lowermost Cambrian limestones. The folds exhibit a dominant NW– SE trend with subordinate NE –SW structures. These folds are often conical (Leblanc 1975) with gently plunging axes and steep axial planes inclined towards the margins of the inlier. Such folds are often associated with reverse faults, which possibly originated from the inversion of former synsedimentary normal faults. Upwards, these folds disappear progressively in the overlying sedimentary sequences. Cleavage is either absent or occurs as a rough spaced cleavage in the pelitic layers of the
442
A. SOULAIMANI & M. BURKHARD
Fig. 6. (a) Location of the Irherm–Tata area in the Anti-Atlas belt. (b) Structural map of the Irherm–Tata area and (c) geological cross-section of the Irherm–Tata area (section location D– D0 in the map). The stereograms (Wulff stereonet, lower hemisphere) represent bedding (S0) and foliation (S1) data in the Talat N’Ouamane domain (dashed inset).
THE ANTI-ATLAS CHAIN, MOROCCO
443
Fig. 7. (a) Location of the Bou Azzer–El Graara area in the Anti-Atlas belt. (b) Geological and structural map of the Bou Azzer– El Graara area, modified from Leblanc (1975) and (c) geological cross-section (section location E– E0 in the map).
Cambrian synclines. Along the northeastern margin of the inlier, the folds are associated with kilometre-scale sinistral strike-slip faults and SW-verging transpressional reverse faults (J. El Hassel), where the ‘Lower Limestones’ formation is stacked upon the ‘Tikirt Sandstones’, equivalent to the ‘Lie-de-vin series’ rocks. The southwestern border of the inlier is defined by a sharp, steep monocline, exposed on the scale of tens of metres. The NW–SE direction of the Bou Azzer – El Graara antiform continues southeastward as far as the southern Tafilalt area and the Algerian ‘aulacogenic’ Ougarta chain, where the Variscan tectonic style has similar geometry and structural features to that of the eastern Anti-Atlas (Donzeau 1974).
The Saghro – Ougnate area The last example chosen to illustrate the Variscan tectonics in the Anti-Atlas belt comes from the vicinity of the Tinghir oasis, just south of the High Atlas (Fig. 1). The Cambrian to Carboniferous strata cropping out along the northern border of the Saghro and Ougnate Proterozoic massifs (Fig. 8b) are deformed by several thrust faults that dip gently northward (Hindermeyer 1954; Choubert 1959). Michard et al. (1982) presented a detailed north–south cross-section of the region south of Tinghir (Fig. 8c), where they distinguished
three allochthonous units overriding a southern autochthonous unit apparently undetached from the Sahgro substratum. The allochthonous units display have sedimentary features to those of the northern margin of the Saghro massif. Lower Ordovician and Silurian shales constitute two ductile de´collement levels accommodating SSEdirected thrusting. A rough, gently north-dipping, east –west-trending cleavage is locally axial planar to south-vergent asymmetrical folds. The cleavage affects all of the Palaeozoic units but is preferentially developed within the pelitic horizons. This cleavage also occurs within the latest Proterozoic volcanoclastic formation along the northern border of the Saghro massif. The cleavage progressively dies out northward and disappears completely in the J. Tisdafine Carboniferous flysch units, which are, however, affected by east –west-trending folds associated with thrusting toward the south (Soualhine et al. 2003). Southvergent thrusts are also developed in the Neoproterozoic cover along the northern side of the Ougnate inlier. To the south, between the Saghro and Ougnate massifs, an uplifted Cambrian outcrop, about 15 km wide, is squeezed between the Saghro and the Ougnate basement blocks. These Cambrian strata between the two Precambrian massifs are locally overprinted by a north – south rough cleavage, noticeable in the Cambrian
444
A. SOULAIMANI & M. BURKHARD
Fig. 8. (a) Location of the Saghro– Ougnate area in the Anti-Atlas belt. (b) Structural map of Saghro– Ougnate area, modified from the 1:500 000 geological map of Ouarzazate (Choubert 1959) and (c) geological crosssection of the northeastern border of the Saghro massif (section location F– F0 in the map), reinterpreted from Michard et al. 1982).
Paradoxides-bearing shales (Fig. 8b). To the south, in the Alnif plain, Ordovician rocks have a south-dipping monoclinal geometry similar to that observable along the southern flank of the central and eastern Anti-Atlas chain. Generally, similar southern-vergent structures are described in the Palaeozoic inliers of the High Atlas, south of the ‘South Atlas Fault’, in the Skoura – Aı¨t Tamlil (Jenny & Le Marrec 1980) and in the Bechar basin, where overlying strata are post-Namurian in age (Ball et al. 1975).
Discussion Para-autochthonous v. autochthonous Anti-Atlas The present study suggests that the Variscan AntiAtlas chain consists of two structural domains, as follows.
To the west, the narrow western Atlantic Bas Draˆa area was subjected to a penetrative deformation (Mattauer et al. 1972; Soulaimani 1998; Belfoul et al. 2001). The Cambrian series has been intensely deformed by east –west- to NW – SE-trending large recumbent folds. The axial planes of these folds are gently west-dipping or even subhorizontal in the westernmost outcrops. The folded units are composed of thrust sheets or duplexes. The deformation intensity decreases progressively eastward. The analysis of lineations and kinematic indicators, the occurrence of large recumbent folds with horizontal foliations, and the widespread development of low-angle faults indicate a horizontal displacement toward the ESE. In this area, there is no evidence for any basement involvement, at least within the exposed supracrustal levels. This thin-skinned tectonic style suggests that the westernmost part of the Anti-Atlas Palaeozoic cover can be regarded as a paraautochthonous (rather than an allochthonous) area
THE ANTI-ATLAS CHAIN, MOROCCO
that is the northern prolongation of the Mauritanides belt. This is the only part of the Anti-Atlas belt that is clearly part of the external zones of the Variscan belt. The remaining areas of the Anti-Atlas further east should be considered as meta-cratonic rather than part of the Variscan mobile belt (Burkhard et al. 2006). Throughout the remaining western and central Anti-Atlas chain, NE– SW-trending Precambrian blocks were deeply buried beneath up to 10 km of Palaeozoic sediments. This Palaeozoic basin was subsequently reactivated during the Late Palaeozoic compressive event (Soulaimani et al. 1997; Helg et al. 2004; Burkhard et al. 2006). The geometry of the Variscan basement uplifts or ‘boutonnie`res’ is controlled by inherited fracture zones of the underlying Precambrian basement. Variscan folding within the Palaeozoic cover series is in turn strongly influenced by the inversion of these basement blocks. Deformation intensity is generally greatest at the base of the cover, near the reactivated basement structures. Folding intensity decreases southeastward as well as upward. Steep thrust faults are rarely associated with this compressive event, they are observed only in some places along the borders of the basement inliers. Cover shortening is almost exclusively accommodated by upright folding rather than by duplex formation or thrusting (Caritg et al. 2004; Helg et al. 2004). Thus, the greater part of the Anti-Atlas is clearly an autochthonous thick-skinned tectonic province characterized by brittle to semi-brittle reactivation of the Precambrian basement and clear decoupling of the deformation mode between the basement and Palaeozoic cover. This tectonic style mostly related to vertical uplift of basement blocks is more typical for a foreland area rather than the Variscan external thrust-and-fold belt.
Basement – cover relationships The structural pattern of the Anti-Atlas, except in its westernmost part, is characterized by the presence of Proterozoic inliers separated by wide flat-floored synclines cored by Cambrian sedimentary rocks. A map-scale outcrop pattern combined with structural data indicates a large-magnitude bend in the structural grain of the Anti-Atlas wrapping around the West African craton. In the western Anti-Atlas, the inliers are oriented NE– SW to north–south; in the central and eastern Anti-Atlas, the structural trends are east –west (Irherm, Saghro) and NW– SE (Bou Azzer –El Graara), belonging to the Ougarta trends. On the whole, the Anti-Atlas and Ougarta domains form a broad arc north of the Reguibat shield (Lefort 1988; Haddoumi et al. 2001) (Fig. 9).
445
It is commonly accepted that the variation in the structural trend of Variscan folds and the shape of inliers throughout the Anti-Atlas are inherited from Proterozoic basement trends (Choubert 1947; Leblanc 1972, 1975; Donzeau 1974; Michard 1976; Jeannette & Pique´ 1981; Soulaimani et al. 1997). Many arguments and observations suggest that the pattern and geometry of the inliers, and even the orientation of their faulted borders, often at right angles, are controlled by ancestral Precambrian basement structures. The best example is that of the Bou Azzer –El Graara boutonnie`re. The Late Pan-African orogeny, and the subsequent rift-related extensional event, controls the Palaeozoic and younger structural trends (Pique´ et al. 1999; Soulaimani et al. 2003). However, a precise synthetic reconstruction of the Late Proterozoic basement –cover configuration has not yet been compiled. The Late Proterozoic–Early Cambrian extensional event provides important constraints on the subsequent development of Variscan structures. Although parts of some Precambrian inliers may be inverted basins filled in with the synrift ‘Ouarzazate Supergroup’ deposits (Faik et al. 2001; Helg et al. 2004), many others were already single highs during the Late Proterozoic continental rifting, as shown by the development of surrounding collapse fault and by the deposition of the post-rift Cambrian sequences directly over their Proterozoic basement. Many inliers in the western Anti-Atlas can be taken as examples of this configuration (Bas Draˆa, Kerdous, Tagragra Akka, Tagragra Tata, etc.). Furthermore, the mechanism of the Late Proterozoic extensional event in the Anti-Atlas is more complex than a classical tilting of basement rigid blocks, if we consider the coeval expanding magmatic and hydrothermal activities affecting the Precambrian basement. As an example, the eastern Kerdous inlier could correspond to a diapiric metamorphic dome exhumed during the Late Precambrian transtensional event (Soulaimani & Pique´ 2004). The other western Anti-Atlas inliers provide evidence for tectonothermal restructuring of their Precambrian basement during this extensional event (Oudra et al. 2006). Although the mechanism of this early extensional event is a matter of debate, it is obvious that the resulting basement –cover configuration has a significant role in subsequent reactivations of these structures during the Variscan orogeny. The mechanism governing the development of basement uplifts during the Variscan epoch seems to be more complex than simple classical positive inversion (Williams et al. 1989) of Late Proterozoic – Early Cambrian extensional grabens. The uplift of the Precambrian basement blocks as a result of the Variscan shortening, as shown by the geometric
446
A. SOULAIMANI & M. BURKHARD
Fig. 9. Position of the Anti-Atlas and related areas in the limit Palaeozoic–Mesozoic, modified from Burkhard et al. (2006). (a) General map of the African– North American system in the Early Mesozoic from Sahabi et al. (2004), with the addition of the main structures of the Late Palaeozoic Alleghanian–Gondwanian collision. (b) Global plate tectonic reconstruction (in the middle) and cross-sections showing Variscan– Alleghanian structures in the Anti-Atlas belt modified from Burkhard et al. (2006). 1, West African craton and Anti-Atlas cratonic basement uplift; 2, North American craton; 3, Acadian chain; 4, Avalon terrane; 5, Meguma terrane; 6, Proterozoic basement uplift; 7, allochthonous Mauritanide– western Anti-Atlas terranes; 8, Palaeozoic cover; 8, Post-Palaeozoic cover.
array of the post-rift Early Cambrian Limestone deposits around the Precambrian inliers and supported by kinematic analyses (e.g. Bas Draˆa and Lakhssas Plateau), is superimposed upon the Late Precambrian initial horst shape. It is therefore difficult to provide quantitative estimates of the amount of Variscan shortening and uplift for single basement blocks. At Lakhsass, the difference in elevation between the top of the uplifted basement massif (J. Lkest, Kerdous) and the bottom of the adjacent Lakhssas basin is at least 4 km. On the scale of the entire Anti-Atlas anticlinorium such assessments of the vertical movements become even more complex, as we also have to take into account a significant Late Miocene Atlas– Alpine uplift affecting all of the Anti-Atlas belt. Proterozoic basement blocks subjected to the Variscan compression behaved rigidly and do not
display any evidence for internal deformation except in peripheral zones. The rheological difference between the metamorphic and granitic basement rocks and the layered cover sediments may have played a significant role, as the cover rocks are more easily deformed than their basement. The shortening of the basement was primarily accommodated by the reactivation of ancient fractures, whereas cover rocks are mostly shortened by poly-harmonic folding without any significant faulting and thrusting. Penetrative plastic deformation with the development of a foliation is restricted to the margins of the uplifted basement (e.g. Bas Draˆa) or to limited areas between basement blocks (e.g. Lakhssas Plateau). Around the Bas Draˆa inlier, folded limestones several hundred metres thick are clearly detached from the weak infra-Adoudounian basal sandstones.
THE ANTI-ATLAS CHAIN, MOROCCO
Such detachments occur at all levels throughout the Palaeozoic succession but are typical of the frequent shaly levels within the Ordovician and Silurian rocks, especially in the western Anti-Atlas. They account for the marked changes in style, wavelengths and amplitudes of folds within the Palaeozoic cover (Helg et al. 2004). The uppermost de´collement level is postulated within the Late Devonian marls to explain the apparent ‘unconformity’ below the J. Ouarkziz Carboniferous strata, which remained undeformed above the folded J. Rich Devonian sandstones. This latter de´collement can be tentatively interpreted as a regional-scale back-thrust (NNW thrusting) at the base of J. Ouarkziz (Burkhard et al. 2001; Helg et al. 2004), although other researchers have suggested the occurrence of a diffuse dextral shear zone (Michard 1976). Given the lack of geophysical data, these interpretations remain provisional, as they are not yet corroborated by any field evidence. Along the southern flank of the western AntiAtlas anticlinorium (J. Bani), some 15–20% of horizontal shortening was deduced from the restoration of Devonian and Ordovician folds in the Tata zone (Caritg et al. 2004; Helg et al. 2004). Deformation increases in the western Atlantic Bas Draˆa, where a minimum rate of 45% of shortening is observed within the units that crop out. Elsewhere in the central and eastern Anti-Atlas, values of 5–10% at most for the Variscan shortening have been calculated (Leblanc 1975; Hassenforder 1987), except for the thrust-dominated area of Tineghir, where some 40% of shortening has been postulated (Michard et al. 1982). Because of the lack of unconformably overlying strata, the age of the folding and synchronous basement uplift remains poorly constrained. It is certainly Variscan (Bonhomme & Hassenforder 1985), as folds and thrusts affect Vise´an strata in the Tineghir area (eastern Saghro; Michard et al. 1982), as well as the south Ougnat domain, and as folds of the entire Anti-Atlas belt are crosscut by Late Triassic dolerite dykes. The regional (western Anti-Atlas) structural contrast between the folded Early and Middle Palaeozoic strata and the mostly undisturbed Carboniferous sequences is not a stratigraphic unconformity. Instead, it represents a gradual or faulted transition from the weakly deformed domain of the southern AntiAtlas towards the mostly undeformed Carboniferous Ouarkziz sequences at the northern border of the Tindouf platform (Soulaimani et al. 1997). Furthermore, many Anti-Atlas rocks have experienced a Variscan lower-greenschist facies metamorphism at 150–300 8C (Buggisch 1988; Soulaimani 1998; Burkhard et al. 2001), which is explained by deep sedimentary burial beneath a 10 km thickness, or more, of Palaeozoic overburden.
447
The southern limit of the Variscan belt As described above, the thin-skinned tectonic style observed along the Atlantic coast differs considerably from the thick-skinned tectonics prominent throughout the Anti-Atlas chain, where the involvement of basement blocks plays a critical role during the Variscan compressive deformations. At a larger scale, the Anti-Atlas appears as the nonsymmetrical hinterland to the Alleghanian foreland fold– thrust belt on the American side of the Appalachian chain (Fig. 9). The east-vergent thrusts of the western Atlantic Bas Draˆa area correspond to the eastern side of the Mauritanides back-thrust terranes and both can easily be seen as the eastern Appalachian mountain front. The Anti-Atlas belt, however, corresponds to the SE metacratonic foreland to the Variscan belt in NW Africa. Indeed, as defined by Abdelsalam et al. (2002), the metacraton corresponds to a craton that has been at least remobilized during an orogenic event, but that is still rheologically recognizable. This is the case for the Anti-Atlas belt and overall the NW border of the West African craton, which are visibly affected by major basement uplifts and inversion tectonics during the Variscan orogeny, after the Pan-African orogeny and subsequent extensional evolution (Coward & Ries 2003; Burkhard et al. 2006). To the north, the Moroccan Variscan belt is represented by the so-called Meseta domain (Pique´ 1989; Pique´ & Michard 1989; Hoepffner et al. 2005). Within the Meseta, the eastern zones differ from the western ones by their early metamorphic evolution (Hoepffner 1987). In the western Meseta, the deformation is heterogeneous and concentrated within sheared zones (Lagarde et al. 1990; Essaı¨fi et al. 2001). As in the Anti-Atlas, a contrast exists in the Meseta between homogeneously and more strongly deformed zones on the one hand (eastern Meseta), and heterogeneously deformed zones, where the deformation is concentrated along block limits (western Meseta) on the other hand. It is important to emphasize significant differences between the Anti-Atlas and the Meseta (Michard 1976; Pique´ & Michard 1989; Hoepffner et al. 2005): (1) generally, the orogenic shortening is much more important in the Meseta than in the Anti-Atlas chain; also, the development of a regional metamorphism, sometimes reaching medium grade (upper greenschist to amphibolite facies), and the emplacement of Variscan granitoids, which are never represented in the Anti-Atlas, are noticeable in the Meseta domain; (2) the dominantly northwestward vergence observed in the eastern Meseta zones differs from the SE-directed structures of westernmost ‘Atlantic’ domains of the Anti-Atlas.
448
A. SOULAIMANI & M. BURKHARD
These fundamental differences between the two segments of the Variscan belt of Morocco call for a basic role of the ‘South Atlas Fault’ as a major Variscan lithospheric fault (Mattauer et al. 1972; Ouanaimi & Petit 1992). This Palaeozoic expression of the South Atlas Fault has been named the Atlas Palaeozoic Transform Fault by Pique´ & Michard (1989). A recent plate tectonic interpretation (Stampfli & Borel 2002) suggests that the Moroccan Meseta was rifted and drifted away from Gondwana at c. 490 Ma. However, many geological observations indicate that the Meseta domain belongs to, or at least was located next to, the West African craton during the Palaeozoic times (Dostal et al. 2005; Hoepffner et al. 2005). The Meseta corresponds to relatively distal parts of the Africa passive margin, which remained close to the Anti-Atlas throughout Palaeozoic times. The present-day position results from the Late Carboniferous–Early Permian collage of the Meseta domain against and onto the Anti-Atlas domain. The traces of this collage are widely exposed in the Tineghir area as folds and thrusts (Michard et al. 1982) and as dextral wrench zones in the Tamlelt area (Houari & Hoepffner 2003). Further west, it is coincident with the Mesozoic –Cenozoic Tizi n’Test-Meltsen fault system. The ‘South Atlas Fault’ zone is a major Late Carboniferous –Early Permian intracontinental fault devoid of any ophiolitic slivers. The strong contrasts in Variscan deformation between the Anti-Atlas and Meseta domains, both in intensity and dominant vergence, imply that their mutual boundary operated mostly as a wrench system. Toward the south, the southwesternmost AntiAtlas belt correlates with the Zemmour belt, considered as the foreland of the northern Mauritanides chain (Sougy 1969; Le´corche´ et al. 1991). In the western Zemmour belt, Devonian units are folded and thrust to the east, where Lower Palaeozoic sediments unconformably overlie the Reguibat shield. Further south, the allochthonous Oulad Dlim area (Adrar Soutouf) consists of crystalline nappes of Pan-African age in the west (Le Goff et al. 2001; Villeneuve 2005; Villeneuve et al. 2006), which were thrust during the Variscan orogeny and override eastward a Palaeozoic sedimentary series (Ordovician to Devonian). Finally, because of its position along the southern border of the Variscan belt, the Anti-Atlas chain constitutes an example of cratonic basement blocks with, at its western side, a cratonward vergence. It fundamentally differs from the northern limit of the Variscan belt in Europe, where the orogenic boundary is a crustal-scale thrust toward the North European craton and its Caledonian margin (Matte 1986). Variscan deformation, analysed above, is different from that of foreland belts
and associated collisional chains such as the Appalachian Valley and Ridge province (Hatcher & Odom 1980; Rodgers 1995) or the Alpine External Crystalline Massifs (Boyer & Elliot 1982). The involvement of basement blocks throughout the Anti-Atlas chain best compares with the Wind River basement uplifts of the Rocky Mountains, which developed far beyond the orogenic front of this mountain belt. Rodgers (1987, 1995) in his comparative anatomy of mountain chains called this style ‘basement uplifts within cratons marginal to orogenic belts’ in contrast to the more common ‘basement uplifts within external parts of orogenic belts’.
Conclusion The main tectonic features of the Anti-Atlas are summarized briefly below. (1) The Anti-Atlas consists of the juxtaposition of two Variscan domains affected by contrasting tectonic styles: (1) a western ‘Atlantic’ domain (para-autochthonous Anti-Atlas) marked by flat foliations, axial planar to large-scale recumbent folds, and east-verging thrusts; (2) a vast folded domain (autochthonous Anti-Atlas) characterized by upright folds of wavelengths varying from less than 100 m to more than 10 km with broad Precambrian basement-involved anticlines. (2) Basement blocks limited by Pan-African reactivated fractures controlled the Variscan deformation within the cover. The vertical uplift of basement inliers in excess of 4 km is the outstanding feature of this tectonic inversion. (3) The deformation intensity within the cover is strongly influenced by the geometry of the basement uplifts. Deformation is most intense near the steep borders of the basement blocks or between closely spaced basement blocks, where a more or less penetrative cleavage develops. Around the basement inliers, bedding and the geometry of folds are strongly influenced by the relative displacements of the underlying rigid block. (4) Shortening within the Palaeozoic cover is decoupled from the basement deformation and accommodated by upright folds of different wavelengths. Strong disharmony is accommodated by a multitude of detachment levels. Deformation intensity generally decreases toward the south and SE and completely vanishes within the stable Tindouf platform. (5) The Anti-Atlas belt represents a particular style of deformation not seen elsewhere in the Variscan chain of Northern Africa or Europe. Basement uplifts and strong tectonic inversion are observed all along the northern border of the West African metacraton. This metacraton occupies a key position
THE ANTI-ATLAS CHAIN, MOROCCO
between the northern domains of Morocco (Meseta Block) and the Mauritanides belt to the south. Together, these orogenic belts define a broadly synchronous circum-Atlantic Variscan –Alleghanian belt developed during the Late Palaeozoic Gondwana–Laurentia collision. Professor Martin Burkhard died on 23 August 2006, when sampling in the Alps. This paper, to which he significantly contributed, is dedicated to his memory. Thanks are due to A. Pique´ and A. Michard for fruitful discussions and a careful review of early versions, and to A. Azor and S. Mazur for constructive reviews of the manuscript.
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 DMOU , H. & J UTEAU , T. 2000. De´couverte d’un syste`me hydrothermal oce´anique fossile dans l’ophiolite ante´cambrienne de Khezama (massif du Siroua, Anti-Atlas marocain). Comptes Rendues de l’Acade´mie des Sciences, 327, 335–340. A¨I T M ALEK , H., G ASQUET , D., B ERTRAND , J. M. & L ETERRIER , J. 1998. Ge´ochronologie U –Pb sur zircon de granitoı¨des e´burne´ens et panafricains dans les boutonnie`res d’Irherm, du Kerdous et du Bas Draˆa (Anti-Atlas occidental, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 327, 819– 826. A LGOUTI , AB ., A LGOUTI , AH ., B EAUCHAMPS , J., C HBANI , B. & T AJ -E DDINE , K. 2002. Pale´ge´ographie d’une plateforme infracambrienne en dislocation: se´rie de base adoudounienne de la re´gion Waoufengha– Igherm, Anti-Atlas occidental, Maroc. Comptes Rendus de l’Acade´mie des Sciences, 330, 155– 160. A LVARO , J. J., E ZZOUHAIRI , H., V ENNIN , E. ET AL . 2006. The Early-Cambrian Boho volcano of the El Graara massif, Morocco: Petrology, geodynamic setting and coeval sedimentation. Journal of African Earth Sciences, 44, 396–410. A ZIZI S AMIR , M. R., F ERRANDINI , J. & T ANE , J. L. 1990. Tectonique et volcanisme tardi-Pan Africains (580–560 M.A.) dans l’Anti-Atlas central (Maroc): Interpre´tation ge´odynamique a` l’e´chelle du NW de l’Afrique. Journal of African Earth Sciences, 10, 549–563. B ALL , E., F ABRE , J., G UELLAL , S., M E´ GARD , F. & M OUSSINE -P OUCHKINE , A. 1975. Sur la pre´sence de cisaillement plats d’aˆge hercynien dans le Carbonife`re de Be´char (Alge´rie). Comptes Rendus de l’Acade´mie des Sciences, 280, 2721– 2724. B ELFOUL , A., F AIK , F. & H ASSENFORDER , B. 2001. Mise en e´vidence d’une tectonique tangentielle ante´rieure au plissement majeur dans la chaıˆne hercynienne de l’Anti-Atlas occidental, Maroc. Journal of African Earth Sciences, 32, 723–739. B ENSSAOU , M. & H AMOUMI , N. 2003. Le graben de l’Anti-Atlas occidental (Maroc): controˆle tectonique de la pale´oge´ographie et des se´quences au Cambrien infe´rieur. Comptes Rendus Ge´oscience, 335, 297– 305.
449
B ONHOMME , M. & H ASSENFORDER , B. 1985. Le me´tamorphisme hercynien dans les formations tardiet post-panafricaines de l’Anti-Atlas occidental (Maroc). Donne´es isotopiques Rb/Sr et K/Ar des fractions fines. Sciences Ge´ologiques, 38, 175–183. B OURCART , J. 1937. L’anticlinal ante´cambrien du Bas Draˆa (Sahara marocain). Comptes Rendus Sommaires de la Socie´te´ Ge´ologique de France, 242–244. B OYER , S. E. & E LLIOT , D. 1982. Thrust systems. AAPG Bulletin, 66, 1196–1230. B OYER , C., C HIKHAOUI , C., D UPUY , M. & L EBLANC , M. 1978. Le volcanisme calco-alcalin pre´cambrien terminal de l’Anti-Atlas (Maroc) et ses alte´rations, Interpre´tation ge´odynamique. Comptes Rendus de l’Acade´mie des Sciences, 287, 427– 430. B UGGISCH , W. 1988. Diagenesis and very low-grade metamorphism of the lower Cambrian rocks in the Anti-Atlas (Morocco). Lecture Notes in Earth Sciences, 15, 123– 128. B UGGISCH , W. & S IEGERT , R. 1988. Paleogeography and facies of the ‘Gres Terminaux’ (Uppermost Lower Cambrien, Anti-Atlas/Morocco). Lecture Notes in Earth Sciences, 15, 107 –121. B URKHARD , M., C ARITG , S. & H ELG , H. 2001. Forced, disharmonic multilayer buckle folding in the late Variscan Antiatlas of Morocco. AAPG Bulletin, 85, 3– 6. M., C ARITG , S., H ELG , U., B URKHARD , R OBERT -C HARRUE , CH . & S OULAIMANI , A. 2006. Tectonics of the Anti-Atlas system. Comptes Rendus Ge´oscience, Special Volume, 338, 11– 24. C ARITG , S., B URKHARD , M., D UCOMMUN , R., H ELG , U., K OPPO , L. & S UE , C. 2004. Fold interference patterns in the late Palaeozoic Anti-Atlas belt of Morocco. Terra Nova, 16, 27–37. C HARLOT , R. 1978. Caracte´risation des e´ve´nements e´burne´ens et panafricains dans l’Anti-Atlas marocain. Apport de la me´thode ge´ochronologique Rb/Sr. PhD thesis, Rennes University. C HOUBERT , G. 1947. L’accident majeur de l’Anti-Atlas. Comptes Rendus de l’Acade´mie des Sciences, 224, 1172– 1173. C HOUBERT , G. 1952. Histoire ge´ologique du domaine de l’Anti-Atlas. In: Ge´ologie du Maroc. Notes et Me´moires du Service Ge´ologique du Maroc, 100, 75–194. C HOUBERT , G. 1959. Carte ge´ologique du Maroc au 1/500 000, feuille Ouarzazate. Service Ge´ologique du Maroc. C HOUBERT , G. 1963. Histoire ge´ologique du Pre´cambrien de l’Anti-Atlas. Tome I. Notes et Me´moires du Service Ge´ologique du Maroc, 162. C HOUBERT , G. & F AURE -M URET , A. 1969. Sur la se´rie stratigraphique pre´cambrienne de la partie sud-ouest du massif du Bas Dra (Tarfaya, Sud marocain). Comptes Rendus de l’Acade´mie des Sciences, 269, 759– 762. C LAUER , N. 1974. Utilisation de la me´thode Rb– Sr pour la datation d’une schistosite´ de se´diments peu me´tamorphise´s: application au Pre´cambrien II de la boutonnie`re de Bou-Azzer– El Graara (Anti-Atlas, Maroc). Earth and Planetary Science Letters, 22, 404– 412. C OWARD , M. P. & R IES , A. C. 2003. Tectonic development of North African basins. In: A RTHUR , T., M AC G REGOR , D. S. & C AMERON , N. R. (eds) Petroleum Geology of Africa; New Themes and
450
A. SOULAIMANI & M. BURKHARD
Developing Technologies. Geological Society, London, Special Publications, 207, 61–83. D ESTOMBES , J., H OLLARD , H. & W ILLEFERTS , S. 1985. Lower Palaeozoic rocks of Morocco. In: H OLLAND , C. H. (ed.) Lower Palaeozoic Rocks of the World, Vol. 4, Lower Palaeozoic of North-Western and West-Central Africa. Wiley, Chichester, 91– 336. D OBLAS , M., L OPEZ -R UIZ , J., C EBRIA , J. M., Y OUBI , N. & D EGROOTE , E. 2002. Mantle insulation beneath the West African craton during the Precambrian – Cambrian transition. Geology, 30, 839–842. D ONZEAU , M. 1974. L’Arc de l’Anti-Atlas-Ougarta (Sahara nord-occidental, Alge´rie –Maroc). Comptes Rendus de l’Acade´mie des Sciences, 278, 417–420. D OSTAL , J., K EPPIE , J. D., H AMILTON , M. A., A ARAB , E. M., L EFORT , J. P. & M URPHY , J. B. 2005. Crustal xenoliths in Triassic lamprophyre dykes in western Morocco: tectonic implications for the Rheic Ocean suture. Geological Magazine, 142, 159–172. D UCROT , J. & L ANCELOT , J. R. 1977. Proble`me de la limite Pre´cambrien –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 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 SSAI¨ FI , A., L AGARDE , J. L. & C APDEVILLA , R. 2001. Deformation and displacement from shear zone patterns in the Variscan upper crust, Jebilet, Morocco. Journal of African Earth Sciences, 32, 335– 350. F AIK , F., B ELFOUL , M. A., B OUABDELLI , M. & H ASSENFORDER , B. 2001. Les structures de la couverture ne´oprote´rozoı¨que terminal et pale´ozoı¨que de la re´gion de Tata, Anti-Atlas centre-occidental, Maroc: de´formation polyphase´e, ou interactions socle/couverture pendant l’orogene`se hercynienne? Journal of African Earth Sciences, 32, 765– 776. F RIZON DE L AMOTTE , D., S AINT B EZAR , B., B RACE` NE , E. & M ERCIER , E. 2000. The two main steps of the Atlas building and geodynamics of the western Mediterranean. Tectonics, 19, 740– 761. G ASQUET , D., C HEVREMONT , P., B AUDIN , T. ET AL . 2004. Polycyclic magmatism in the Tagragra d’Akka and Kerdous–Tafeltast inlier (Western Anti-Atlas, Morocco). Journal of African Earth Sciences, 39, 267–275. 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 ENTIL , L. 1918. Notices sur les titres et travaux scientifiques de L. Gentil. Larose, Paris. H ADDOUMI , H., G UIRAUD , R. & M OUSSINE P OUCHKINE , A. 2001. Hercynian compressional deformations of the Ahnet–Mouydir Basin, Algerian Sahara platform: far-field stress effects of the Late Palaeozoic orogeny. Terra Nova, 13, 220– 226. H ASSENFORDER , B. 1987. La tectonique panafricaine et varisque de l’Anti-Atlas dans le massif de Kerdous
(Maroc). PhD thesis, Louis-Pasteur University, Strasbourg. H ATCHER , R. D. & O DOM , A. L. 1980. Timing of thrusting in the southern Appalachians, USA: model for orogeny. Journal of the Geological Society, London, 137, 321– 327. H EFFERAN , K., K ARSON , J. A. & S AQUAQUE , A. 1992. Proterozoic collisional basin in a Pan-African suture zone, Anti-Atlas Mountains, Morocco. Precambrian Research, 54, 295– 319. H EFFERAN , K., A DMOU , H., K ARSON , J. A. & S AQUAQUE , A. 2000. Anti-Atlas (Morocco) role in Neoproterozoic Western Gondwana reconstitution. Precambrian Research, 103, 89– 96. H ELG , U., B URKHARD , M., C ARITG , S. & R OBERT -C HARRUE , C. 2004. Folding and inversion tectonics in the Anti-Atlas of Morocco. Tectonics, 23, 17, paper number TC4006. H INDERMEYER , J. 1954. De´couverte du Tournaisien et tectonique premonitoire hercynienne dans la re´gion de Tinghir. Comptes Rendus de l’Acade´mie des Sciences, 239, 1824–1826. H OEPFFNER , C. 1987. La tectonique hercynienne dans l’Est du Maroc. PhD thesis, Louis-Pasteur University, Strasbourg. H OEPFFNER , C., S OULAIMANI , A. & P IQUE , A. 2005. The Moroccan Hercynide. Journal of African Earth Sciences, 43, 144–165. H OUARI , M. R. & H OEPFFNER , C. 2003. Late Carboniferous dextral wrench-dominated transpression along the North African craton margin (Eastern High-Atlas, Morocco). Journal of African Earth Sciences, 37, 11–24. J EANNETTE , D. & P IQUE´ , A. 1981. Le Maroc hercynien: plate-forme disloque´e du craton ouest-Africain. Comptes Rendus de l’Acade´mie des Sciences, 293, 79–82. J ENNY , J. & L E M ARREC , A. L. 1980. Tectonique de la boutonnie`re d’Aı¨t Tamlil. Haut Atlas central. Mines, Ge´ologie et Energie, 48, 38–44. L AGARDE , J. L., A¨I T O MAR , S. & R ODAZ , B. 1990. Structural characteristic of granitic plutons emplaced during weak regional deformation: Examples from late Carboniferous plutons. Journal of Structural Geology, 12, 805– 821. L EBLANC , M. 1972. Sur le style disharmonique des plis hercyniens de la couverture, dans l’Anti-Atlas central (Maroc). Comptes Rendus de l’Acade´mie des Sciences, 275, 803– 806. L EBLANC , M. 1975. Ophiolites pre´cambriennes et gıˆtes arse´nie´s de Cobalt (Bou-Azzer, Maroc). Notes et Me´moires du Service Ge´ologique du Maroc, 280. L EBLANC , M. & L ANCELOT , J. R. 1980. Interpre´tation ge´odynamique du domaine pan-Africain (Pre´cambrien terminal) de l’Anti-Atlas (Maroc) a` partir des donne´es ge´ologiques et ge´odynamiques. Canadian Journal of Earth Sciences, 17, 142– 155. L ECOINTRE , G. 1926. Recherches ge´ologiques dans la Meseta marocaine. Notes et Me´moires du Service Ge´ologique du Maroc, 14. L E´ CORCHE´ , J. P., B RONNER , G., D ALLMEYER , R. D., R OCCI , G. & R OUSSEL , J. 1991. The Mauritanide Orogen and its northern extensions (Western Sahara and Zemmour), West Africa. In: D ALLMEYER , R. D. &
THE ANTI-ATLAS CHAIN, MOROCCO L E´ CORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives, Springer, Berlin, 187–227. L EFORT , J. P. 1988. Imprint of the Reguibat uplift (Mauritania) on the central and southern Appalachians of the U.S.A. Journal of African Earth Sciences, 7, 433– 442. L E G OFF , E., G UERROT , C., M AURIN , G., J OHAN , V., T EGYEY , M. & B EN Z ARGA , M. 2001. De´couverte d’e´clogites hercyniennes dans la chaıˆne septentrionale des Mauritanides (Afrique de l’Ouest). Comptes Rendus de l’Acade´mie des Sciences, 333, 711– 718. L E R OY , P. & P IQUE´ , A. 2001. Triassic–Liassic Western Moroccan synrift basins in relation to the Central Atlantic opening. Marine Geology, 172, 359– 381. Maroc Service Ge´ologique 1985. Carte ge´ologique du Maroc, scale 1/1 000 000. Notes et Me´moires du Service Ge´ologique du Maroc, 260. M ATTAUER , M., P ROUST , F. & T APPONNIER , P. 1972. Major strike-slip fault of late Hercynian age in Morocco. Nature, 237, 160–162. M ATTE , P. 1986. La chaıˆne varisque parmi les chaıˆnes pale´ozoı¨ques pe´ri atlantiques, mode`le d’e´volution et position des grands blocs continentaux au PermoCarbonife`re. Bulletin de le Socie´te´ Ge´ologique de France, II, 8, 9 –24. M AZE´ AS , J. P. & P OUIT , G. 1968. Marques de mouvements hercyniens a` composante tangentielle de grande amplitude dans la boutonnie`re pre´cambrienne et infracambrienne du Bas oued Dra (Maroc me´ridional). Comptes Rendus de l’Acade´mie des Sciences, 267, 1549– 1552. M ICHARD , A. 1976. Ele´ments de Ge´ologie marocaine. Notes et Me´moires du Service Ge´ologique du Maroc, 252. M ICHARD , A., Y AZIDI , A., H OLLARD , H., B ENZIANE , F. & W ELLEFRT , S. 1982. Foreland thrusts and olistromes on the pre-Sahara margin of the Variscan orogen, Morocco. Geology, 10, 253–256. M ISSENARD , Y., Z EYEN , H., F RIZON DE L AMOTTE , D., L ETURMY , P., P ETIT , C., S E´ BRIER , M. & S ADDIQI , O. 2006. Crustal versus asthenospheric origin of relief of the Atlas Mountains of Morocco. Journal of Geophysical Research, 111, doi:10.1029/ 2005JB003708. O UANAIMI , H. & P ETIT , J. P. 1992. La limite sud de la chaıˆne hercynienne dans le Haut Atlas (Maroc): reconstitution d’un saillant non de´forme´. Bulletin de la Socie´te´ Ge´ologique de France, 163, 63–72. O UDRA , M., B ERAAOUZ , E. H., I KENNE , M., G ASQUET , D. & S OULAIMANI , A. 2006. La tectonique panafricaine du secteur d’Igherm: implication des doˆmes extensifs tardi- a` post-oroge´niques (Anti-Atlas occidental, Maroc). Estudios Geologicas, 61, 177–189. P IQUE´ , A. 1989. Variscan terranes in Morocco. Geological Society of America, Special Papers, 230, 115– 129. P IQUE´ , A. & M ICHARD , A. 1989. Moroccan Hercynides: A synopsis. The Palaeozoic sedimentary and tectonic evolution at the northern margin of West Africa. American Journal of Science, 289, 286–330. P IQUE´ , A., B OUABDELLI , M., S OULAIMANI , A., Y OUBI , N. & I LLIANI , M. 1999. Les conglome´rats du PIII (Prote´rozoı¨que terminal) de l’Anti-Atlas (Sud du Maroc): Molasses tardi-Panafricaines, ou marqueurs
451
d’un rifting fini-Prote´rozoı¨que. Comptes Rendus de l’Acade´mie des Sciences, 328, 409– 414. R OBERT -C HARRUE , C. 2006. Tectonique de l’Anti-Atlas oriental (Maroc). PhD thesis, Neuchaˆtel University. R OCCI , G., B RONNER , G. & D ESCHAMPS , M. 1991. Crystalline basement of the West Africa Craton. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berlin, 31–61. R ODGERS , J. 1987. Chains of basement uplifts within cratons marginal to orogenic belts. American Journal of Science, 287, 661–692. R ODGERS , J. 1995. Lines of basement uplifts within the external parts of orogenic belts. American Journal of Science, 295, 455– 487. S AHABI , M., A SLANIAN , D. & O LIVET , J. L. 2004. Un nouveau de´part pour l’histoire de l’Atlantique central. Comptes Rendus Ge´oscience, 336, 1041– 1052. S AQUAQUE , A., A DMOU , H., K ARSON , J. A., 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 AQUAQUE , A., B EHARREF , M., A BIA , H., N RINI , Z., R EUBER , I. & K ARSON , J. A. 1992. Evidence for a Pan-African volcanic arc and wrench fault tectonics in the Jbel Saghro, Anti-Atlas, Morocco. Geologische Rundschau, 81, 1 –13. S EBAI , A., F ERAUD , 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 OUALHINE , S., T EJEA D E L EON , J. & H OEPFFNER , CH . 2003. Les facie`s se´dimentaires carbonife`res de Tisdafine (Anti-Atlas oriental): remplissage deltaı¨que d’un bassin en ‘pull-apart’ sur la bordure me´ridionale de l’accident sud-Atlasique. Bulletin de l’Institut Sciences, Rabat, 25, 31– 41. S OUGY , J. 1969. Grandes lignes structurales de la chaıˆne des Mauritanides et de son avant-pays (socle pre´cambrien et sa couverture infracambrienne et pale´ozoı¨que), Afrique de l’Ouest. Bulletin de la Socie´te´ Ge´ologique de France, 7, 133–149. S OULAIMANI , A. 1998. Interactions socle/couverture dans l’Anti-Atlas occidental (Maroc): rifting finiProte´rozoı¨que et orogene`se hercynienne. PhD thesis, Marrakech University. S OULAIMANI , A. & P IQUE´ , A. 2004. The Tasrirt structure (Kerdous inlier, wenstern Anti-Atlas, Morocco): a Late Pan-African transtensive dome. Journal of African Earth Sciences, 39, 247– 255. S OULAIMANI , A., L E C ORRE , C. & F ARAZDAQ , R. 1997. De´formation hercynienne et relation socle/couverture dans le domaine du Bas Draˆa (Anti-Atlas occidental, Maroc). Journal of African Earth Sciences, 24, 271– 284. S OULAIMANI , A., B OUABDELLI , M. & P IQUE´ , A. 2003. L’extension continentale au Ne´o-Prote´rozoı¨que supe´rieur–Cambrien infe´rieur dans l’Anti-Atlas (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, 174, 88– 92. S OULAIMANI , A., E SSAI¨ FI , A., Y OUBI , N. & H AFID , A. 2004. Les marqueurs structuraux et magmatiques de
452
A. SOULAIMANI & M. BURKHARD
l’extension crustale au Prote´rozoı¨que terminal– Cambrien basal autour du Massif de Kerdous (Anti-Atlas occidental, Maroc). Comptes Rendus Ge´oscience, 336, 1433– 1441. S OULAIMANI , A., J AFFAL , M., M AACHA , L., K CHIKACH , A., N AJINE , A. & S AIDI , A. 2006. Mode´lisation magne´tique de la suture ophiolitique de Bou Azzer– El Graara (Anti-Atlas central, Maroc). Implications sur la reconstitution ge´odynamiques panafricaine. Comptes Rendus Ge´oscience, 338, 153– 160. S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic model for the Palaeozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Science Letters, 196, 17– 33. T EIXELL , T., A YARZA , P., Z EYEN , H., F ERNANDEZ , M. & A RBOLEYA , M. L. 2005. Effects of mantle upwelling in a compressional setting: the Atlas Mountains of Morocco. Terra Nova, 17, 456– 461. 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 ILLENEUVE , M. 2005. Paleozoic basins in West Africa and the Mauritanide thrust belt. Journal of African Earth Sciences, 43, 166– 195. 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. Eve´nements panafricains dans l’Adrar Souttouf (Sahara marocain). Comptes Rendus Ge´oscience, 338, 359– 367. W ALSH , G. J., A LEINIKOFF , J. N., B ENZIANE , F., Y AZIDI , A. & A RMSTRONG , T. R. 2002. U–Pb zircon geochronology of the Palaeoproterozoic Tagragra de Tata inlier and its Neoproterozoic cover, western Anti-Atlas, Morocco. Precambrian Research, 117, 1–20. W EIJERMARS , R. 1993. Estimation of paleostress orientation within deformation zones between two mobile plates. Geological Society of America Bulletin, 105, 1419– 1510. W ILLIAMS , G. D., P OWELL , C. M. & C OOPER , M. A. 1989. Geometry and kinematics of inversion tectonics. In: C OOPER , M. A. & W ILLIAMS , G. D. (eds) Inversion Tectonics. Geological Society, London, Special Publications, 44, 3– 15. Y OUBI , N. 1998. Le volcanisme ‘post-collisionnel’: un magmatisme intraplaque relie´ a` des panaches mantelliques. Etude volcanologique et ge´ochimique. Exemple d’application dans le Ne´oprote´rozoı¨que terminal de l’Anti-Atlas et le Permien du Maroc. PhD thesis, Marrakech University.
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) LAHSSEN BAIDDER1, YOUSSEF RADDI2, MOHAMED TAHIRI2 & ANDRE´ MICHARD3 1
Laboratoire de Ge´odynamique, Faculte´ des Sciences Aı¨n Chok, BP 5366 Maaˆrif, Casablanca, Morocco (e-mail:
[email protected],
[email protected])
2
Direction du De´veloppement Minier, Ministe`re de l’Energie et des Mines, BP 6208, Rabat Agdal, Morocco 3
Laboratoire de Ge´ologie, Ecole Normale Supe´rieure, 24 rue Lhomond, 75231 Paris cedex 05, France
Abstract: The Anti-Atlas belt belongs to the northern fringe of the West African craton, moderately deformed during the Variscan orogeny south of the Meseta Block. Field-based investigations into the stratigraphy and structure of the Palaeozoic cover have been performed in the eastern part of Anti-Atlas, with emphasis on the Devonian terranes. The Pan-African basement, which crops out in the Ougnat massif, was fragmented into a mosaic of tilted blocks during a sequence of extensional faulting events that occurred from Cambrian to (mostly) Late Devonian times. The Devonian normal fault pattern indicates a multi-directional extension, with a dominant northward direction. The Variscan compression resulted in the inversion of the palaeofaults as strike-slip– reverse faults, the kinematics of which points to a NE-trending regional direction of shortening, probably Permian in age. The occurrence of the Late Devonian palaeofault array accounts for the thick-skinned style of the (eastern) Anti-Atlas belt. The Devonian paleogeography of the Anti-Atlas can be correlated with that of the Meseta, but the lack of any Late Devonian compressional event in the Anti-Atlas shows that the two domains were not mechanically coupled at that time.
Three segments of the Caledonian– Variscan orogen superimposed onto the Pan-African crust are shown around the West African craton (WAC; Fig. 1a): (1) the Mauritanide belt, which is thrust over the western border of the continent and continues northward to the Zemmour region; (2) the Anti-Atlas and Ougarta belts, moulded around the northern and northeastern borders of the West African craton; (3) the Meseta Block, included in the foreland domain of the Alpine belt, north of the South Atlas Fault (SAF). The strongly deformed, partly metamorphic Meseta Block collided obliquely against the Anti-Atlas domain by the end of Carboniferous along a reverse, dextral strike-slip zone (Atlas Palaeozoic Transform Zone, APTZ), nearly coincident with the SAF (Michard et al. 1982; Pique´ & Michard 1989; Houari & Hoepffner 2003), so that the weakly deformed Anti-Atlas belt can be regarded as the foreland belt of the Mesetan Variscides (Simancas et al. 2005). However, Helg et al. (2004) and Burkhard et al. (2006) emphasized that the Anti-Atlas differs from a typical, thin-skinned foreland fold–thrust belt to the Appalachian– Variscan orogen, and that it is rather to be
considered as a severely inverted intracratonic basin. This raises the problem of the significance of the inverted palaeofaults of this dominantly thick-skinned belt. Another current issue is that of the preCarboniferous displacements of the Meseta Block after it was rifted away from Africa during Late Proterozoic–Cambrian times (Pique´ 2001; Soulaimani et al. 2003). Many workers have argued that the Meseta domain never drifted far from Africa (e.g. Pique´ 2001; El Hassani et al. 2003; Hoepffner et al. 2005, 2006), whereas others have assumed that large oceans occurred between these domains during Ordovician –Early Carboniferous time (e.g. Stampfli & Borel 2002; Burkhard et al. 2006). In the present paper, we focus on the fragmentation of the eastern Anti-Atlas Pan-African basement through pre-Carboniferous time, with emphasis on the Devonian period. The Palaeozoic sequences of the eastern Anti-Atlas have been well investigated from the stratigraphic and palaeontological points of view (Hollard 1960, 1967, 1974, 1981; Wendt et al. 1984; Destombes et al. 1985; Wendt & Aigner 1985; Wendt 1985, 1988; Wendt & Belka 1991).
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 453–465. DOI: 10.1144/SP297.21 0305-8719/08/$15.00 # The Geological Society of London 2008.
454
L. BAIDDER ET AL.
Fig. 1. Location and geological setting of the eastern Anti-Atlas. (a) Schematic structural map of Morocco. SAF, South Atlas Fault; AMA, ‘Accident majeur de l’Anti-Atlas’ (Main Anti-Atlas Fault). The Zemmour folds represent the northern tip of the Mauritanide belt. (b) Simplified geological map of the eastern Anti-Atlas, with location of Figures 2a,b and 3, after the geological maps at scale 1/200 000, sheet Todhra– Maider (Destombes & Hollard 1988) and Tafilalt–Taouz (Destombes & Hollard 1986).
Moreover, excellent maps at 1/200 000 scale are available for the area (Destombes & Hollard 1986, 1988). Based on a general survey of the Tafilalt– Maider areas (Fig. 1b) and on mapping at 1/50 000 scale south of the Ougnat massif, we
show that multidirectional extension occurred from Cambrian to Devonian time, being particularly active during the Late Devonian. The Variscan deformation is heavily controlled by the occurrence of this palaeofault array. The Late Devonian
Fig. 2. Pre-Devonian stratigraphy of the Palaeozoic cover of the eastern Anti-Atlas. (a) Oukhit region, SE of Ougnat. (b) Bouadil region, SW of Ougnat. The commonly used, informal lithostratigraphic units are labelled on the right-hand side of (a), with ages after Destombes et al. (1985), Destombes & Hollard (1986, 1988).
DEVONIAN EXTENSION IN ANTI-ATLAS
455
Fig. 3. Devonian facies pattern of the eastern Anti-Atlas during the late Early–Late Devonian interval after Hollard (1974), Wendt (1985) and personal observations, with location of the stratigraphic columns and photographs of Figures 4 –6 (italic numbers). (For location of the map see Fig. 1.)
extensional setting of the eastern Anti-Atlas compares with that of the western Anti-Atlas, and contrasts with the compression –transpression regime then prevailing in the Meseta Block.
Geological setting The Palaeozoic sequences of the eastern Anti-Atlas overlie a Precambrian (Pan-African) basement that crops out widely in the Jebel (J.) Saghro and Ougnat massifs (‘boutonnie`res’), and in some tiny outcrops east of Erfoud (Fig. 1b) and south of J. Ougnat (J. Angal; see Fig. 10). The older, Cambrian –Ordovician Palaeozoic terranes extend around the Precambrian massifs and in the SE –NW-trending Ougnat-Ouzina Ridge (OOR), whereas the younger ones (Devonian –Early Carboniferous) extend in to the Maider and Tafilalt areas, bounded to the south by the Oum JeraneTaouz (OJT) fault; that is, the east –west branch of the Main Anti-Atlas Fault (AMA; Choubert 1947). North of the Saghro–Ougnat axis, the Early Carboniferous Tineghir– Tisdafine units belong to the South Meseta Zone (Michard et al. 1982; Soualhine et al. 2003; Hoepffner et al. 2006) thrust southward onto the Anti-Atlas domain. The folded Palaeozoic sediments of the eastern Anti-Atlas are intruded by the 200 Ma sills and dykes of the Central Atlantic Magmatic Province (Hollard 1973; Knight et al. 2004), and
overlain by horizontal Cretaceous – Neogene sediments (hamadas) to the south and east. To the north, the transgressive Mesozoic –Cenozoic deposits are tilted northward and eventually folded close to the Atlas belt as a result of the Oligocene– Neogene compression. Elsewhere, only discrete flexures and faults in the Cenomanian –Turonian limestones and Pliocene –Quaternary deposits convey the weak compressional reactivation of some of the basement faults. The exhumation of the Saghro–Ougnat Precambrian axis to the surface results from the superimposed effects of its Late Triassic –Jurassic uplift as the southern shoulder of the Atlas rift and of the Atlas orogeny (Se´brier et al. 2006), combined with a large-scale mantle anomaly (Missenard et al. 2006).
Palaeozoic stratigraphy and paleogeography The Palaeozoic sequence (Fig. 2) begins with coarse clastic deposits (Lie-de-vin conglomerates, Gre`s Terminaux sandstones) dated from the late Early Cambrian, as the older, ‘Adoudounian’– Lower Cambrian carbonates of western Anti-Atlas are lacking here (Destombes & Hollard 1986). The Gre`s Terminaux thickness abruptly changes from c. 400 m SW of the Ougnat massif to c. 150 m going eastward across the Oued Smile
456
L. BAIDDER ET AL.
Fig. 4. Contrasted stratigraphic columns of the Devonian formations of the Tafilalt platform (a) and Maider basin (b).
fault (see Fig. 7), and the whole sequence vanishes in northeasternmost Ougnat. The Early Cambrian sandstones are followed by finer-grained clastic deposits of Middle Cambrian age associated with
alkaline mafic volcanic rocks, remarkably developed SE of the Ougnat massif, between the N120-trending Smile and South Oukhit faults (Destombes 2006).
DEVONIAN EXTENSION IN ANTI-ATLAS
457
Fig. 5. North– south variations in the Late Devonian stratigraphy within the Tafilalt platform, modified from Wendt (1988). Same lithological signatures as in Figure 4. The abrupt stratigraphic changes are related to the east–west-trending faults shown in Figures 7 and 9. A, J. Amelane; IF, Bou Ifarou; AT, J. Atrous; JD, Jdaid.
After a possible gap corresponding to the Late Cambrian (Destombes et al. 1985; Destombes & Feist 1987), clastic sedimentation resumed and lasted until the end of Ordovician. The Tremadoc onlap is restricted to the southern part of the study area, and the isopachs for Early Ordovician sediments parallel the east –west trend of the Saghro–Ougnat axis, which was partly emerged at that time (Destombes et al. 1985). In contrast, the isopachs of the later Ordovician formations (Caradoc –Ashgill) parallel the SE –NW-trending Ougnat– Ouzina Ridge (OOR). The chaotic conglomerates of Caradoc age, probably associated with palaeofault scarps, also crop out along the OOR. After the Late Ordovician Saharan glaciation, Silurian graptolite-bearing black shales accumulated over the entire region as a response to the post-glacial eustatic transgression. However, the Early Silurian (Llandovery) pelites are preserved only within a faulted ‘channel’ extending along the OOR (Destombes et al. 1985). Carbonate sedimentation progressively increased during the Early and Middle Devonian, as did the differentiation of the facies pattern (Figs 3 & 4), as shown by Hollard (1967, 1974, 1981), and Kaufmann (1998a, b) and Destombes & Hollard (1986). The Lochkovian–Emsian sedimentary pile is a rather homogeneous sequence of mixed, well-bedded fine-grained limestones, marls and rare siliciclastic deposits (Fig. 4a and b). Calcalkaline basaltic flows and peperites were emplaced north of Hamar Laghdad (Fig. 3) in the Lochkovian sediments. Some coral biostromes
and most spectacular Stromatactis mud mounds (Brachert et al. 1992; Montenat et al. 1996; Mounji et al. 1998; Kaufmann 1998a, b) developed on the rims of the shallow Tafilalt –Maider shelf during the Emsian–Givetian. A deep-water environment is suggested in the south Tafilalt basin by the occurrence of dark mudstones and nodular limestones. Subsidence was still stronger in the Maider basin, with slump folds facing the centre of the basin (Fig. 4b). The Late Devonian of the Eastern Anti-Atlas, studied in detail by Hollard (1967, 1974, 1981), Wendt et al. (1984), Bensaid et al. (1985), Wendt (1985, 1988), Wendt & Aigner (1985), Becker (1990), Walliser et al. (1995) and Wendt & Belka (1991) corresponds to even more differentiated sequences (Figs 3 and 5). The previous Tafilalt – Maider shelf changed into a restricted Tafilalt shelf bounded to the north, SE and west by deep basins (Oued Rheris – Jorf, South Tafilalt and Maider basins, respectively). Part of the Middle Devonian shelf became emergent on the Ougnat – Ouzina Ridge (e.g. J. Gherghiz; sometimes wrongly spelled ‘J. Rheris’) during the Early Frasnian, so that Upper Frasnian cephalopod limestones or Famennian phosphatic conglomerates (Fro¨hlich 2004) and sandy limestones unconformably overlie the eroded Lower – Middle Devonian limestones (Fig. 6a), and locally the Ordovician layers (Fig. 5d; Destombes & Hollard 1986). The basins were filled with clays and bioclastic wackestones during the
458
L. BAIDDER ET AL.
Fig. 6. Sedimentary records of the Devonian normal fault activity (see Fig. 3 for location). (a) Famennian ferruginous conglomerates (right) unconformably overlying Pragian limestones (left), Wadi Smile north of J. Gherghiz (pencil and notebook indicate scale). (b) Famennian seismites close to the N-Mecissi fault, J. Gherghiz massif. (c) Extensional tilted blocks in Middle Devonian limestones, sealed by Upper Frasnian–Famennian layers, Jebel Mrakib north of the Oum Jerane–Taouz main normal fault. (d) Boudinage and tilting of the Givetian limestones by NE– SW-trending minor palaeofaults, and associated debris flows, J. Issimour, north Maider basin. (e) Minor NW–SE-trending palaeofaults sealed by Famennian layers, J. Bou Tcharafine. (f) Upper part of a minor east– west-trending palaeofault sealed by Famennian layers, J. Amelane. (Note the disrupted Frasnian limestones on the right of the picture.)
Frasnian, whereas on the northern Tafilalt platform a few metres of cephalopod limestones were deposited (Figs 4a and 5a), replaced by quartzose brachiopod layers on the southern platform (Fig. 5c). Differential subsidence reached its maximum during the Famennian, when the Maider basin was infilled by several hundred metres of clays with calcareous mudstones and sandstones intercalations (Fig. 4b), contrasting
with the veneer of condensed cephalopod limestones deposited upon the Tafilalt platform (Figs 4a and 5a). The Saghro – Ougnat axis was again emerged at that time and contributed to the infilling of the Late Devonian basins. The transition from the latest Famennian clays and limestones to the Early Carboniferous deltaic sequence is generally transitional (Destombes & Hollard 1986), but the Tournaisian sandy pelites
DEVONIAN EXTENSION IN ANTI-ATLAS
unconformably transgress onto the Lower Devonian limestones at the southern border of the Ougnat massif. The Upper Vise´an chaotic breccias and olistostromes that occur in the Tineghir– Tisdafine area are not considered here as they belong to the South Meseta zone, thrust southward onto the Anti-Atlas foreland (Michard et al. 1982; Hoepffner et al. 2006).
Devonian fault pattern The differentiated sedimentary facies and contrasted palaeogeography of the Devonian period record the activity of palaeofaults that ‘disintegrated’ the eastern Anti-Atlas crust at that time, particularly during the Late Devonian (Wendt 1985; Wendt & Belka 1991). However, the whole palaeofault array and the geometry of the related crustal blocks have not been considered yet. We present in this section the most significant examples of Devonian palaeofaults that can be recognized in the Tafilalt –Maider area, and the broad organization of the resulting fault array. The faults are qualitatively ordered into first, second and third order according to their mean throw and regional extension.
First-order ENE – WSW-trending faults South of the Ougnat massif, and north of the Mecissi (Msissi) village, the J. Gherghiz–Tamjout Devonian massifs normally overlie the Lower Palaeozoic series to the north, whereas they are bounded to the SSE by a N70-trending major fault (N-Mecissi fault), which connects eastward with the N120-trending Signit fault (Fig. 7). The N-Mecissi fault was reactivated as a reverse, dextral strike-slip fault during the Variscan compression (Raddi et al. 2007a, c), but its Late Devonian (Frasnian) activity as a NW-dipping normal fault can be inferred from the northward wedging of the pre-Famennian layers, unconformably overlain by the Famennian phosphatic conglomerates (Figs 6a and 8). Consistently, seismites (Montenat et al. 2007) are observed in the Famennian formation close to the palaeofault (Fig. 6b). During the Famennian, the N-Mecissi fault separated two southward-tilted blocks: the Mecissi block, which remains undeformed to the present day, and the Angal–Gherghiz block, now crushed between the adjacent blocks. The Oum Jerane –Taouz (OJT) fault belongs to the same group of ENE –WSW-trending major faults as the N-Mecissi fault. Its early Late Devonian normal throw is recorded in the J. Mrakib, immediately north of the OJT fault, by a system of extensional minor faults that affect the Middle Devonian limestones and are sealed by the Upper Frasnian– Famennian layers (Fig. 6c). However, a northwarddipping slope had already formed during the
459
Eifelian, as shown by the development of mud mounds north of the platform border marked by the J. Mrakib reef mound (Kaufmann 1998a, b).
Second-order NW – SE- and NNE – SSW-trending faults The occurrence of NW–SE- and NNE –SSW-trending Late Devonian palaeofaults on both sides of the Maider basin (Fig. 7) is strongly suggested by the dramatic contrast between the stratigraphy of the Frasnian –Famennian formations of the subsiding basin and that of the adjacent platforms (Fig. 5), as well as by the slump folds converging towards the basin axis in the Famennian layers (Wendt 1985). It is worth noting that the palaeofault activity began as early as the Middle Devonian, as dismembered limestone blocks and debris flows of Middle Devonian age can be also observed (Fig. 6d). West of the Hamar Laghdad mud mounds (Hollard 1967; Brachert et al. 1992; Montenat et al. 1996; Mounji et al. 1998), a submeridian, east-dipping normal fault was active during the Emsian, and partly controlled the mud mound formation and preservation in the downthrown block (Montenat et al. 1996).
Third-order ENE – WSW- to ESE – WNW-trending faults These relatively minor palaeofaults are frequently preserved beneath an unconformable sealing succession of Famennian layers, in contrast to the major faults, which were reactivated during the Variscan compression. The Erfoud and Bou Tcharafine faults are clear examples of such sealed palaeofaults (Figs 6d,e and 7). Synsedimentary breccias are associated with another of these roughly east –west faults at J. Amelane (Fig. 6f). Similar faults are observed at J. Kreir, close to the Jorf–Rheris basin (Fig. 7), where a NE–SWtrending hectometre-scale graben affecting the Middle Devonian limestones is sealed by the Upper Frasnian layers.
Middle– Late Devonian multi-directional extension The broad organization of the reported palaeofaults (Fig. 7) allows us to infer a multi-directional extension of the eastern Anti-Atlas area during the Middle –Late Devonian. The first- and third-order ENE –WSW- to ESE –WNW-trending faults point to a dominant NNW–SSE to NNE–SSW direction of extension north of the OJT branch of the AMA (Fig. 9). Contemporaneously, another direction of extension, oriented ENE –WSW, corresponds to the second-order, NW–SE-trending faults. This
460
L. BAIDDER ET AL.
Fig. 7. Sketch map of the Devonian fault pattern. Straight lines with teeth, normal palaeofaults with clear Devonian kinematics; undecorated straight lines, undifferentiated faults. Stereograms are lower hemisphere; arrows indicate striae directions. ABF, Aberchane fault; AF, Aguelmous n’Fezzou syncline; AKF, Akerouz fault; AMA, Main Anti-Atlas fault; EF, Erfoud fault; ESF, East Signit fault; FF, Fezzou fault; JA, Jebel Atrous syncline; JT, Jebel Tijekht anticline; NMF, N-Mecissi fault; OJTF, Oum Jerane– Taouz fault; SF, Wadi Smile fault; TF, Tisserdmine fault. Location of Figure 10 is shown.
normal fault pattern results in a mosaic of crustal blocks of various scales tilted either southward or eastward.
Discussion Lower Palaeozoic and Devonian palaeofaults The ENE –WSW trend of the first-order and part of the third-order Devonian faults (Fig. 9) corresponds
to that of the Anti-Atlas axis, uplifted as the shoulder of the Cambrian High Atlas–Meseta rift (Ouanaimi & Petit 1992; Ouali et al. 2000). We note that the thickness of the Early Cambrian clastic deposits abruptly changes crossing the NW– SW-trending Smile fault, suggesting synsedimentary activity of this fault. The same direction, which is that of the second-order Devonian faults as well as that of the Ougnat–Ouzina Ridge, controls the palaeogeography during the Middle Cambrian (alkaline volcanism extension between
DEVONIAN EXTENSION IN ANTI-ATLAS
461
Fig. 8. An example of first-order, NNW-dipping normal fault of Late Devonian age: the N-Mecissi fault (see Fig. 7 for location). (a) Diagrammatic representation. (b) Schematic stratigraphic column from J. Gherghis, immediately north of the fault. (c) Schematic stratigraphic column, about 3 km further north in the Wadi Smile tilted block. (See also Fig. 6a and b for details.)
the Oued Smile and South Oukhit faults), and the Caradoc, with the Imzioui fault scarp conglomerates (Destombes 2006). Therefore the observed Devonian palaeofault pattern is essentially inherited from Early Palaeozoic faults, which in turn can reactivate late Neoproterozoic structures as in the western Anti-Atlas (Jeannette & Pique´ 1981; Soulaimani et al. 1997, 2003). The metacratonic crust dislocation appears as a long-lasting process that culminated during the Late Devonian. In the same way, the importance of the Early–Middle Devonian extension and normal faulting has been described by Ouanaimi & Lazreq (2008) in the western –central Anti-Atlas, although it is less easily deciphered than in the eastern Anti-Atlas because of the stronger Variscan folding.
Implications for the Variscan deformation The Anti-Atlas and Ougarta Variscan deformation results from the collision of the WAC indenter
(Reguibat Arch) with the Appalachian – Mauritanide belt, the Meseta Block and the Hoggar craton from west to north, to east, respectively (Haddoum et al. 2001; Hoepffner et al. 2005; Simancas et al. 2005). The mosaic of tilted blocks achieved during the Late Devonian extension reacted to the compressional stress by palaeofault inversion, strike-slip displacements, and crushing of some of the blocks. This can be illustrated through the map of the top of basement depth around the Ougnat culmination (Fig. 10). The top of basement isobaths are deduced from the dip and thickness measurements in the Palaeozoic cover of each block (Raddi et al. 2007c). Most of the blocks are monoclinal (e.g. Central Ougnat, Mecissi or Tarhouilast block), whereas others are crushed and sheared, such as the Angal –Gherghiz triangle or Ighil n’Ighiz block. Structural analysis of the folds and fold– fault relationships (en e´chelon and/or overturned minor folds) allows us to decipher the fault kinematics
Fig. 9. Diagrammatic representation of the Late Devonian palaeofault pattern in the Tafilalt– Maider regions.
462
L. BAIDDER ET AL.
during the Variscan compression (Raddi et al. 2007c). Most of the palaeofaults are partly or totally inverted and changed into strike-slip faults, either right-lateral when they trend N120 to N70, or left-lateral when their orientation is close to N10–N30. We emphasize that the horizontal displacements are probably of the same order of magnitude as the vertical displacements on the inverted faults. The kinematic pattern (Fig. 10) is consistent with a regional, NE–SW maximum stress direction, which is comparable with the Stephanian– Autunian shortening direction of the Ougarta chain (Donzeau 1974; Haddoum et al. 2001; Fabre 2005) and Meseta domain (Saber 1994; Saidi et al. 2002). However, a south-directed compression is recorded along the northern border of the Saghro –Ougnat axis (Tineghir –Tisdafine imbrications; Michard et al. 1982; Hoepffner et al. 2006) and Bechar
basin further east (Djebel Bechar; Fabre 2005, p. 358). This event is comparable with the southverging, last shortening event of the western AntiAtlas (Caritg et al. 2004; Hoepffner et al. 2006), and also with the late Westphalian event recognized in the Meseta Block (Hoepffner et al. 2006). The curved fold axis of the south Tafilalt area and the cross-folds of the Maider basin (Fig. 7) record the interference of these superimposed shortening phases in the detached Palaeozoic cover. A similar interference pattern can be inferred at a regional scale from the distribution of the Precambrian and Cambrian– Ordovician axes compared with Late Palaeozoic basins (Robert-Charrue 2006). In contrast, the earlier (late Namurian– Westphalian), ESE–WNW-directed main compression event recognized in the western Anti-Atlas (Soulaimani et al. 1997; Belfoul et al. 2002; Caritg et al. 2004; Fabre 2005; Burkhard et al. 2006) and
Fig. 10. The deformed mosaic of tilted basement blocks of the eastern Anti-Atlas around the Ougnat culmination. The top of basement isobaths have been mapped based on the stratigraphic thicknesses and mean dip of the Palaeozoic cover. Kinematic information (arrows) is inferred from structural observations in the folded cover (mainly en e´chelon and reclining folds). (See Raddi et al. (2007a –c) for details.)
DEVONIAN EXTENSION IN ANTI-ATLAS
western Meseta (Hoepffner et al. 2006) is lacking in the eastern Anti-Atlas. It must be noted that our analysis is only partly consistent with the recent proposals by Robert-Charrue (2006), who made a point of the occurrence of a Triassic extensional reactivation of the north-dipping Cambrian palaeofaults inverted during the Variscan compression. In contrast, we emphasize that the Palaeozoic palaeofaults display three different orientations and varied dips (Fig. 7), and are strongly dependent on the Devonian crustal dislocation event (Figs 5, 6 and 8). We also insist on the importance of strike-slip movements in the faulted basement besides the vertical movements (reverse faults, pop-up structures). It is hoped that these different approaches will converge in the near future.
Mesetan correlations The contrasted Middle–Late Devonian palaeogeography observed in the eastern and western AntiAtlas is comparable with that of the western Meseta, which is characterized by a western carbonate shelf and an eastern, fault-bounded turbiditic basin (Pique´ 1987; Bouabdelli & Pique´ 1996). Moreover, the faunal similarities are important between the platform deposits of western Meseta and Anti-Atlas (Bohemian facies; Hollard 1974, 1981). This is consistent with the Meseta Block lying close to the Anti-Atlas during the Devonian, as argued by Pique´ & Michard (1989), El Hassani et al. (2003) and Hoepffner et al. (2006), and contrary to alternative proposals by Stampfli & Borel (2002) and Burkhard et al. (2006). However, the palaeotectonic, Late Devonian setting of the Meseta Block differs from that of the Anti-Atlas in the occurrence of varied records of tectonic shortening (i.e. from west to east): (1) a possible compression event in the Doukkala basin (Echarfaoui et al. 2002); (2) a compressional –transtensional regime during the opening of the western Meseta basins (Pique´ 1987; Bouabdelli & Pique´ 1996); (3) more significantly, the onset of the Eovariscan synmetamorphic folding phases in the eastern Meseta (Pique´ & Michard 1989; Hoepffner et al. 2006). This suggests that the Meseta Block was not yet coupled to the Anti-Atlas at that time.
Conclusion Middle–Late Devonian multi-directional extension, dominantly oriented NNW–SSE and ENE– WNW, resulted in the formation of a mosaic of crustal blocks, tilted either southward or eastward, at the expense of the Pan-African crust of the
463
eastern Anti-Atlas. This major crustal dislocation succeeded earlier extensional events, associated with alkaline mafic volcanism during the Middle Cambrian. The fragmented crust deformed easily in response to the Variscan compression, through horizontal displacements of the blocks combined with inversion of the palaeofaults. Consequently, at the southern front of the Variscan Meseta Block, the eastern Anti-Atlas foreland belt displays a ‘thick-skinned’ tectonic style. The eastern Anti-Atlas is comparable with the western parts of the chain as far as the major role of the Devonian extension in crustal dislocation is concerned. In contrast, the eastern Anti-Atlas differs from the western Anti-Atlas in the timing and intensity of the Variscan deformation. The western, relatively intense folding events can be ascribed to Namurian–Westphalian and Late Westphalian –Stephanian phases, whereas the weaker deformation of the eastern Anti-Atlas results from probably younger, Late Westphalian– Stephanian and Early Permian events. The last event corresponds to the NE–SW-trending compression that gave birth to the Ougarta chain. The Middle –Late Devonian dislocation of the whole Anti-Atlas domain is comparable with that of the western Meseta, as are the associated sedimentary facies and faunal associations. At that time, the Anti-Atlas would represent a proximal passive margin domain, whereas the Meseta formed a distal system of blocks somewhat separated from the Gondwana mainland. A.M. acknowledges grants from the Bureau de Recherches Ge´ologique et Minie`res, programme Ge´oforma, and from the Ministe`re de l’Energie et des Mines, Rabat, Programme National de Cartographie ge´ologique. The authors are indebted to J. P. Prian (BRGM), L. Tabit, A. El Khlifi and A. Charik (MEM) for technical support at Rabat, and to P. Morzadec, H. Ouanaimi and an anonymous reviewer for their constructive reviews.
References B ECKER , R. T. 1990. Stratigraphische Gliederung und Ammonoideen-Faunen im Nehdenium (Oberdevon II) von Europa und Nord-Afrika. Dissertation, Ruhr-Universita¨t Bochum. B ELFOUL , M. A., F AIK , F. & H ASSENFORDER , B. 2002. Evidence of a tangential tectonic event prior to the major folding in the Variscan belt of western antiAtlas. Journal of African Earth Sciences, 32, 723– 739. B ENSAID , M., B ULTYNCK , P., S ARTENAER , P., W ALLISER , O. H. & Z IEGLER , W. 1985. The Givetian–Frasnian boundary in Pre-Sahara Morocco. Courier Forschungsinstitut Senckenberg, 75, 287– 300. B OUABDELLI , M. & P IQUE´ , A. 1996. Du bassin sur de´crochement au bassin d’avant-pays: dynamique du
464
L. BAIDDER ET AL.
bassin d’Azrou– Khenifra (Maroc hercynien central). Journal of African Earth Sciences, 22, 213– 224. B RACHERT , T. C., B UGGISCH , W., F LU¨ GEL , E., H U¨ SSNER , H. M., J OACHIMSKI , M. M., T OURNEUR , E. & W ALLISER , O. H. 1992. Controls of mud mounds formation: the Early Devonian Kess-Kess carbonates of the Hamar Laghdad, Anti-Atlas, Morocco. Geologische Rundschau, 81, 15–44. M., C ARITG , S., H ELG , U., B URKHARD , R OBERT -C HARRUE , Ch. & S OULAIMANI , A. 2006. Tectonics of the Anti-Atlas of Morocco. Comptes Rendus Ge´oscience, 338, 11–24. 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. C HOUBERT , G. 1947. L’accident majeur de l’Anti-Atlas. Comptes Rendus de l’Acade´mie des Sciences, 224, 1172–1173. D ESTOMBES , J. 2006. Carte Ge´ologique au 1:200 000 de l’Anti-Atlas marocain. Pale´ozoı¨que infe´rieur: Cambrien moyen et supe´rieur, Ordovicien, base du Silurien. Feuille Tafilalt Taouz. Chap. E. Notes et Me´moires du Service Ge´ologique du Maroc, 244bis. D ESTOMBES , J. & F EIST , R. 1987. De´couverte du Cambrien supe´rieur en Afrique (Anti-Atlas central, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 304, 719–724. D ESTOMBES , J. & H OLLARD , H. 1986. Carte Ge´ologique du Maroc au 1:200 000. Feuille Tafilalt– Taouz. Notes et Me´moires du Service Ge´ologique du Maroc, 244. D ESTOMBES , J. & H OLLARD , H. 1988. Carte Ge´ologique du Maroc au 1:200 000. Feuille Todgha–Ma’der. Notes et Me´moires du Service Ge´ologique du Maroc, 243. D ESTOMBES , J., H OLLARD , H. & W ILLEFERT , S. 1985. Lower Palaeozoic rocks of Morocco. In: H OLLAND , C. H. (ed.) Lower Palaeozoic rocks of North-Western and West–Central Africa. Wiley, Chichester, 91–336. D ONZEAU , M. 1974. L’Arc Anti-Atlas– Ougarta (Sahara nord-occidental, Alge´rie –Maroc). Comptes Rendus de l’Acade´mie des Sciences, 278, 417–420. E CHARFAOUI , H., H AFID , M. & A IT S ALEM , A. 2002. Structure sismique du socle pale´ozoı¨que du bassin des Doukkala, Moˆle coˆtier, Maroc occidental. Indication en faveur de l’existence d’une phase e´ovarisque. Comptes Rendus Ge´oscience, 334, 13– 20. E L H ASSANI , A., T AHIRI , A. & W ALLISER , O. H. 2003. The Variscan crust between Gondwana and Baltica. Courier Forschungsinstitut Senckenberg, 242, 81– 87. F ABRE , J. 2005. Ge´ologie du Sahara occidental et central. Tervuren African Geoscience Collection, 108. F RO¨ HLICH , S. 2004. Phosphatic black pebbles and nodules on a Devonian carbonate shelf (Anti-Atlas, Morocco). Journal of African Earth Sciences, 38, 243– 254. H ADDOUM , H., G UIRAUD , R. & M OUSSINE P OUCHKINE , A. 2001. Hercynian compressional deformations of the Ahnet–Mouydir Basin, Algerian Saharan Platform: far-field stress effects of the Late Palaeozoic orogeny. Terra Nova, 13, 220– 226. H ELG , U., B URKHARD , M., C ARITG , S. & R OBERT -C HARRUE , C. H. 2004. Folding and
inversion tectonics in the western Anti-Atlas of Morocco. Tectonics, 23, paper number TC4006. H OEPFFNER , C., H OUARI , M. R. & B OUABDELLI , M. 2006. Tectonics of the North African Variscides (Morocco, Western Algeria), an outline. Comptes Rendus Ge´oscience, 338, 25– 40. H OEPFFNER , H., S OULAIMANI , A. & P IQUE´ , A. 2005. The Moroccan Hercynides. Journal of African Earth Sciences, 43, 144–165. H OLLARD , H. 1960. Une phase tectonique intrafamennienne dans le Tafilalet et le Maider (Maroc pre´saharien). Comptes Rendus de l’Acade´mie des Sciences, 250, 1303–1305. H OLLARD , H. 1967. Le De´vonien du Maroc et du Sahara nord occidental. International Symposium on the Devonian System. Alberta Society of Petroleum Geology, Calgary, I, 203– 244. H OLLARD , H. 1973. La mise en place au Lias des dolerites dans le Pale´ozoı¨que moyen des plaines du Draˆa et du bassin de Tindouf (Sud de l’Anti-Atlas central, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 277, 553– 556. H OLLARD , H. 1974. Recherche sur la stratigraphie des formations du De´vonien moyen, de l’Emsien supe´rieur au Frasnien, dans le Sud du Tafilalt et dans le Ma’der (Anti-Atlas oriental, Maroc). Notes et Me´moires du Service Ge´ologique du Maroc, 264, 7–68. H OLLARD , H. 1981. Principaux caracte`res des formations de´voniennes de l’Anti-Atlas. Notes et Me´moires du Service Ge´ologique du Maroc, 308, 15– 21. H OUARI , M. R. & H OEPFFNER , C. 2003. Late Carboniferous dextral wrench-dominated transpression along the North African craton margin (Eastern High-Atlas, Morocco). Journal of African Earth Sciences, 37, 11–24. J EANNETTE , D. & P IQUE´ , A. 1981. Le Maroc hercynien: plate forme disloque´e du craton Ouest-Africain. Comptes Rendus de l’Acade´mie des Sciences, 293, 79–82. K AUFMANN , B. 1998a. Middle Devonian reef and mudmounds on a carbonate ramp: Ma’der Basin (Eastern Anti-Atlas, Morocco). In: W RIGHT , V. P. & B URCHETT , T. P. (eds) Carbonate Ramps. Geological Society, London, Special Publications, 149, 417– 435. K AUFMANN , B. 1998b. Facies, stratigraphy and diagenesis of Middle Devonian reef and mud-mounds in the Ma’der (Eastern Anti-Atlas, Morocco). Acta Geologica Polonica, 48, 43–106. K NIGHT , K. B., N OMADE , S., R ENNE , P. R., M ARZOLI , A., B ERTRAND , H. & Y OUBI , N. 2004. The Central Atlantic Magmatic Province at the Triassic– Jurassic boundary: paleomagnetic and 40Ar/39Ar evidence from Morocco for brief, episodic volcanism. Earth and Planetary Science Letters, 228, 143 –160. M ICHARD , A., Y AZIDI , A., B ENZIANE , F., H OLLARD , H. & W ILLEFERT , S. 1982. Foreland thrust and olistostromes on the pre-Saharan margin of the Variscan orogen, Morocco. Geology, 10, 253–256. M ISSENARD , Y., F RIZON DE L AMOTTE , D., Z EYEN , H., L ETOURMY , P., P ETIT , C. & S E´ BRIER , M. 2006. Crustal versus asthenospheric origin of the relief of the Atlas mountains of Morocco. Journal of Geophysical Research, 111, paper number B03401.
DEVONIAN EXTENSION IN ANTI-ATLAS M ONTENAT , C., B AIDDER , L., B ARRIER , P., H ILALI , A., L ACHKHEM , H. & M ENNING , J. 1996. Controˆle tectonique de l’e´dification des monticules biose´dimentaires de´voniens du Hmar Lakhdad d’Erfoud (Anti-Atlas oriental, Maroc). Comptes Rendus de l’Acade´mie des Sciences, 323, 297– 304. M ONTENAT , C., B ARRIER , P., O TT D ’ ESTEVOU , P. & H IBSCH , C. 2007. Seismites: An attempt at critical analysis and classification. Sedimentary Geology, 196, 5–30. M OUNJI , D., B OURQUE , P. A. & S AVARD , M. M. 1998. Hydrothermal origin of Devonian mud-mouds of Hmar–Lakhdad. Evidence from architectural and geochemical constraints. Geology, 26, 1123– 1126. O UALI , H., B RIAND , B., B OUCHARDON , J. L. & E L M AATAOUI , M. 2000. Mise en e´vidence d’un volcanisme alcalin intraplaque d’aˆge Acadien dans la Meseta nord-occidentale (Maroc). Comptes Rendus de l’Acade´mie des Sciences, 330, 611– 616. O UANAIMI , H. & L AZREQ , N. 2008. The ‘Rich’ group of the Draˆa Basin (Lower Devonian, Anti-Atlas, Morocco): an integrated sedimentary and tectonic approach. In: E NNIH , N. & L IE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 467–482. O UANAIMI , H. & P ETIT , J. P. 1992. La limite sud de la chaıˆne hercynienne dans le Haut-Atlas marocain: reconstitution d’un saillant non de´forme´. Bulletin de la Socie´te´ Ge´ologique de France, 163, 63–72. P IQUE´ , A. 1987. Un e´le´ment majeur de la Meseta marocaine nord occidentale: le bassin de´vono-dinantien de Sidi Bettache. Notes et Me´moires du Service Ge´ologique du Maroc, 323, 41– 64. P IQUE´ , A. 2001. Geology of Northwest Africa. Borntraeger, Berlin. P IQUE´ , A. & M ICHARD , A. 1989. Moroccan Hercynides, a synopsis. The Palaeozoic sedimentary and tectonic evolution at the northern margin of west Africa. American Journal of Science, 289, 286–330. R ADDI , Y., T AHIRI , M., D ERRE´ , C., L ECOLLE , M. & B AIDDER , L. 2007a. Carte ge´ologique du Maroc a` 1:50 000, feuille Oukhit. Notes et Me´moires du Service Ge´ologique du Maroc, 475 (in press). R ADDI , Y., T AHIRI , M., D ERRE´ , C., L ECOLLE , M. & B AIDDER , L. 2007b. Notice explicative de la carte ge´ologique du Maroc a` 1:50 000, feuille Oukhit. Notes et Me´moires du Service Ge´ologique du Maroc, 475bis (in press). R ADDI , Y., B AIDDER , L., T AHIRI , M. & M ICHARD , A. 2007c. Variscan deformation at the northern border of the West African Craton, eastern Anti-Atlas, Morocco: Compression of a mosaic of tilted blocks. Bulletin de la Socie´te´ Ge´ologique de France, 178, 343–352. R OBERT -C HARRUE , C. 2006. Ge´ologie structurale de l’Anti-Atlas oriental, Maroc. PhD thesis, Universite´ Neuchaˆtel. S ABER , H. 1994. Se´dimentologie et e´vidence d’une tectonique tardi-hercynienne d’aˆge Permien infe´rieur dans le bassin des Ida Ou Ziki, sud-ouest du massif ancien du Haut Atlas (re´gion d’Argana, Maroc). Journal of African Earth Sciences, 19, 99– 108.
465
S AIDI , A., T AHIRI , A., A IT B RAHIM , L. & S AIDI , M. 2002. Etats de contraintes et me´canismes d’ouverture et de fermeture des bassins permiens du Maroc hercynien. L’exemple des bassins des Jebilet et des Rehamna. Comptes Rendus Ge´oscience, 334, 221– 226. S E´ BRIER , M., S IAME , L., Z OUINE , E. M., W INTER , T., M ISSENARD , Y. & L ETURMY , P. 2006. Active tectonics in the Moroccan High Atlas. Comptes Rendus Ge´oscience, 338, 65– 79. S IMANCAS , J. F., T AHIRI , A., A ZOR , A., L ODEIRO , F. G., M ARTINEZ P OYATOS , D. J. & E L H ADI , H. 2005. The tectonic frame of the Variscan–Alleghanian orogen in Southern Europe and Northern Africa. Tectonophysics, 398, 181–198. S OUALHINE , S., T EJERA DE L EON , J. & H OEPFFNER , C. 2003. Les facie`s se´dimentaires carbonife`res de Tisdafine (Anti-Atlas Oriental): remplissage deltaı¨que d’un bassin en ‘pull-apart’ sur la bordure me´ridionale de l’accident sud-atlasique. Bulletin de l’Institut Scientifique de Rabat, 25, 31– 41. S OULAIMANI , A., L E C ORRE , C. & F ARAZDAQ , R. 1997. De´formation hercynienne et relation socle/couverture dans le domaine du Bas Braa (Anti-Atlas occidental, Maroc). Journal of African Earth Sciences, 24, 271– 284. S OULAIMANI , A., P IQUE , A. & B OUABDELLI , M. 2003. L’extension continentale au Prote´rozoı¨que terminal – Cambrien basal dans l’Anti-Atlas (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, 174, 83– 92. S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic model for the Palaeozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Sciences Letters, 196, 17–33. W ALLISER , O. H., B ULTYNCK , P., W EDDIGE , K., B ECKER , R. T. & H OUSE , M. R. 1995. Definition of the Eifelian–Givetian stage boundary. Episodes, 18, 107– 115. W ENDT , J. 1985. Disintegration of the continental margin of north-western Gondwana: Late Devonian of the eastern Anti-Atlas (Morocco). Geology, 13, 815– 818. W ENDT , J. 1988. Facies pattern and paleogeography of the Middle and Late Devonian in the eastern Anti-Atlas Morocco). In: M C M ILLAN , N. J., E MBRY , A. F. & G LASS , D. G. (eds) Devonian of the World. Canadian Society of Petroleum Geology, Memoirs, 14, 467– 480. W ENDT , J. & A IGNER , T. 1985. Facies pattern and depositional environments of Palaeozoic cephalopod limestones. Sedimentary Geology, 44, 263–300. W ENDT , J. & B ELKA , Z. 1991. Age and depositional environment of Upper Devonian (Early Frasnian to Early Famennian) black shales and limestones (Kellwasser facies) in the Eastern Anti-Atlas, Morocco. Facies, 25, 51–90. W ENDT , J., A IGNER , T. & N EUGEBAUER , J. 1984. Cephalopod limestone deposition on a shallow pelagic ridge: the Tafilalt Platform (Upper Devonian, eastern Anti Atlas, Morocco). Sedimentology, 31, 601– 625.
The ‘Rich’ group of the Draˆa Basin (Lower Devonian, Anti-Atlas, Morocco): an integrated sedimentary and tectonic approach HASSAN OUANAIMI1 & NEZHA LAZREQ2 1
De´partement de Ge´ologie, Ecole Normale Supe´rieure, BP S 2400, Marrakech 40000, Morocco (e-mail:
[email protected] or
[email protected]) 2
Faculte´ des Sciences Semlalia, De´partement de Ge´ologie, BP 2390, Marrakech 40 000, Morocco
Abstract: In this study, based on a sedimentary analysis, an interpretation in terms of sequence stratigraphy is proposed for the ‘Rich’ group of the Early Devonian Draˆa Basin (of Pragian to late Eifelian age). This group is composed of four depositional sequences (third order) containing a transgressive systems tract and a highstand systems tract, both deposited on a storm- and tidedominated shelf. The depositional sequences are related to global sea-level changes but their geometric architecture is controlled by tectonic subsidence and by sediment supply. The basin forms an ENE –WSW channel with two internal high blocks acting on the depocentre shifts. The channel is surrounded by high points and by marine distal condensed series, and bordered by inherited various regional faulted axes. The regional setting corresponds to an extensional passive margin on the northern edge of the West African craton.
The Lower Devonian sequence of the Draˆa Basin forms part of a great Devonian marine deposit on the northern flank of the Tindouf Basin and the southern edge of the Moroccan Anti-Atlas (Fig. 1). The series appear as a continuous Palaeozoic succession, without Caledonian influence, but with slight Hercynian folding, exhumed on common morphological ridges called ‘Rich’ by Hollard (1967, 1981). They exhibit thick clastic formations (1000 m), called ‘Rhenish facies’, in contrast to the thin and more marly– limy series, named ‘Hercynian facies’, of the eastern Anti-Atlas and the High Atlas. Previous studies concentrated especially on the stratigraphic and palaeontological aspect of the Lower Devonian sequences, and the faunal provinces are correlated essentially with European and German ones (e.g. Bultynck & Hollard 1980; Hollard 1981; Weddige 1998; Becker et al. 2004). The great lateral variation of facies in the Anti-Atlas involves various nomenclatures for formations and members of the same age. Sedimentological and structural studies are lacking and a global conceptual framework is needed to simplify understanding of the sedimentary basin evolution and to permit more efficient correlations. Sequence stratigraphy is now a paradigm and can constitute a fundamental approach for regional and global correlations (Wilgus et al. 1988; Einsele et al. 1991; Vail et al. 1991; Posamentier et al. 1993; Weimer & Posamentier 1994). Depositional sequences and systems tracts
can contain various sedimentary facies of varying environments. However, they correspond to genetic units, whose definition can allow us to understand the spatio-temporal distribution of the Lower Devonian formations and the opposition between the western (‘Rhenish’) and eastern (‘Hercynian’) facies of the Anti-Atlas. As faults are rarely exhumed in the Devonian sequences of the Draˆa Basin, this distribution may also allow us to clarify local and regional tectonic context, through the variations of the subsidence or uplift. The purpose of this study is to present the main results of Ouanaimi (2004) on the Lower Devonian sequence stratigraphy of the Draˆa Basin, based essentially on sedimentological observations. This has allowed relative sea-level change interpretation, compared with other global interpretations, and has given a sequence stratigraphic setting for the lateral facies changes in the basin and in the Anti-Atlas. These changes are considered within the tectonic setting of the northern edge of the West African craton (WAC).
Stratigraphic framework As defined by Hollard and presented in his latest synthesis (Hollard 1981), the Rich Group rests on a basal Devonian unit (Lmhaifid Formation), composed essentially of sandy argillaceous rocks and limestones of Lochkovian age. On a synthetic vertical succession, this group consists of four nearly
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 467–482. DOI: 10.1144/SP297.22 0305-8719/08/$15.00 # The Geological Society of London 2008.
468
H. OUANAIMI & N. LAZREQ
Fig. 1. Location of the studied Devonian sites in southern Morocco and northern WAC. AT, Aouinat Torkoz; As, Assa; Ak, Akka; Ta, Tata; Tis, Tissint; FZ, Foum Zguid; Hb, High blocK. Localities (GPS points): 1, Kh. Talha (288210 0000 N, 098580 0300 W), Hassi Boutkerkourt (288230 4500 N, 098540 0400 W), Ma’der Anesis (288240 5200 N, 09851’0700 W to 288220 2700 N, 098500 5700 W), Tazegzaout (288270 0000 N, 098380 5000 W); 2, Sidi Boulaa`ka (288370 4500 N, 098230 4900 W) Kh. Adeken (288300 4000 N, 098260 5600 W); 3, Moumersal (298060 3800 N, 088400 0500 W); 4, Om El Aa`lg (298210 3300 N, 088110 2500 W); 5, Anorhrif (298380 5600 N, 078580 1100 W), Oued Maskaou (298260 0300 N, 088040 1300 W), El Aioun (298390 0600 N, 078530 5200 W); 6, South Sidi Rezzoug (298360 1200 N, 078410 3300 W); 7, Oued Kharouaa` (298460 3200 N, 078240 3500 W) and Reguiba El Allaka (298520 1500 N, 078180 4300 W); 8, El Haidourya– Hamsailikh (298560 1400 N, 078030 2200 W).
identical sedimentary formations (Fm); these consist of, from the base to the top, shelly thin limestones, thin-bedded rhythmites and fine sandstones, and thick-bedded sandstones. These formations are named (Fig. 2) Rich 1 (Assa Fm), Rich 2 (Merzaˆ Akhsai Fm), Rich 3 (El Annsar Fm or Mdaouer El Kbir Fm, Lower Timrhanrhart Fm) and Rich 4 (Nkheila Fm or Khebchia Fm ). They are respectively denoted R1, R2, R3 and R4 in the present paper. Hollard (1981) assigned the Rich Group to a Late Siegenian– Early Eifelian age: a Late Siegenian age for R1 and the basal limestone of R2, an Emsian age for R2 and R3 and the basal limestone limestones of R4, and an Early Eifelian age for R4. New palaeontological results have been published (Lazreq & Ouanaimi 1998; Weddige 1998; Jansen 2000; Becker et al. 2004), and they have been reviewed by El Hassani (2004). The stratigraphical ranges of the ‘Riches’ are (Fig. 2): Upper Lochkovian to Pragian for R1, Upper Pragian to lowermost Emsian for R2, Lower to Middle Emsian for R3 and Upper Emsian to Lower Eifelian for R4 (Becker et al. 2004). Great lateral variation characterizes these Rich Formations and allows distinction between eastern and western facies in the Draˆa Basin. These variations concern essentially R1, R3 and R4 and are summarized in the table of Figure 2, which is modified from Becker et al. (2004).
In this study, numerous sections and parts of sections were considered in the Draˆa Basin between the western Torkoz region and the eastern Foum Zguid
Fig. 2. Synthetic stratigraphic table of the lower Devonian sequences of the Draˆa Basin (according to Becker et al. 2004).
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA
area (Fig. 1). The sections are summarized in Figure 3 and are compared with a section of the central High Atlas (Tizi n’Tichka), representing the eastern reduced facies.
Facies and depositional environments A Rich is a sedimentary unit containing a limy basal part Ra (10– 30 m), a fine-grained and thin-bedded middle part Rb (100 –150 m) and a thick-bedded sandstones upper part Rc (70–120 m) (Figs 3 and 11a). This definition allows a simplified description because the upper clastic units Rb and Rc have identical facies and sedimentary structures and therefore indicate very similar sedimentary dynamics and environmental contexts. The main variations concern essentially the Ra basal limy facies.
The basal limy facies (Ra) The basal limestones of Rich 1 (R1a). They correspond to the decametre-scale ‘calcaires gre´seux’ of Torkoz (Hollard 1981). The basal limestones R1a clearly crop out near Foum El Hassan, at site 3 (Fig. 1) and are about 12 m thick. They rest on Lochkovian green siltstones and their base is characterized by an erosional surface (D1). This basal unit consists of centimetre- to metre-scale layers of bioclastic limestones (rudstones and floatstones), separated by thin silty intercalations. The limestones are very ferruginous and contain an often reworked fauna of brachiopods, trilobites and tentaculites. They show large size stratifications, sometimes of metre scale, corresponding to trough, oblique, sigmoid and herringbone laminations (Fig. 4). These structures indicate a littoral sedimentation under energetic currents, probably induced by waves and/or tidal processes. The basal limestones of Rich 2 (R2a). They appear over the entire basin with typical sections between the Torkoz and Akka areas. In these western regions, the limy level is very distinctive because of rapid changes (D2) over the clastic units of Rich 1 (see Fig. 11b). However, in the eastern areas (El Aioun, oued Karoua`a sites), it is difficult to distinguish because it belongs to a limy condensed Pragian section (Lazreq & Ouanaimi 1998). This level starts with lenticular mudstones and generally forms two shallowing upward decametre-scale sequences (20 –30 m). In these sequences, the lower thin-bedded facies corresponds to wackestones and intercalated mudstones, with tentaculites and small varied fragmented tests, and the upper massive beds are crinoidal packstones and rudstones, with brachiopods, trilobites, bryozoans, nautiloids and sometimes coral elements.
469
Typical sedimentary structures are observed in the ferruginous crinoidal calcarenites of the A. Torkoz and correspond to cross- and herringbone stratifications caused by tidal currents (see Fig. 11c). The parasequences show the passage from a relatively distal and quiet environment to a proximal more agitated environment. Various communities of opensea organisms, crinoidal ‘prairies’ and corals are reworked and mixed by storm or tide currents. The basal limestones of Rich 3 (R3a). These limestones are well defined in the eastern Draˆa Basin (Tata –Tissint) where the Rich 3 formation is complete. They locally start with a microconglomeratic layer of several decimetres thickness (site 6, for example), indicating a basal erosional surface (D3). The limestones are lenticular with varying thickness from a few metres (2–5 m) to a few decimetres. The most condensed beds correspond to black grainstones or rudstones, mostly formed of reworked tentaculites. Laterally, they are laminated calcarenites, containing some channels, HCS (hummocky cross-stratification) and SCC (swaly crossstratification), indicating storm-induced currents. In the southwestern Draˆa Basin (Torkoz–Assa), there is generally a lack of Rich 3 clastic units (R3b and R3c) and the R3 formation is condensed and sometimes amalgamated with the base of Rich 4 in the Ouin– Mesdour Formation. Ouanaimi (2004) considered that the calcareous basal unit R3a corresponds to the first limy level that directly overlies the top of the R3c sandstones. This facies is 10 –15 m thick and shows a local basal erosion, corresponding to a thin microconglomeratic layer (site 1), overlain by extendive fossiliferous and pyritic black limestones. Their microfacies correspond to wackestone with reworked bioclasts and lithoclasts under moderate currents, in a relatively quiet environment. In the Assa region, the basal level was called the ‘Akhal Tergoua Member’ by Becker et al. (2004). The basal limestones of Rich 4. These limestones correspond to the base of the lower member of the Timrhanrhart Formation of the eastern Draˆa facies, at the ‘calcaires a` Sellanarcestes’ level of Hollard (1981). According to various researchers, the Sellanarcestes wenckenbachi level corresponds to a late Emsian age (Hollard 1981; Weddige 1998; Becker et al. 2004; Ebbighausen et al. 2004). In the western Draˆa basin area, this level, about 1.5 –2 m thick, is well exposed on the top of R3c sandstones, except at site 7, where it is separated from these sandstones by decametre-scale green siltstones. It is characterized by rapid lateral changes of thickness and facies. To the south of Tata (site 5), it locally starts with a thin level of ferruginous reddish limestone
470 H. OUANAIMI & N. LAZREQ
Fig. 3. Lithostratigraphic correlation and sedimentary structures of the principal sections of the Devonian Riches (see Fig. 1 for location). Ra, calcareous member; Rb, middle member; Rc, sandy upper member; D, unconformity (sequence boundary); Dm, Eifelian transgressive limestones.
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA
471
Fig. 4. The basal limy facies R1a in the Moumersal section (site 3): detailed section (a) and view of large-scale stratification of the ferruginous calcarenites (b).
(30 cm), with trilobites and stromatolitic laminae. This level is rapidly overlain by decametre-scale grey limestones with goniatites (Sellanarcestes), orthoceras and tentaculites, passing at the top into nodular siltstones or marls. The limy facies corresponds to mudstones, wackestones or floatstones with tentaculites, trilobites, goniatites and crinoids. The sedimentary environment is relatively distal with some intermittent currents reworking some angular quartz and ferruginous grains. In the same locality, Ebbighausen et al. (2004) has reported some thin breccia beds at the base of the Sellanarcestes limestones, representing a basal erosional event. To the north of this locality, the basal limestones correspond locally to a reddish ferruginous bed and massive grey limestones (a few decimetres to 1 m), which are underlain by a minor erosional surface and covered by nodular limestones (7 m) or directly by nodular grey shales. In the west (sites 6 and 7), the basal limestones R4a start locally with coarse-grained beds of marly conglomerates or sandy cross-bedded limestones with ferruginous elements and soft pebbles (1.5 m) (D4). In the southwestern part of the study area (Assa– Torkoz), the R3 clastic deposits are sometimes totally lacking and the basal limy facies R4a (calcaire a` Sellanarcestes) and R3a are both represented in the Ouin Mesdour Formation. However, in some localities, the R4a limy level is separated from the first (R3a) by about 30 m of dark shales and marls, corresponding to the base of the R3b clastic unit. Therefore, the R4a level has a thickness of 25– 30 m and contains limy beds of decimetre to several centimetres thickness, alternating with fossiliferous marls (trilobites, goniatites, fish, etc.) of an open-sea environment. However, the existence of some teracorals suggests the proximity of reef constructions.
A similar marly limestone level is also observed near Assa (15 m), but the lower part of the section is mostly covered by Quaternary sequences. Elsewhere, a detailed section has been described by Becker et al. (2004) in the Bou Tserfine zone, where they distinguished two limy successive members: the Hollardops member and Sellanarcestes member (Fig. 2). This distinction is not clear in the Torkoz area but palaeontological records are well exposed (Jansen et al. 2004). The Dm basal limestones. The basal Dm limestone is the calcareous facies that clearly appears at the top of the R4c sandstones and closes the Rich Group sedimentation in the western area. This unit consists of a thinning and fining-upward sequence, of yellowish limestones and interbedded siltstones or mudstones (about 30 m). Generally, the sedimentation corresponds to floatstones and rudstones containing various communities of fauna (brachiopods, goniatites, crinoids, trilobites, bryozoans, tentaculites, corals, ostracodes and conodonts). It indicates a shallow-water environment probably near coral ‘prairies’. The energy decreases progressively upward, and the facies changes to marly and argillaceous. The Dm limestones terminate the clastic sedimentation characterizing the Rich Group, and start a homogeneous muddy, marly and limy sedimentation of the late Eifelian–Givetian (Becker et al. 2004). In the eastern part of Draˆa Basin, the level is not easily distinguishable because there is a lack of R4b–R4c sandstones and continuous marly – limy sedimentation over the Sellanarcestes R4a level. The limestones show various fossils communities of essentially latest Emsian age (Jansen et al. 2004).
472
H. OUANAIMI & N. LAZREQ
The clastic facies Rb – Rc The fine-grained middle facies Rb of the Riches. These facies correspond to the clastic formations (Rb) that abruptly overlie the basal limestone facies Ra in each Rich. They generally show a thickening and grading upward trend, with mudstones or marls grading up into siltstones and silt– sand fine rhythmites. Some thickness and facies variations characterize in particular the bases of these units. The base of the R1b clastic formation is richer in thin-bedded intercalations (30 cm) of nodular shelly limestones, containing fragments of crinoids, ostracodes, trilobites, bryozoans, tentaculites and corals. Their texture corresponds to floatstones, grainstones or rudstones (site 3, Fig. 1). Some fossiliferous nodular beds also exist at base of the R4b clastic beds. In Rich 2 and 3, the basal mudstones facies are lacking and the Rb units start directly with green or dark siltstones. The upper part of the four Rb units consists of silt –sand laminated graded-rhythmites with small wave and current ripples. They show some decimetre-scale intercalations of bio-quartzarenites with brachiopods, tentaculites and ferruginous sparitic cement. This succession corresponds first to offshore sedimentation, below the storm wave base, with a predominance of decantation processes. However, sometimes abrupt agitations occur, reworking fauna and ‘prairies’ of echinoderms, as ‘storm lags’. At the top, and in spite of bioturbation, characteristic sedimentary structures are well preserved in the upper graded-rhythmites of the Rb members, such as grooves, channels, flutes, asymmetrical ripples, undulating and symmetrical ripples, and interfering cogenetic polygonal ripples. They show evidence of marine unidirectional and wave currents, sometimes typical of storm and/or tide processes (Pedersen 1985; Guillocheau & Hoffert 1988). Therefore, from the base to the top, the middle members indicate a gradual trend from pelagic or predominant offshore facies to more proximal ones. These members gradually become richer in sandstones and pass without interruption into R1c sandy members of the Riches. The upper sandy R1c members. These members are formed essentially of sandstones and intercalated siltstones. With the Rb units, they form thickening and coarsening upward successions. Their thickness varies from one Rich to another, and sometimes within the same Rich until the total disappearance of the member (R1c, R3c and R4c). The sandy beds are often lenticular, with a decimetrescale thickness at the base of the Rc members and
.1.5 m thickness at the top. The sandstones and the intercalated siltstones form metre- to decametre-scale parasequences and show thickening and graining upward trends (shallowing upward sequences). The R4c member is clearly characterized in the field by some frequent seismite levels in the eastern areas (see below). The Rich sandstone units Rc are often laminated, with abundant shell lenses, bioturbation and sedimentary structures: grooves, channels (Fig. 11e), current ripples, wave ripples (Fig. 11d) and interfering cogenetic polygonal ripples. Some erosional and accretional HCS (Fig. 11f) and SCS can also be observed. Oblique laminae with a sigmoid trend are common at the top of the R1c beds. These sedimentary structures are often organized into various elementary sequences of decreasing energy. A typical complete sequence often starts with an erosive base covered by coarse-grained or shelly intervals. This passes into laminated sandstones (e.g. planar laminas, ripples, HCS), which evolve to planar laminae and siltstones. Generally, these sequences are incomplete; they are often thickest at the top of Rc members, where they are amalgamated into massive multi-event metrescale beds. These field structures are characteristic of a storm-dominated shelf (Dott & Bourgeois 1982; Brenchley 1985). The presence of HCS and SCS structures indicates shallow sedimentation above the storm wave base (Bourgeois 1980; Dukes 1985; Greenwood & Sherman 1986). In total, the Rb –Rc members show a continuous and progressive passage, on a storm- dominated shelf, from a distal and relatively deep environment (offshore) toward more proximal shoreface conditions.
Sequence stratigraphy The terminology used in this paper corresponds to that developed in sequence stratigraphy (Van Wagoner et al. 1988; Posamentier & Vail 1988; Wilgus et al. 1988; Vail et al. 1991; Einsele et al. 1991; Posamentier et al. 1993; Weimer & Posamentier 1994) and commonly used in various basin studies. A sequential synthesis is shown in Figures 5 and 6.
Summary of sequence boundaries (D1 – D5) The sequence boundaries are located at the base of the limy members Ra of the Riches (D1–D5; Fig. 2, and are evident at local and regional scales, as follows. The D1 surface is covered by a decimetre-scale layer of limestones with ferruginous reworked
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA
Fig. 5. Sequence stratigraphy synthesis and relative sea-level changes. Comparison with sea-level changes in other continents. Ia, Ib, Ic, transgressive– regressive cycles; Lo-Pr, E0 -Em, M0 -Em and Em/Ei are bio-events. Thickness is about 200 m for a complete R sequence.
Fig. 6. Schematic subsidence variations of the depositional sequences during Dm limestones deposition. D, sequence boundaries; TST, transgressive systems tract; HST, Highstand systems tract; Hb, high block.
473
474
H. OUANAIMI & N. LAZREQ
elements, sometimes of centimetre scale (calcirudite). D2 is a slight erosional surface on an irregular substratum, the products of which occur in the basal thin and lenticular conglomerates or enrich R2a with siliclastic and ferruginous elements. The strata are conformable, but R2a rests on various older units of Rich 1 (Figs 3 and 6). D3 occurs at the base of R3a limestones, which also rest on various levels of the Rich 2 sandstones (R2c). Erosional surfaces are marked locally by thin microconglomeratic layers and by infilling of the palaeotopography. D4 erosion is clearly seen in the eastern basin areas, where the basal limestones R4a locally start with coarse-grained sandstones or conglomerates, sometimes with ferruginous elements. Regional unconformity is indicated by the superimposition of the R4a level on the R3b and R3a facies in the western areas. D5 corresponds to the base of the Dm facies, which transgresses numerous facies of Rich 4, and it is well defined locally by thin basal conglomerates. In the western areas, a correlative unconformity is not well defined because R4 clastic deposits are lacking and homogeneous marly– limy facies directly cover the Sellanarcestes beds (R4a). These unconformities are probably submarine because no true indication of emersion has been detected for this interval, except for discontinuity D1, which is underlain by a basaltic volcanism (tuffs) in the eastern Anti-Atlas and could coincide with the general hiatus separating the Lochkovian from the Pragian in Morocco (Hollard 1967, 1981; Michard 1976). However, the elevated rate of ferruginous deposition at the base of Ra units reveals a subaerial or aerial character in the surrounding areas of the basin (Hollard 1967; Sougy 1969; Beuf et al. 1971; Lecorche´ et al. 1991). These surfaces could have a wide extent and could be correlated with discontinuities described in the lower Devonian units of the central Sahara and western Libya (Beuf et al. 1971; Massa 1988); and probably in western and southwestern neighbouring areas of high relief. They are also transgressive surfaces (TS), with the calcareous facies Ra covering various detrital and limestone facies of the preceding Rich. The subaerial exposure of marine sediments at the sequence boundary is an important criterion for recognizing sea-level cycles as opposed to supply cycles (Schlager 1993). The wide lateral extent of D discontinuities and their features permit us to consider them as type 2 sequence boundaries (SB2) (Posamentier & Vail 1988). They record sea-level falls at a rate that does not exceed the rate of subsidence of the platform in the studied regions.
Depositional sequences and systems tracts Every Rich unit corresponds to a depositional sequence (third-order sequence) in which a limy Ra member constitutes a transgressive systems tract (TST), and the Rb–Rc clastic members, respectively, correspond to early and late highstand systems tracts (HSTs). These systems tracts are deposited during the sea-level rises following two successive falls (D unconformities).
Depositional sequence Rich 1 (R1) This is located between the D1 and D2 boundaries. The TST corresponds to the basal limestones R1a of the Moumersal section, which represent tidal and/or wave-induced deposits. The maximum flooding surface (MFS) occurs in the 10 m above the TST, in deeper nodular and limy marls. The latter pass into a fine-grained and thin-bedded distal facies (R1b) that represents the early HST. This grades up into shallower sandstones R1c, generally storm-generated and forming the late HST, which fill the basin and close the first depositional sequence. This complete sequence R1, of about 250 m thickness, covers only the western basin of Draˆa (Torkoz– Assa). Eastward, it becomes gradually thinner and is truncated at the summit, first by the disappearance of the R1c sandstone, and then by reduction of the R1b unit. In the Tata –Tissint area, the whole sequence is represented by only a few metres at the top of the open-sea limy–marly facies of the eastern domain. This is a good example of proximal to distal evolution in a typical Rich sequence (Figs 3 and 6).
Depositional sequence Rich 2 (R2) The depositional sequence Rich 2 (R2) is located between the D2 and D3 unconformities. The R2a limestones, which contain local phosphatic elements (Hollard 1967; Becker et al. 2004), clearly represent a TST. This shows a SW –NE facies evolution, from tidal-influenced deposits (Torkoz), to shallow-water massive limestone (Assa –Akka) and finally to a very condensed pelagic facies (Tata –Tissint). The MFS interval is situated at the top of R2a, where the early and late HST are respectively represented by R2b and R2c clastic units. This HST corresponds to a progressive vertical passage from distal rhythmites and storm lags to shoreface storm-dominated facies. The early and late HST have a wide lateral extenst and cover the entire Draˆa Basin, without lateral facies changes but with a major increase in thickness towards the NE, near the Tissint –Foum Zguid area.
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA
Depositional sequence Rich 3 (R3) This sequence is bounded by the D3 and D4 unconformities and it is complete in the eastern area, with a TST, represented by R3a basal limestones, and an HST, corresponding to the R3b unit (early HST) and R3c sandstones (late HST). The TST shows important lateral changes of facies and thickness, from black marly limestones in the SW (basal Ouin Mesdour Fm) to condensed limestones with tentaculites in the NE. The filling HST goes vertically from distal mudstones to thin-bedded sandstones (early HST), then to coastal shelly sandstones (late HST). The western equivalent facies of this HST corresponds to condensed black shales and marls of the Ouin Mesdour Formation. According to Becker et al. (2004), the Hollardops limestone mb (new local formation) of eastern Assa also is a late HST.
Depositional sequence Rich 4 (R4) The sequence is well defined between the D4 and D5 unconformities, especially in the western Draˆa Basin, where all members of Rich 4 are represented. The TST corresponds to the Sellanarcestes wenckenbachi limestones, which are remarkably continuous over the basin and cover various units of the underlying systems tracts of the R3 sequence. The MFS occurs in the pelagic nodular black marls and shales of the Torkoz–Assa area, which pass upward into a thick early HST formed of the offshore R4b clastic rhythmites. The R4c sandy coastal sediments are late HST deposits that close the sequence. The eastern lateral equivalent of this HST may correspond to the upper part of the lower Timrhanrhart and the upper Timrhanrhart marls and limestones with Anarcestes and Phacops (Hollard 1981; Becker et al. 2004).
The Dm base of depositional sequence R5 The base of this sequence corresponds to the Dm marly limestones overlying the R4 HST (Rich 4 Sandstones) in the western Torkoz–Assa zone. However, it is not well defined in the eastern areas, where dominantly fine-grained and homogeneous facies follow the R4 sequence. According to Becker et al. (2004), the eastern equivalent base of the R5 sequence may correspond to the Eifelian– Givetian Pinacites beds or ‘Horizon d’Aherich’ (Hollard 1981; Fig. 2).
Eustatic control on depositional sequences The D unconformities, sequences and systems tracts can be represented by a relative sea-level schematic
475
curve. This curve includes five episodes of sea-level fall represented by discontinuities (D), which are followed by five episodes of sea-level rise, corresponding to four R sequences and Dm transgressive limestones (Fig. 5). The amplitude of the relative rises decreases from R1 to R3, as shown by the northeastward shift of their HST (Figs 6 and 7). However, with the HST of the R4 sequence, this amplitude becomes higher and it returns to the R1 state, before the more amplified rise of the Basal R5 sequence. Geometrically, the depositional sequences include only TST and HST in a shelf margin state. There are no lowstands or lowstand wedges systems tracts, because of the wide extent of the Devonian platform. The problem is to determine whether the sealevel changes are global or local. This operation is simpler for the post-Paleozoic, with the existing eustatic chart (Haq et al. 1988). For Palaeozoic basins, we can only compare local relative sea-level changes with those defined in the neighbouring regions and continents. Currently, there are only three curves of references: two curves for Euramerica (Johnson et al. 1985; Walliser, cited by Weddige 1998) and a Barrandian curve (Chlupac & Kukal 1986). Comparatively, it appears that the D1, D3 and D5 unconformities of the Draˆa Basin exist in all curves and/or correspond to chronostratigraphic limits and major Barrandian events. They can be respectively linked with the Lochkovian–Pragian global event (Lo –Pr), the Zlichov basal event (E’ Em; early Emsian) and the Chotec basal event (Em –Ei). Despite their coincidence with periods of sea-level rise in Bohemia, the other discontinuities, D2 and D4, could be also global. D2 could be correlated with the ‘Ia’ second sea-level fall of the Euramerica sequence, and D4 with one of the two intra-Pragian discontinuities of the curve of Euramerica D4, which is also recognized in the Tafilalt (M’-Em middle Emsian; Weddige, cited by Walliser, 1998), could correspond to one of the two intra-Emsian sea-level falls of the Euramerica (‘Ib’ cycle). One of the two latter could be global, as has been suggested by Johnson et al. (1985). The problem of the origin of the early Devonian sea-level variations has been discussed by some workers (Johnson et al. 1985; Brand 1989) but is still not clear. For modern basins, Vail et al. (1991) favoured the oceanic basin volume changes for second- or first-order cycles and climatic and/or water volume changes for thirdorder sequences. However, the relatively hot climate of the early Devonian excludes glacio-eustatic effects for the third-order cycles of the Draˆa Basin. Genetic parameters must be studied in the models of plate tectonics (ridge volume, volcanism and mid-oceanic thermal uplift).
476
H. OUANAIMI & N. LAZREQ
Fig. 7. Restoration of subsidence setting at the end of successive R depositional sequences (at the following transgression) with respect to their relative geographical position. Numbers indicate the locations shown in Figs. 1 and 6. TST, transgressive systems tract; HST, highstand systems tract. Maximum thickness is about 200 m for a complete R sequence.
Sediment supply, subsidence and original geometry of the sequences With eustacy, subsidence (or uplift) and sediment supply are the two other parameters that act on the stacking and architecture of the sequences (Posamentier & Vail 1988; Vail et al. 1991; Steckler et al. 1993). The sequence stacking pattern (Fig. 6) shows that the Rich group is much thicker in the Draˆa Basin compared with the surroundings, and forms an ENE –WSW channel, bordered by high blocks, in the northern edge of the WAC. The most important thickness variations affect the early and late HST, during the periods of filling and high sedimentary flux rates. The relatively constant bathymetry of the HST facies suggests that the irregular subsidence is always accommodated by the sedimentary supplies, which come from high points situated to the south (Reguibat shield), the west and the SW. At the end of every HST subsidence and inputs balance, and the following sea-level fall and consecutive erosion end with restoration of a flattened shelf where a new TST (Ra) is deposited (Fig. 7). This stacking pattern has two major stratigraphical implications, as follows. First, there are horizontal changes from sandstones to shales or calcareous facies with the same
margin profile and polarity for R1, R2 and R4 sequences. Thus the opposition between the eastern and western facies in the Draˆa Basin, on one hand, and between the ‘Rhenish’ and ‘Hercynian’ facies, on the other hand, becomes more comprehensible. This is especially the case for the passages (1) from the R1 thick rhythmites and sandstones to El Aioun condensed marly limestones, (2) from the R3 HST to upper Oui-n-Mesdour Fm and Hollardops mb, and (3) from the R4 HST to reduced facies of the lower Timrharhart Fm. Second, if the subsidence is removed, the model of the original geometry of depositional sequences R and systems tracts could be restored (Fig. 8a). This is conformable with the seismic stratigraphy and sequence stratigraphy models of type II depositional sequences (Fig. 8b). In these models, the TST presents a coastal onlap with diachronous sedimentation (landward younger sets), whereas the HST is prograding, with sigmoid time-lines, toward the open sea, where the two systems tracts (TST and HST) are condensed. The condensation may also affect the fauna that may migrate basinward in highstands because of the high supply rates (Brett 1998). Such TST and HST diachronous sedimentation must be taken into account in stratigraphical correlations. The major problem is that of the distal condensed sedimentation, where sequences are
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA
477
Fig. 8. (a) A2D geometric reconstruction of depositional sequence R after removing subsidence. (b) Similar 2D model for sequence stratigraphy. Maximum thickness is about 200 m for a complete R sequence TST, transgressive systems tract; HST, highstand systems tract.
amalgamated and correlative unconformities are difficult to define precisely. They may correspond to thin hiatus surfaces, hardgrounds and condensed sections (Loutit et al. 1988; Kidwell 1991).
Tectonic control on depositional sequences It is generally acknowledged, without structural evidence, that the Early Devonian is an extensional period. The Rich group was folded during the Hercynian exent, but there is no field evidence of major faulting. However, the stratigraphical signature of tectonic control on third-order sequences is common (Vail et al. 1991). Arguments for tectonic control are deduced from the previously shown subsidence, systems tract shifts, land source shifts, and from some seismite levels, associated with some minor synsedimentary faults. The HST depocentres shift from one sequence to another, and this shift was probably accompanied by shoreline displacements. These changing sources (Fig. 7) may indicate active tectonics in
the surrounding land. The main subsidence variations occur close to the Ordovician basement of the only two Hercynian anticlines that occur in the vast Siluro-Devonian plain of the Draˆa (Fig. 1): Addana –Adrar Zouggar in the SW and Hamsaı¨likh in the NE (Figs 1 and 6). The R1 and R4 HSTs begin to disappear against the first, and the R3 HST starts to subside. The second coincides with the start of the R3 HST subsidence. This suggests a probable tectonic significance in the Early Devonian for the anticline zones. The Hamsailikh zone already seems to be a zone of maximum thinning of the Lochkovian sequences (less than 50 m), which are very thick to the SW of the basin (more than 700 m in Ain De´liouine) and in the Ougarta chain (more than 900 m) (Hollard 1967). Thus extensional high zones are considered in these two regions of the Basin, with alternative activation during R1, R2, R3 and R4 sequences (Fig. 9). The active faults are not expressed in the lower Devonian units because of the rheological behaviour of the Silurian – Lochkovian series. These are predominantly clayey and provide a level of decoupling that
Fig. 9. Block diagram showing east–west and north–south subsidence variations in the Rich group. In the Draˆa Basin, subsidence migration is controlled by two high blocks Hb1 and Hb2, probably located in the Ordovician substratum. Maximum thickness is about 200 m for every complete R sequence. Sil-Lochk, Silurian–Lochkovian shales.
478
H. OUANAIMI & N. LAZREQ
would absorb the vertical propagation of the substratum faults that probably affect the Ordovician and older rocks. The traces of distension are especially recorded in the R4 HST sandstones by syndiagenetic deformed levels (1–5 m thick) considered as seismites (Ouanaimi 2004). They extend over more than 50 km and appear clearly because of their strongly disturbed nature. The levels contain various deformed structures (Plaziat et al. 1990), especially various-sized ball-and-pillow structures (Fig. 11g), sheath folds, monogenetic conglomerates, slumps and convoluted bedding. The deformations took place generally in situ, by water-escape dominated processes, without major gravitational displacement. Their great geographical extent indicates the regional instabilities that generated them, probably in distension, as proved by some synsedimentary associated small normal faults (Fig. 11h). The direction of these faults (N150) is in agreement with the major NE– SW trend of subsidence along the basin. The underground faults would have the same NW–SE to NNW–SSE trend. However, the north–south subsidence variations also suggest the activation of ENE– WSW to ESE– WNW faults, some examples of which occur in the Addana –Adrar Zouggar anticlines (Desthieux 1977) and in the Assa outcrops (Hollard 1967). A subequatorial fault trend is also implicated in the tectonics of blocks that uplifted the Anti-Atlas axis during the Early Devonian (Hassenforder 1987).
Regional tectonic implication in the northern WAC borders At the scale of the WAC, the Draˆa Basin constitutes an NE–SW channel that is bordered by reduced pelagic or more littoral facies and by high points, notably the following: (1) the thinnest shales and carbonates of Maider, in the Jbel Issimour section for example (Plodowski et al. 2000; Jansen 2001; Becker et al. 2004) and Central High Atlas (Gigout 1937; Hollard 1967; Laville 1980; Lazreq & Ouanaimi 1998) to the NE and north; (2) the reduced detrital deposits (Hollard 1981; Bitam et al. 1996) in the Zemmour to the SW; (3) the reduced detrital deposits, fluvial facies or even hiatuses in the south of the Tindouf Basin (Beuf et al. 1971; Bitam et al. 1996). As in the Hoggar, where the Lower Devonian sequence shows a differential subsidence controlled by faults (Beuf et al. 1971), the Draˆa channel may be bordered by three principal fault directions (Fig. 10a). (1) A NW– SE direction is represented by the faulted axis of Bou Azzer –Ougarta and Zemmoul. To the NE of this axis, the Lower
Fig. 10. (a) Sketch map for the Devonian outcrops (according to Hollard 1967) and major basement fault axes in the northern border of the WAC. (b) Tectonic interpretation: the NW–SE normal faults delimiting high blocks (1 and 2) require a transtensional context for the subsiding Draˆa Basin (model shown in inset).
Devonian clastic deposits pass into the eastern Anti-Atlas facies (Maider, Tafilalt). Some major faults with the same trend have also controlled the Late Devonian sedimentation in these areas (Wendt 1985; Wendt & Belka 1991; Baidder et al. 2008). (2) ENE– WSW to east–west directions occur in the central Anti-Atlas (Hassenforder 1987), which could have permitted the passage to the northern reduced facies of the Central High Atlas; also a ENE –WSW fault zone is suspected in the northern margin of the Tindouf Basin (Michard 1976; Hassenforder 1987; Weijermars 1993; Burkhart et al. 2006) that separates the Draˆa depocentres from the reduced and incomplete series of the Saharan Djebilet (Bitam et al. 1996). (3) A western NNE –SSW faulted zone (Zemmour and Atlantic Anti-Atlas) is inherited from Neoproterozoic and Cambrian times (Pique´ 2003; Soulaimani et al. 2003; Belfoul 2005). In this context, and as a multidirectional extension is mechanically improbable, the transtensional
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA
479
Fig. 11. Various aspects of the Lower Devonian Riches: (a) example of upper part of Rich unit: Rich 3 of Anorhrif (Tata); (b) transgressive limestones R2a on the sandy member R1c in the Torkoz area; (c) herringbone stratifications in the transgressive calcareous facies R2a in the same area (Hassi Boutkerkourt); (d) wavy ripples in the sandy member R2c of Tata (Anorhrif); (e) shelly channels in the sandy member R2c (Anorhrif); (f) hummocky cross-stratification in the sandy member R2c (Anorhrif); (g) Metric-sized balls in deformed beds of the R4c seismites (Kheneg Adeken, Assa); (h) N150– 70 N small normal fault associated with the R4c Torkoz seismites.
480
H. OUANAIMI & N. LAZREQ
tectonic regime of the Draˆa Basin is privileged in the hypothesis of reactivated pre-existing fractures (Fig. 10, model shown in the inset). Such a pattern, which is common in the extensional regime of passive margins (Vail et al. 1991), emphasizes the continental stretching in the northern edge of the WAC (Baidder et al. 2008), in a global context of the western Palaeotethys margin (Jurdy et al. 1995; Guiraud & Bosworth 1999). A sequence stratigraphic and tectonic approach including the Lower Devonian sequences of all NW Africa is needed to situate more precisely the margins of the WAC in a global context. Apparently, such work has been done (for Morocco, Algeria and Libya) and recently presented in the PhD thesis of Lubseder (2005), but it has not been published.
Conclusion The Rich group of the Draˆa Basin consists of four depositional sequences of Pragian– early Eifelian age. Each sequence contains a transgressive systems tract overlain by a highstand systems tract, both deposited on a storm- or tide-dominated shelf, on the northern border of the WAC. Depositional sequences are related to global sea-level changes, but their geometric architecture is controlled by tectonic subsidence and a high rate of sediment supply from southern and western lands. The effects of tectonic subsidence, sediment supply and eustatic changes give an integrated explanation for the classical lateral facies changes in the Draˆa Basin and in the Anti-Atlas. The subsidence and supply signature correspond to an extensional context. The latter induces a generally ENE – WSW-trending channel surrounded by lands to the south and the west and by distal condensed series to the north and NE. Into this channel two internal high blocks have an influence on highstand subsidence migration. The regional setting corresponds to a subsiding passive margin on the northern edge of the WAC, where early Palaeozoic and Precambrian fault axes are reactivated in an extensional or transtensional regime on the northern edge of the WAC. The authors thank A. Michard and S. Lubeseder for their constructive revisions, and D. Petit and R. Mjilla for their help.
References B AIDDER , L., R ADDI , Y., T AHIRI , M. & M ICHARD , A. 2008. 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). In: E NNIH , N. &
L IE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 453– 465. B ECKER , G., L AZREQ , N. & W EDDIGE , K. 2004. Ostracods of Thuringian provenance from the Devonian of Morocco (Lower Emsian– middle Givetian; southwestern Anti-Atlas). Palaeontographica, Abteilung A, 271, 1 –109. B ELFOUL , M. A. A. 2005. Cine´matique de la de´formation hercynienne et ge´odynamique de la marge NW du Gondwana (Anti-Atlas, occidental, Sahara marocain: Zemmour– Ouled Dlim et Mauritanides septentrionales). The`se es Sciences, University Ibno Zohr, Agadir. B EUF , S., B IJU -D UVAL , B., DE C HARPAL , O., R OGNON , P., G ARIEL , E. & B ENNACEF , A. 1971. Les gre`s du Pale´ozoı¨que infe´rieur au Sahara. Se´dimentation et discontinuite´s. Evolution structurale d’un craton. Technip, Paris. B ITAM , L., G OURVENNEC , & R OBARDET , M. 1996. Les formations pale´ozoı¨ques ante´-carbonife`res du sousbassin de Djebilet (flanc sud du Bassin de Tindouf, Nord-Ouest du Sahara alge´rien). In: B ITAM , L. & F ABRE , J. (eds) Ge´odynamique du craton ouest africain central ets oriental: he´ritage et e´volution postpanafricains. Me´moires du Service Geologique d’Alge´rie, 8, 91–111. B OURGEOIS , J. 1980. A transgressive shelf sequence exhibiting hummocky stratification: the Cape Sebastian sandstone (Upper Cretaceous), south western Oregon. Journal of Sedimentary Petrology, 50, 681–702. B RAND , U. 1989. Global climatic changes during the Devonian –Mississipian: stable isotope biogeochemistry of brachiopods. Palaeogeography, Palaeoclimatology, Palaeoecology, 75, 311– 329. B RENCHLEY , P. J. 1985. Storm-influenced sandstone bed. Modern Geology, 9, 369–396. B RETT , C. E. 1998. Sequence stratigraphy, paleoecology, and evolution; biotic clues and responses to sea-level fluctuations. Palaios, 13, 241–262. B ULTYNCK , P. & H OLLARD , H. 1980. Distribution compare´ de conodonts et goniatites de´voniens des plaines du Dran du Ma’der et du Tafilalt. Aardkundige Mededlingen, 1, 9 –73. M., C ARITG , S., H ELG , U., B URKHART , R OBERT -C HARRUE , C. & S OULAIMANI , A. 2006. Tectonics of the Anti-Atlas of Morocco. Comptes Rendus Ge´oscience, 338, 11– 24. C HLUPAC , I. & K UKAL , Z. 1986. Reflecting of possible global Devonian events in the Barrandian area, C.S.S.R. In: W ALLISER , O. (ed.) Global Bio-Events. Lecture Notes in Earth Sciences, 8, 169–179. D ESTHIEUX , F. 1977. Etude tectonique et me´talloge´nique du Jbel Addana, Ordovicien des Basines du Dra, Maroc pre´saharien. Notes et Me´moires du Service Ge´ologique du Maroc, 268, 209–236. D OTT , R.H. & B OURGEOIS , J. 1982. Hummocky stratification: significance of its variable bedding sequences. Geological Society of America Bulletin, 93, 663–680. D UKES , W. 1985. Hummocky cross-stratification, tropical hurricanes, and intense winter storms. Sedimentology, 32, 167–194.
ˆ A BASIN, MOROCCO ‘RICH’ GROUP OF DRA E BBIGHAUSEN , V., B OCKWINKEL , J., B ECKER , R. T., A BOUSSALAM , Z. S., B ULTYNCK , P., E L H ASSANI , A. & N U` BEL , H. L. 2004. Late Emsian and Eifelian stratigraphy at Oufrane (Tata region, eastern Dra Valley, Morocco). In: E L H ASSANI , A. (ed.) Devonian neriticpelagic correlation and events in the Draˆa valley (western Anti Atlas, Morocco). International Meeting on Stratigraphy, IUGS. Documents de l’Institut Scientifique de Rabat, 19, 57–78. E INSELE , G., R ICKEN , W. & S EILACHER , A. (eds) 1991. Cycles and Events in Stratigraphy. Springer, Berlin. E L H ASSANI , A. (ed.) 2004. Devonian neritic-pelagic correlation and events in the Draˆa valley (western Anti Atlas, Morocco). International Meeting on Stratigraphy, IUGS. Documents de l’Institut Scientifique de Rabat, 19. G IGOUT , M. 1937. Sur trois affleurements de terrains anciens situe´s entre le Rdat et la Tassaout. Diploˆmes Etudes Supe´rieures. Vigot, Paris. G REENWOOD , B. & S HERMAN , D. J. 1986. Hummocky cross-stratification in the surf zone: flow parameters and bedding genesis. Sedimentology, 33, 33–46. G UILLOCHEAU , F. & H OFFERT , M. 1988. Zonation des de´poˆts de tempeˆtes en milieu de plate-forme: le mode`le des plate-formes nord gondwanienne et armoricaine a` l’Ordovicien et au De´vonien. Comptes Rendus de l’Acade´mie des Sciences, 307, 1909–1916. G UIRAUD , R. & B OSWORTH , W. 1999. Phanerozoic geodynamic evolution of northwestern Africa and the northwestern Arabian platform. Tectonophysics, 315, 73– 108. H AQ , B. U., H ARDENBOL , J. & V AIL , P.R. 1988. Mesozoic and Cenozoic chronostratigraphy and cycles of sea-level changes. In: W ILGUS , C. K., H ASTINGS , B. S., K ENDALL , C. G. St. C., P OSAMENTIER , H. W., R OSS , C. A. & V AN V AGONER , J. C. (eds) Sealevel Changes—an Integrated Approach. Society of Economic Paleontologists and Mineralogists, Special Publications, 42, 71– 108. H ASSENFORDER , B. 1987. La tectonique panafricaine et varisque de l’Anti-Atlas dans le massif de Kerdous (Maroc). The`se es Sciences, University of Strasbourg. H OLLARD , H. 1967. Le De´vonien du Maroc et du Sahara nord-occidental. In: Proceedings of International Symposium on Devonian System. Alberta Society of Petroleum Geologists, Calgary, 1, 203– 244. H OLLARD , H. 1981. Principaux caracte`res des formations de´voniennes de l’Anti-Atlas. Notes et Me´moires du Service Ge´ologique du Maroc, 308, 15– 22. J ANSEN , U. 2000. Stratigraphy of the Early Devonian in the Dra Basins (Moroccan Pre-Sahara). Bulletin de l’Institut Scientifique de Rabat, Se´rie Ge´ologie et Ge´ographie Physique, 20, 36– 44. J ANSEN , U. 2001. Morphologie, Taxonomie and Phylogenie unter-devonisher Brachiopoden aus der Dra-Ebene (Marokko, Pra¨-Sahara) und dem Rheinischen Schiefergebirge (Deutschland). Abhandlungen der Senckenbergishen Naturforschenden Gesellschaft, 554, 1– 389. J ANSEN , U., P LODOWSKI , G., S CHINDLER , E. & W EDDIGE , K. 2004. The Pragian at Assa (SW Dra Valley, Morocco). Bulletin de l’Institut Scientifique de Rabat, Se´rie Ge´ologie et Ge´ographie Physique, 20, 85– 91.
481
J OHNSON , J. G., K LAPPER , G. & S ANDBERG , C. A. 1985. Devonian eustatic fluctuations in Euramerica. Geological Society of America Bulletin, 96, 567 –587. J URDY , D. M, S TEPHANIK , M. & S COTESE , C. R. 1995. Paleozoic dynamics. Journal of Geophysical Research, 100, 17965– 17975. K IDWELL , S. M. 1991. Condensed deposits in siliclastic sequences: expected and observed features. In: E INSEILE , G., R ICKEN , W. & S EILACHER , A. (eds) Cycles and Events in Stratigraphy. Springer, Berlin, 682– 695. L AVILLE , E. 1980. Tectonique et microtectonique d’une partie du versant sud du Haut Atlas marocain (boutonnie`re de Skoura, nappe de Toundout). Notes et Me´moires du Service Ge´ologique du Maroc, 285, 81–183. L AZREQ , N. & O UANAIMI , H. 1998. Le De´vonien infe´rieur de Tizi-n-Tichka (Haut Atlas) et de Laa`youne (Tata, Anti-Atlas, Maroc): nouvelles datations et implications pale´oge´ographiques. Senckenbergiana Lethea, 77, 223–231. L ECORCHE´ , J. P., B RONNER , G., D ALLMEYER , R. D., R OCCI , G. & R OUSSEL , J. 1991. The Mauritanide orogen and its northern extensions (Western Sahara and Zemmour), West Africa. In: D ALLMEYER , R. D. & L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlatives. Springer, Berin, 187– 227. L OUTIT , T. S., H ARDENBOL , P. R., V AIL , P. R. & B AUM , G. R. 1988. Condensed sections: the key to age dating and correlations of continental margin sequences. In: W ILGUS , C. K., H ASTINGS , B. S., K ENDALL , C. G. ST . C., P OSAMENTIER , H. W., R OSS , C. A. & V AN V AGONER , J. C. (eds) Sea-level Changes—an Integrated Approach. Society of Economic Paleontologists and Mineralogists, 42, 183– 213. L UBSEDER , S. 2005. Silurian and Devonian Sequence Stratigraphy of North Africa: Regional Correlation and Sedimentology (Morocco, Algeria, Libya). PhD thesis, University of Manchester. M ASSA , D. 1988. Stratigraphie de Libye occidentale: stratigraphie et pale´oge´ographie. The`se es Sciences, Universite´ de Nice. M ICHARD , A. 1976. Ele´ments de ge´ologie marocaine. Notes et Me´moires du Service Ge´ologique du Maroc, 252. O UANAIMI , H. 2004. Contribution a` l’e´tude ge´ologique du sud marocain. Stratigraphie se´quentielle de bassins de´tritiques pale´ozoı¨ques. Distribution des re´seaux de diaclases dans les chaıˆnes hercynienne et atlasique. The`se es Sciences, University of Marrakech. P EDERSEN , G. K. 1985. Thin, fine grained storm layers in a muddy shelf sequence: an example from the Lower Jurassic in the Stenlille 1 well, Denmark. Journal of the Geological Society, London, 142, 357– 374. P IQUE´ , A. 2003. Evidence of an important extensional event during the Latest Proterozoic and Earliest Paleozoic in Morocco. Comptes Rendus Ge´oscience, 335, 865– 868. P LAZIAT , J. C., P URSER , B. H. & P HILOBBOS , E. 1990. Seismic deformation structures (seismites) in the syn-rift sediments of the NW Red Sea (Egypt). Bulletin de la Socie´te´ Ge´ologique de France, VI, 419–434. P LODOWSKI , G., B EKCKER , G., B ROCKE , R. ET AL . 2000. The section at Jebel Issimour (NW Maı¨der,
482
H. OUANAIMI & N. LAZREQ
Early to Early Middle Devonian). First results with respect to lithology and biostratigraphy. In: E L H ASSANI , A & T AHIRI , M. (eds) Moroccan Meetting of the SDS-IGCP 421. Excursion Guidebook. Notes et Me´moires du Service Ge´ologique du Maroc, 329. P OSAMENTIER , H. W. & V AIL , P. R. 1988. Eustatic control on clastic deposition II—Sequence and system tract models. In: W ILGUS , C. K., H ASTINGS , B. S., K ENDALL , C. G. St. C., P OSAMENTIER , H. W., R OSS , C. A. & V AN V AGONER , J. C. (eds) Sealevel Changes—an Integrated Approach. Society of Economic Paleontologists and Mineralogists, Special Publications, 42, 125 –155. P OSAMENTIER , H. W, S UMMERHAYES , C. P., H AQ , B. U. & A LLEN , P. (eds) 1993. Sequence Stratigraphy and Facies. Special Publication, International Association of Sedimentologists, 18. S CHLAGER , W. 1993. Accommodation and supply—a dual control on stratigraphic sequences. Sedimentary Geology, 86, 111–136. S OUGY , J. 1969. Grandes lignes structurales de la chaıˆne des Mauritanides et de son avant-pays (socle pre´cambrien et sa couverture infra-cambrienne et pale´ozoı¨que): Afrique de l’Ouest. Bulletin de la Socie´te´ Ge´ologique de France, 7, 133–149. S OULAIMANI , A., B OUABDELLI , M. & P IQUE´ , A. 2003. L’extension continentale au Ne´o-Prote´rozoı¨que supe´rieur–Cambrien infe´rieur dans l’Anti-Atlas (Maroc). Bulletin de la Socie´te´ Ge´ologique de France, 174, 83– 92. S TECKLER , M. S., R EYNOLDS , D. J., C OAKLEY , B. J., S WIFT , B. A. & J ARRARD , R. 1993. Modelling passive margin sequence stratigraphy. In: P OSAENTIER , H. W., S UMMERHAYES , C. P., H AM , B. U. & A LLEN , P. (eds) Sequence Stratigraphy and Facies Associations. Special Publication, International Association of Sedimentologists, 18, 19– 41. V AIL , P. R., A UDEMARD , F., E ISNER , P. N. & P EREZ -C RUZ , C. 1991. The stratigraphic signatures
of tectonics, eustacy and sedimentology—An overview. In: E INSELE , G., R ICKEN , W. & S EILACHER , A. (eds) Cycles and Events in Stratigraphy. Springer, Berlin, 617 –659. V AN W AGONER , J. C., P OSAMENTIER , H. W., M ITCHUM , R. M., V AIL , P. R., S ARG , J. F., L OUTIT , T. S. & H ARDENBOL , J. 1988. An overview of the fundamentals of sequence stratigraphy and key definitions. In: W ILGUS , C. K., H ASTINGS , B. S., K ENDALL , C. G. St. C., P OSAMENTIER , H. W., R OSS , C. A. & V AN V AGONER , J. C. (eds) Sea-level Changes—an Integrated Approach. Society of Economic Paleontologists and Mineralogists, Special Publications, 42, 39– 45. W EDDIGE , K. 1998. Devon-Korrelationstabelle. Senckenbergiana Lethea, 77, 289– 326. W EIJERMARS , R. 1993. Estimation of paleostress orientation within deformation zones between two mobile plates. Geological Society of America Bulletin, 105, 1491– 1510. W EIMER , P & P OSAMENTIER , H. (eds) 1994. Siliclastic Sequence Stratigraphy. Recent Developments and Applications. American Association of Petroleum Geologists, Memoirs, 58. W ENDT , J. 1985. Disintegration of the continental margin of north western Gondwana: Late Devonian of the eastern Anti-Atlas (Morocco). Geology, 13, 815–818. W ENDT , J. & B ELKA , Z. 1991. Age and depositional environment of Upper Devonian (Early Frasnian to Early Famennian) black shales and limestones (Kellwasser facies) in the Eastern Anti-Atlas, Morocco. Facies, 25, 51–90. W ILGUS , C. K., H ASTINGS , B. S., K ENDALL , C. G. St. C., P OSAMENTIER , H. W., R OSS , C. A. & V AN V AGONER , J. C. (eds) 1988. Sea-level Changes—an Integrated Approach. Society of Economic Paleontologists and Mineralogists, Special Publications, 42.
Orogen-parallel tectonic transport: transpression and strain partitioning in the Mauritanides of NE Senegal M. DABO1, M. GUEYE2, P. M. NGOM1 & M. DIAGNE2 1
De´partement de Ge´ologie, Faculte´ des Sciences et Techniques, Universite´ Cheikh Anta Diop de Dakar, BP 5005, Dakar-Fann, Senegal (e-mail:
[email protected])
2
Institut des Sciences de la Terre, Faculte´ des Sciences et Techniques, Universite´ Cheikh Anta Diop de Dakar, BP 5005, Dakar-Fann, Senegal Abstract: New structural, metamorphic, finite-strain analysis and kinematic data from the southern part of the Mauritanides in Senegal (Bakel area) have revealed a history reflecting different tectonic regimes. The structural framework is expressed in the form of shear-zone systems comprising rectilinear NE–SW trending thrusts and NE– SW-trending ductile sinistral strikeslip faults associated with north– south- to NE–SW-trending transpression belts. The analysis of ductile fabrics reveals the presence of lineation trajectories that show a change from a tangential deformation towards the marginal zones in the Northern and Southern Domains of the belt to sinistral oblique transpression in the Central Domain. The deformation within the belt is strongly partitioned into shear zones distributed across the belt defining structural domains with accompanying large-scale folds and thrusts. This exemplifies wrench-style tectonics dominated by crustal-scale strike-slip shear zones in the orogen core and temporal and spatial progression outwards to the marginal zones. Thus, the Variscan orogeny in the southern part of the Mauritanides belt in Senegal developed through sinistrally oblique transpressive shearing along the borders of the orogen core leading to the exhumation of ultramafic rocks between large shear zones.
Transpression describes a style of deformation that involves collisional orogenesis accompanied by strike-slip shear across a zone. Two end-members of tranpression identify deformation of an upright body within a vertical zone boundary (Sanderson & Marchini 1984) or deformation of an oblique body in non-vertical zones, known as inclined transpression (Dutton 1977; Jones et al. 2004). Transpressional deformation shows strike-slip shearing in the internal part, through progressively more oblique convergence to high-angle overthrusting onto the foreland (e.g. Tapponnier et al. 1982; Woodcock 1986; Holdsworth & Strachan 1991; Jones & Strachan 2000; Vassallo & Wilson 2002). The processes and results of transpression have been widely studied in both the brittle and ductile sections of continental crust (e.g. Harland 1971; Hudleston et al. 1988; Oldow et al. 1990; D’Lemos et al. 1992; Tikoff & Greene 1997; Holdsworth et al. 1998, 2002). One of the specific consequences of transpression is the partitioning of strain into domains that are predominantly transcurrent associated with domains that are predominantly compressive (e.g. Lister & Williams 1983; Cobbold et al. 1991; Jones & Tanner 1995; Tikoff & de Saint Blanquatt 1997). However, in oblique transpression, the geometric relationships among the resulting structures are too complex to be modelled successfully (Little et al. 2002b;
Jones et al. 2004). The consequences and importance of oblique and orogen-parallel tectonic motions have led to new insights into mountain-building processes. In this study, new data are presented together with detailed structural analysis of superimposed deformation in the Mauritanide belt of southeastern Senegal, allowing the characterization of Hercynian-age structures and their relative importance during the Variscan deformation.
Geological setting The southern part of the Mauritanides belt in Senegal defines the NE–SW trending arm of the Hercynian orogen, which extends for some 40 km from Thianaff village in the north to Kidira in the south (Fig. 1). Several geological studies have presented details of the lithostratigraphy of the Mauritanides belt in this region (Bassot 1966; Agassiz 1970; Chiron 1973; Dia 1984; Le Page 1988), the metamorphic deformation and structural evolution of the belt (Petkovic 1971; Dia et al. 1979; Dia 1984; Le Page 1988; Burg et al. 1993; Diop 1996) and the geochronological framework of this part of the Mauritanides belt (Dallmeyer & Lecorche´ 1989, 1990a, b). Recent studies have shown that the geodynamic evolution of the region
From: 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 0305-8719/08/$15.00 # The Geological Society of London 2008.
484
M. DABO ET AL.
(640–550 Ma); (5) the development of Variscan thrusts and nappes and the reactivation of older structures at c. 325–275 Ma (Lecorche´ et al. 1991; Blanc et al. 1992; Villeneuve 1993; Diop 1996). Rocks of the Mauritanides belt were investigated near Bakel, using conventional mapping, structural analysis and strain measurements, to characterize the nature and style of deformation in the Mauritanides belt NE of Senegal.
The Bakel Mauritanides belt
Fig. 1. Simplified geological map of the Mauritanides belt. Modified after Lecorche´ et al. (1989). 1, Basement; 2, West African mobile belts; 3, Taoudeni basin; 4, Mesozoic and Cenozoic basins.
involved: (1) deposition of continental sediments over the Birimian basement at c. 1100– 700 Ma; (2) extension and rifting at c. 700 –680 Ma resulting in the outpouring tholeiitic basalts and emplacement of ophiolites and associated alkaline formations at c. 680 Ma (Remy 1985); (3) collision and oceanic closure between 680 and 640 Ma resulting in the accretion of the volcano-plutonic arc complexes; (4) erosion and molasse deposition
In the Bakel area the Mauritanides belt comprises the following rock units: (1) the Faleme Series (or paraautochthonous unit), which crops out south of the Marsa Thrust Fault, comprises power Cambrian age volcanosedimentary and sedimentary rocks (Bassot 1966; Deynoux 1978; Le Page 1983) overlying the early Proterozoic series of the Kenieba Inlier; (2) the Bakel series (or internal and external units) north of the Marsa Thrust Fault comprising conglomerates, sandstones, quartzites associated with tholeiitic basalts and ultramafic klippen (Bassot 1966; Le Page 1983; Diop 1996). Radiometric dates on quartzites are c. 350–206 Ma (Bassot et al. 1963). The dominant tectonic feature in the area is an ESE-verging thrusting event, which resulted in folding and epizonal metamorphism of the various rock units (Fig. 2). The structure of the area is strongly controlled by the general effects of the Mauritanides orogeny, characterized by NNE–SSW shortening resulting in thrust-related tectonics; this is responsible for the development of the regional-scale gentle folds, which are characterized by a NNE–SSW-trending subhorizontal axis and NW-dipping axial plane (Fig. 3). The thrusts propagate along lateral, frontal and oblique ramps, which developed contemporaneously with greenschist-facies metamorphism. The intrusion of magmatic rocks within the thrust package is associated with the formation of sericite-schist, chlorite-schist, serpentinite and talc-schist. The thrust planes are filled by thick flat quartz vein systems. The NE – SW-trending structures in the cover units developed thin-skinned tectonics, associated with greenschist metamorphic conditions (Burg et al. 1993; Diop 1996). These Hercynian thrusts resulted in a partitioned strain gradient with the intensity of deformation increasing towards the thrust planes (Burg et al. 1993). Moreover, there is a slight increase in the degree of metamorphism and deformation from NW to SE. Structural features and kinematic indicators such as cleavage – bedding relationships, asymmetrical shear fabrics and flat – ramp and duplex geometries (e.g. Boyer & Elliot 1982) indicate a NW to SE tectonic transport (Burg et al. 1993).
TECTONIC EVOLUTION OF THE MAURITANIDES
485
Fig. 2. Geological map of the area south of Bakel, showing major lithological divisions. 1, Foliation orientation and dip; 2, minor fault; 3, mega-fault; 4, quartzites; 5, jaspilites; 6, schists; 7, chlorite schists; 8, metagabbros; 9, serpentinites; 10, parautochthonous sedimentary rocks (conglomerates and sandstones).
The southern part of the Mauritanides belt in the Bakel area can be subdivided into four domains on the basis of tectonic style. These are three NE – SW-trending parallel domains (Northern, Central and Southern Domains). In addition, three subdomains are defined within the Central Domain (Fig. 4).
The Northern Domain The Northern Domain (ND) is underlain by quartzites that are deformed into upright folds as a result of NE–SW-directed shortening. The eastern margin of the Northern Domain is marked by a 0.5 km wide shallow NW-dipping, SE-vergent
Fig. 3. Regional transect across the Mauritanides belt from Bakel to Marsa (location is indicated in Fig. 2). Key as in Figure 2. BTF, Bakel Thrust Fault; ASZ, Aı¨re´ –Diabal Shear Zone; KSZ, Koughany Shear Zone; SSZ, Samba–Kontaye Shear Zone; GSZ, Gabou Shear Zone; MTF, Marsa Thrust Fault.
486
M. DABO ET AL.
Fig. 4. Map showing the main domains and subdomains. 1, fault; 2, Northern and Southern Domain; 3, Central Domain; 3a, Diabal– Gounia subdomain; 3b, Samba Kontaye– Gabou subdomain; 3c, Samba Niame´ –Gue´tie´ subdomain; 4, Eastern Domain (parautochthonous).
thrust system, known as the Aı¨re´ – Diabal Shear Zone (ASZ) (Fig. 5). The Aı¨re´ –Diabal Shear Zone exposes a complex system of SE-verging folds and associated top-to-the-SE thrusts thought to have formed at relatively low metamorphic grade (Burg et al. 1993; Diop 1996).
The Central Domain
Structural elements within the Northern Domain. The dominant structure in the Northern Domain is a foliation defined by an alignment of micas and quartz + feldspar ribbons. The large-scale structures in this domain are thrust faults, responsible for the imbrications of the original sedimentary pile (Fig. 6a and b). The foliation in this domain is flat-lying and contains a welldeveloped N3208E-trending stretching lineation. In many shear zones in this domain, the foliation developed as fine-grained shear bands and foliation seams, or as mylonitic to ultramylonitic zones in the high-strain regions. The mineral stretching lineation is defined by mineral aggregates and ribbons of quartz and feldspar subgrains, trains of fine micas and aligned laths of mica (Fig. 6c). The lineation is regionally pervasive, typically subhorizontal to shallow NW-plunging, and interpreted to be the vector of the regional tectonic transport (Fig. 5). The analysis of thin sections of oriented samples reveal asymmetrical S – C fabrics, mylonitic
The Central Domain (CD) is a 10 –12 km wide and 30 km long structure trending NE–SW (Fig. 4). This zone is bounded by Senegal river in the NE and the Mesozoic to Cenozoic SenegaloMauritanian basin in the west. It comprises a wide network of anastomosing NNE–SSW-trending shear zones developed in micaceous quartzites in the west and chlorite-schist in the east. The rocks show intense S-fabrics that developed during sinistral transpression culminating in the exhumation of the ultramafic rocks. The CD is delineated by the Gabou Shear Zone (GSZ) in the east and the Koughany Shear Zone (KSZ) in the west. This domain is considered to be a wrench-style orogen core composed of shear zone bounded panels of a deformed and metamorphosed harzburgite sequence (Petkovic 1971) with a high proportion of chromite often occurring as serpentinite. Detailed mapping of the central corridor within this domain shows that the pattern of deformation is markedly heterogeneous on a mesoscale, particularly in the area between Gabou and Diabal. The CD is subdivided into three subdomains bounded by
fabrics and intense strain softening features such as dynamic recrystallization of quartz and feldspars.
TECTONIC EVOLUTION OF THE MAURITANIDES
487
Fig. 5. Structural map with the trend of the main shear zone (lower hemisphere), and stereographic projection showing the main foliation and lineation in the study area. The foliations show a constant NE–SW orientation despite the variable orientation of the stretching lineation. 1a, Foliation pole; 1b, lineation orientation and dip; 2, minor fault; 3, mega-fault.
NE–SW-trending faults (Fig. 4): (1) the Diabal Gounia subdomain in the north bounded by the Aı¨re´ –Diabal and Koughany Shear Zones; (2) the Samba Niame´ –Gue´tie´ subdomain between the Koughany and Samba –Kontaye– Gabou Shear Zones in the centre; (3) the Samba Kontaye– Gabou subdomain in the south. In these subdomains, regional deformation is accommodated by ductile to ductile –brittle strike-slip shear zones, lateral thrusts and isoclinal folding within a 3 km wide zone. Basement rocks represented by ultramafic rocks are exposed within a metamorphic complex bounded by NE–SW-trending shear zones in the Diabal–Guethie shear system (Fig. 5). Structural elements within the Central Domain. The structural geometry of the Central Domain is characterized by variable orientations of the lineation
and relatively constant orientation of foliations. The dominant structural feature of the Central Domain is the west-dipping, crustal-scale mylonite to ultramylonite sinistral strike-slip shear zones, which dissect it into panels or subdomains. The consequences are the formation of an anastomosing network of shear zones, with lineations defined by slicken-fibres and stretched minerals along the fault and shear zone planes. In the highstrain zones, lineations plunge subhorizontally to the NE. The associated kinematic indicators preserve a predominantly sinistral strike-slip sense of shear defined by shear band cleavages, asymmetric mantled porphyroclasts, flanking folds and asymmetric boudins (Lister & Snoke 1984; Platt 1984; Simpson 1984; Passchier & Simpson 1986; Passchier 2001; Goscombe et al. 2004) (Figs 4 and. 5).
488
M. DABO ET AL.
Fig. 6. Schematic sketch synthesizing the typical features of the major shear zones in the marginal domains. (a) Marsa Thrust Fault (MTF); (b) Bakel Thrust Fault (BTF); (c) thin section of d-type recrystallized trails illustrating top-to-SE sense of shear during thrusting. Fo, foliation; Ft, compression joint; Le, stretching lineation; Lc, crenulation lineation; Js, stylolite joint; Pb, boudinaged fold; Qz, quartz; Se, sericite; P, foliation pole.
The Samba Niame´ –Gue´tie´ subdomain is a NE– SW-trending anastomosing system of faults that preserve serpentinite zones and show evidence of both strike-slip and compressional deformation. The interaction between the main discontinuities and synthetic oblique shear zones such as the Samba Kontaye Shear Zone (SSZ) (Fig. 5) forms contractional strike-slip duplexes, which show an
anastomosing pattern. In the north, near the Senegal river, the shear zones are bent to a NNW– SSE trend. In the Samba Niame´ area (Fig. 5), secondary folds axes and mineral lineation occur at low angles to each other defining type II interference fold patterns in the form of upright to gently inclined early folds and recumbent isoclinal folds (Fig. 7a and b). The foliation (S2) is axial
TECTONIC EVOLUTION OF THE MAURITANIDES
planar to small isoclinal folds with axes parallel to the lineation. Sheath folds are observed in sections perpendicular to the lineation. The quartzites of Samba Niame´ area are cut by oblique faults that expose slickenside striations that plunge between 408E and 508E to the SSE. The analyses of the fracture sets show that the quartzites were deformed in semiductile conditions, as the obtuse rather than acute angle between conjugate sets contains the axis of principal compression (Fig. 8). In this subdomain the compression axis bisecting conjugate arrays lies a few degrees clockwise of being orthogonal to the local zone boundaries. This is consistent with high-angle oblique sinistral shortening (i.e. pure shear dominated transpression) (Fig. 8). The Samba Kontaye – Gabou subdomain is dominated by folds with highly curvilinear hinges that are a product of locally constrictional strain (1 , k , 1). A sheath fold origin for the curvilinear folds seems likely. The parallelism of the fold axes to the GSZ is evidence of regional strain partitioning, and
489
the folds define a domain of coaxial shear strain in which horizontal shortening across the deformation zone caused the domain to thicken vertically. The coexistence of folding and shearing can be understood as the effect of deformation partitioning (Bell 1981) at mid-crustal depth. Upright folding in the rocks within this subdomain has produced vertical thickening and horizontal thinning and stretching parallel to the GSZ, the latter being facilitated by displacement accumulated in an anastomosing system of high-strain zones. The structural record within the Central Domain indicates that deformation is related to a sinistral strike-slip regime. The stretching lineations associated with this regime plunge 108 to the SW.
Southern Domain The Southern Domain (SD) is dominated by jaspilites and is delineated by the shallow SE-verging Marsa Thrust Fault (MTF). The thrust zone is some metres thick and is marked by mylonites derived from jaspilites.
Fig. 7. (a) Curvilinear folds of the Samba Kontaye area. (b). Type II interference pattern resulting from the refolding of first generation folds. (c) Asymmetrically boudinaged lenticular pods of massive serpentinite. (d) Extensively fractured feldspar with quartz- and chlorite-filled cracks. ECC, Extension crenulation cleavage; Fd, feldspar.
490
M. DABO ET AL.
Fig. 8. Lower hemisphere equal-area stereoplots of the best-fit finite strain axes and principal planes of strain discussed in the text. Estimate of the strain axes: best-fit orientation: x: N0528 –328NE; y: N2208 –608SW; z: N3208– 058NW. 1, Measured slickensides; 2, maximum (x); 3, intermediate (y); 4, Minimum (z); 5, zy-plane; 6, xy-plane; 7, minor strike-slip conjugate fault; 8, normal microfault.
Structural elements within the Southern Domain. The dominant structure in this domain is the presence of a NE – SW-trending shear zone system. The Marsa thrust Fault is a 100 m wide steeply dipping shear zone with a welldeveloped foliation and stretching mineral lineation (Fig. 6b). The sense of shear in this zone is parallel to the lineation direction and the movement is interpreted as a NW-side-up thrust fault as deduced from the shear sense indicators parallel to the lineation (Burg et al. 1993; Diop 1996). The main regional planar and linear fabrics are preserved in most rocks and consist of mylonitic foliation, schistosity, chocolatetablet boudinage, stretching and mineral lineations, and intrafolial, sheath and isoclinal folds. Kinematic indicators such as S – C foliation, mica fish, asymmetrical boudins and shear bands suggest a dominant top-to-the-SE thrust (Fig. 6c).
Finite-strain analysis In this section we present results of finite-strain measurements within the various zones and we compare the results on finite strain in each zone with its position in the domain. The occurrence of deformed quartz grains in a fine-grained matrix within the quartzites and deformed pebbles in conglomerate layers provides us with a means of estimating the finite strain in the various domains. The shape of the finite-strain ellipsoid was measured from ellipsoidal markers, or approximated from microtectonic observations (Flinn 1962).
Method Measurements were carried out on sample surfaces parallel to the finite-strain axes, assuming that the principal axes of the fabric of the rock coincide
TECTONIC EVOLUTION OF THE MAURITANIDES
with those of the finite-strain ellipsoid; that is, S is parallel to the xy-plane and L is parallel to the x-axis (x . y . z) (Flinn 1962). Quartzites samples were selected in each domain. Two sections in each sample were studied: (1) perpendicular to the foliation and parallel to the mineral lineation (xz section of the finite-strain ellipsoid); (2) perpendicular to both foliation and lineation (the yz section of the finite-strain ellipsoid).
Finite-strain ratios Finite-strain ratios were calculated from the two principal strain ellipsoid planes using a harmonic method (Lisle 1977; Ramsay & Huber 1983, p. 80). The Ramsay plot (Ramsay 1967) was used in analysing the fabrics. These ratios were used to construct the k-parameter of Flinn (1962) to estimate the shape of the finite-strain ellipsoid.
Fry analyses The method of Fry analyses (Fry 1979) is based on the centre –centre separations of clasts in a deformed matrix and how these approximate a finite-strain ellipse in two dimensions. Threedimensional strain analyses are achieved by combining the measured results of the least two orthogonal planes (i.e. xz-plane and xy-plane). The points (centre of the objects) located on sample surfaces are digitized. The cut-off ellipse is traced by a high density of points on the calculated Fry plot. The ellipse axes obtained are always parallel to the macroscopic fabric axes (L, S).
Results The results presented in Table 1 are plotted in Figure 9.
491
In the Northern Domain, the values of axial ratios of quartzites and conglomerate pebbles give oblate finite-strain ellipsoids. The datas suggest a shape parameter k , 1 for the finite-strain ellipsoid (Table 1). In the two sections, the pressure shadows around quartz clasts are conspicuous and the strain markers are strongly elliptical. In the Central Domain, the finite-strain ellipsoid increases in roundness from the borders to the centre. Fry analyses plot very close to the k ¼ 1 line, suggesting that at low-strain sites, bulk strain was plane strain. Other mylonites within the Central Domain showed exceptionally developed linear fabrics (L-tectonites), thus implying a significant amount of constriction in their deformational history. The lineations are subhorizontal and trend NE –SW, parallel to the x-axes of the finite-strain ellipsoid. In the Southern Domain conventional strain markers are lacking; however, we used the chocolate-tablet boudinage of some quartz veins to estimate the strain (Fig. 10). Numerous boudinage structures occur throughout this area. Many of the boudin networks may be described as chocolate-tablet boudinage structures with a subvertical long axis (S . L tectonites (Flinn 1962). This suggests a shape parameter k , 1 for the finitestrain ellipsoid. The finite-strain analysis shows that rocks from the Gabou– Bakel area exhibit ellipsoids ranging from slightly oblate symmetry to plane strain (k-parameters vary from 0.4 to 1.2 for quartz). The increase of deformation intensity is connected with the disappearance of augen structures and the development of banded rocks in the Central Domain. An L2 lineation developed, characterized by stretching of quartz aggregates. In the Central Domain the fabric development can be attributed to a late transcurrent dominated transpressional phase of the deformation in the
Table 1. Results of finite-strain measurements Domain and subdomain Axial ratios analysis Northern Domain (ND) Central Domain (CD): Diabal –Gounia subdomain CD: Samba Niame´ –Gue´tie´ subdomain CD: Samba Kontaye –Gabou subdomain Fry analysis ND CD: Diabal Gounia sub domain CD: Samba Niame´ –Gue´tie´ subdomain CD: Samba Kontaye –Gabou subdomain
Sample
x
y
z
k
Ol: conglomerate BK: quartzites D17: quartzites SD5: quartzites SN4: quartzites SN31: quartzites SK3: quartzites Go7: quartzites
2.4 2.55 3.02 2.6 4.88 4.39 2.49 2.68
4.62 1.9 2.03 1.85 2.07 1.91 1.9 2.09
1 1 1 1 1 1 1 1
0.66 0.37 0.47 0.47 1.26 1.40 0.34 0.31
BK: quartzites SD5: quartzites SN4: quartzites SK63: quartzites
2.4 2 4.6 2.54
1.7 1.6 1.84 2
1 1 1 1
0.7 0.41 1.78 0.27
492
M. DABO ET AL.
Fig. 9. Map showing the average position of the strain ellipse within and between the different domains. Flinn plot illustrates flattening in the marginal domains and constrictional deformation in the Central Domain.
Mauritanides belt, which created the current structural geometry of the belt by progressive sinistral shearing. In contrast to the previous models (polyorogenic folding and penetrative deformation (Le Page 1988) or thin-skin tectonics (Burg et al. 1993)), we argue that the tectonic evolution of the area is the result of a progressive deformation involving a distinct early period of SSE-directed thrusting preceding the development of the Central Domain. Lower temperature flow and high fluid activity during deformation are evident from syntectonic mineral growth. Serpentinites were transformed to chlorite and/or talc schist, and feldspars in gneisses were entirely replaced by dissolution and precipitation of quartz in asymmetrical boudins necks and chlorite-coated discrete shear bands.
Discussion and tectonic interpretation The Bakel area of the Mauritanides belt is composed of distinct terranes, many of which are
bounded by high-strain zones with a regional foliation folded into upright folds and a stretching lineation generally striking N3208E. However, east of Diabal, the lineation trends towards N0458E. There is an almost 908 swing in mineral stretching lineation orientation, from NE with shallow to horizontal plunges in the Central Domain, which is regarded as the strike-slip core of the orogen, to down-dip NW plunges in the Northern and Southern Domains (Figs 4 and 5). In general, the variation in lineation trends in the study area (Fig. 5) is more or less progressive and shows some similarity to the results of models for strain partitioning in inclined or oblique transpressional settings (e.g. Dewey et al. 1998; Holdsworth et al. 2002). The variation in strain ellipsoid shape observed in the study area may be related to strain partitioning. In the Central Domain the slight subhorizontal stretching and the k ¼ 1 finite-strain ellipsoid indicate a moderate finite simple shear. In this zone sinistral oblique transpression is the main component of strain. Further evolution from oblate to plane strain may be related to an increase
TECTONIC EVOLUTION OF THE MAURITANIDES
493
Fig. 10. Geological cross-sections for the Northern Domain in the Diabal area. The fabric is characterized by thrust faults with exhumed ultrabasic rocks. The fault zone shows numerous folded and chocolate-tablet boudinaged quartz veins (a, b) suggesting its strong extensional flattening; this was followed by folding during the late deformational event. F, fault; S1, schistosity; other symbols as in Figure 6.
of the transcurrent shear associated with an extensional laminar flow, as indicated by the occurrence of shear bands. Other notable changes are the lithological and metamorphic grade; for example, the Northern and Southern Domains are underlain by siliceous rocks (quartzites, jaspilites) whereas the Central Domain is dominated by metamorphosed harzburgite and chlorite sequences (Fig. 2). The juxtaposition in the study area of transcurrent and compressive domains is viewed as a strain partitioning linked to both the transpressive character of the main regional deformation event and the rheological contrast between the chlorite schists and quartzite country rocks. We believe that strain partitioning within the area was linked to the ductility contrast between these formations; that is, the lithological competence contrasts have induced weakness that has created the zones required for partitioning the regional strain (e.g. Jones & Tanner 1995). Thus, the Bakel area yields an outstanding example of
strain partitioning inside a single and relatively small area. Partitioning is coeval with the late transpressional event that characterizes the southern Mauritanides belt, and resulted from a rheological contrast in the country rocks. In the Central Domain the shear sense indicators such as shear band cleavages, asymmetric mantled porphyroclasts, flanking folds and asymmetric boudins everywhere are consistently sinistral along shallow lineation zones (Fig. 7c and d). In summary, the simple shear component of the regional transpression was partitioned into the Central Domain (subdomain dominated by pure shear) and the other domains dominated by simple shear deformation. In the Samba Niame´ area we can demonstrate that strain represented by ductile – brittle fracturing is triaxial, with a maximum extensional strain that is horizontal and oriented broadly parallel to the transpressionzone boundaries (Fig. 8). In summary, because the simple-shear component of regional transpression was partitioned into the
494
M. DABO ET AL.
Central Domain, the flanking domains are dominated by pure shear deformation (Figs 11 and 12). The new results also indicate that the Northern and Eastern Domains formed within a simple shear dominated shearing involving top-to-the-SE thrust displacement. The formation of the thrust represents D1 deformation resulting in a nappe pile that was emplaced horizontally to form a classical fold-and-thrust belt characterized by flats and ramps (e.g. Burg et al. 1993). The thrusting identified in the Bakel area has the same tectonic transport direction as described elsewhere in the southern Mauritanides belt (e.g. Le Page 1978; Dia 1984). The comparison of the style of deformation between the domains points to a partitioning of the regional strain that resulted from strike-slip deformation in the Central Domain and a compressive thrust motion towards the Northern and Southern Domains. The structural evolution and architecture of the southern part of the Mauritanides belt is therefore typical of that of oblique
transpressional orogens as described elsewhere (e.g. Holdsworth & Strachan 1991; Vassallo & Wilson 2002; Jones et al. 2004). We propose a two-stage tectonic evolution of the Central Domain. Shear zones separate rocks with different tectonometamorphic histories during the early part of the Variscan collision: (1) northeastward ductile thrusting, dominant in the Northern and Southern Domains is shown by the general northwestward dip of the foliation; (2) sinistral wrenching, dominant in the Central Domain, is indicated by northeastward thrusting and sinistral oblique transpression. Thrusting and oblique transpression were combined during the tectonic evolution, as suggested by the spatial continuity between domains with steeply plunging lineation and zones with horizontal lineations. Similar structural assemblages have been briefly described in the Guidimakha region; for example, by Lahondie`re et al. (2004). Those workers described a N708E- to N1108E-oriented lineation that is oblique to the regional one.
Fig. 11. Three-dimensional sketch showing a schematic transpression zone. Partitioning of strain occurred into contraction-dominated and wrench-dominated domains.
TECTONIC EVOLUTION OF THE MAURITANIDES
495
Fig. 12. Three-dimensional diagram showing the main domains and associated folding patterns. The development of a flat-lying early fabric suggests an early period of thrusting. In the Central Domain, thrusting changed into sinistral transpression. This is explained by NW –SE shortening associated with strong foliation-parallel lithologies and accompanied by strong NE–SW shortening. (Redrawn after Holdsworth et al. 2002).
Conclusions The Mauritanides orogenic belts in SE Senegal expose regional- and local-scale structures that reveal the mechanism of accommodating 3D deformation. The results of our structural, metamorphic, finite-strain and kinematic analyses show that the superimposed structures in the Bakel area result from a progressive non-coaxial strain history. The Central Domain is the most highly metamorphosed and strongly deformed area. The tectonic transport direction changes from NE in the internal part of the orogen to NW in the marginal region. This change in displacement is a result of oblique convergence similar to that observed in other transpressional orogens (Little et al. 2002a; Jones et al. 2004). The exhumation of the ultramafic rocks may be related to NW–SE oblique convergence, which was accommodated by displacement partitioning and strain partitioning with the fault system. We interpret the gradual change in the orientation of lineations from NE in the internal part of the orogen to NW in the marginal domains, and folding in the central zone, as a result of a progressive change from a wrench-dominated to a more convergence-dominated collision. Recognition of a major sinistral oblique transpression in the Central
Domain of the investigated area provides new insights into the Hercynian tectonic evolution of the Mauritanides belt.
References A GASSIZ , J. F. 1970. Etude des ressources minie`res du Se´ne´gal oriental—Rapport sur la prospection du cuivre de la re´gion de Gabou (Ope´ration transitoire Septembre–De´cembre 1970). Rapport ine´dit, Projet Nation Unis de De´veloppement. B ASSOT , J. P. 1966. Etude ge´ologique du Se´ne´gal oriental et des ses confins guine´o-maliens. The`se Doctorat d’Etat, Universite´ Clermont Ferrand. B ASSOT , J. P., B ONHOMME , M., R OQUES , M. & V ACHETTE , M. 1963. Mesures d’aˆges absolus sur les se´ries pre´cambriennes et pale´ozoı¨ques du Se´ne´gal oriental. Bulletin Ge´ologique de France, 7, 401– 405. B ELL , T. H. 1981. Foliation development—the contribution, geometry and significance of progressive, bulk, inhomogeneous shortening. Tectonophysics, 75, 263– 296. B LANC , A., B ERNARD -G RIFFITHS , J., C ABY , R. ET AL . 1992. U –Pb dating and isotopic signature of the alkaline ring complexes of Bou Naga (Mauritania): its bearing on late Proterozoic plate tectonics around the West African Craton. Journal of African Earth Sciences, 14, 301– 311.
496
M. DABO ET AL.
B URG , J. P., C ORSINI , M., D IOP , C. & M AURIN , J. C. 1993. Structure et cine´matique du Sud de la chaıˆne des Mauritanides: un syste`me de nappe te´gumentaire varisque. Comptes Rendus de l’Acade´mie des Sciences, Se´rie I, 317, 697–703. C HIRON , J. C. 1973. Etude ge´ologique de la chaıˆne des Mauritanides entre le paralle`le de Moudje´ria et le fleuve Se´ne´gal (Mauritanie). The`se Doctorat e`s-Sciences, Universite´ de Lyon. C OBBOLD , P. R., G APAIS , D. & R OSSELLO , E. A. 1991. Partitioning of transpressive motions within a sigmoidal foldbelt: the Variscan Sierras Australes, Argentina. Journal of Structural Geology, 13, 743– 758. D ALLMEYER , R. D. & L ECORCHE´ , J. P. 1989. 40Ar– 39Ar polyorogenic mineral age 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. 40 Ar– 39Ar polyorogenic mineral age record within the southern Mauritanide orogen (Mbout– Bakel region) West Africa. American Journal of Science, 290, 1136– 1168. D ALLMEYER , R. D. & L ECORCHE´ , J. P. 1990b. 40Ar– 39Ar polyorogenic mineral age record within the northern Mauritanide orogen, West Africa. Tectonophysics, 177, 81–107. D EWEY , J. F., H OLDSWORTH , R. E. & S TRACHAN , R. A. 1998. Transpression and transtension zones. In: H OLDSWORTH , R. E., S TRACHAN , R. A. & D EWEY , J. F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 1– 14. D EYNOUX , M. 1978. Les formations glaciaires du Pre´cambrien terminal et de la fin de l’Ordovicien en Afrique de l’Ouest. Deux exemples de glaciation d’inlandis sur une plate forme stable. The`se Doctorat d’Etat, Universite´ de Marseille. D IA , O. 1984. La chaıˆne panafricaine et hercynienne des Mauritanides face au bassin Prote´rozoı¨que supe´rieur a` De´vonien de Taoude´ni dans le secteur-cle´ de Mejeria (Taganet, R. I. Mauritanie): lithostratigraphie et tectonique: un exemple de tectoniques tangentielles superpose´es. The`se d’Etat, Universite´ Aix– Marseille III. D IA , O., L ECORCHE , J. C. & L E P AGE , A. 1979. Trois e´ve´nements oroge´niques dans les Mauritanides d’Afrique occidentale. Revue de Ge´ographie Physique et de Ge´ologie Dynamique, 21, 403– 409. D IOP , C. 1996. Structure et circulation de fluides dans un avant pays schisteux: le syste`me de chevauchement des Mauritanides du Se´ne´gal. The`se Doctorat Ge´osciences, Institut National Polytechnique de Lorraine. D’L EMOS , R. S., B ROWN , M. & S TRACHAN , R. A. 1992. Granite magma generation, ascent and emplacement within a transpressional orogen. Journal of the Geological Society, London, 149, 487– 490. F LINN , D. 1962. On folding during three dimensional progressive deformation. Quarterly Journal of the Geological Society of London, 118, 385 –433. F RY , N. 1979. Random point distributions and strain measurement in rocks. Tectonophysics, 60, 89–105. G OSCOMBE , B. D., P ASSCHIER , C. & H AND , M. 2004. Boudinage classification: end member boudin types
and modified boudin structures. Journal of Structural Geology, 26, 739– 763. H ARLAND , W. B. 1971. Tectonic transpression in Caledonian Spitsbergen. AAPG Bulletin, 108, 27–42. H OLDSWORTH , R. E. & S TRACHAN , R. A. 1991. Interlinked system of ductile strike slip and thrusting formed by Caledonian sinistral transpression in northeastern Greenland. Geology, 19, 510–513. H OLDSWORTH , R. E., S TRACHAN , R. A. & D EWEY , J. F. (eds) 1998. Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135. H OLDSWORTH , R. E., T AVARNELLI , E., C LEGG , P., P INHEIRO , R. V. L., J ONES , R. R. & M C C AFFREY , K. J. W. 2002. Domainal deformation patterns and strain partitioning during transpression: an example from the Southern Uplands terrane, Scotland. Journal of the Geological Society, London, 159, 401–415. H UDLESTON , P. J., S CHULTZ -E LA , D. & S OUTHWICK , D. L. 1988. Transpression in an Archean greenstone belt, northern Minnesota. Canadian Journal of Earth Sciences, 25, 1060– 1068. J ONES , K. A. & S TRACHAN , R. A. 2000. Crustal thickening and ductile extension in the NE Greenland Caledonides: a metamorphic record from anatectic pelites. Journal of Metamorphic Geology, 18, 719– 735. J ONES , R. R. & T ANNER , P. W. G. 1995. Strain partitioning in transpression zones. Journal of Structural Geology, 17, 793– 802. J ONES , R. R., H OLDSWORTH , R. E., C LEGG , P., M C C AFFREY , K. & T AVARNELLI , E. 2004. Inclined transpression. Journal of Structural Geology, 26, 1531– 1548. L AHONDIE´ RE , D., C ABY , R., B OUAMATOU , M. A. & C ORSINI , M. 2004. Thin-skin tectonics and the latePaleozoic thrusting of oceanic lithosphere above the west African craton, southern Mauritania. Colloquium of African Geology. Orleans, France, Abstract 20e. L E P AGE , A. 1978. Les unite´s structurales de la re´gion de Bakel (Se´ne´gal oriental). Leur place dans la chaıˆne des Mauritanides. Comptes Rendus de l’Acade´mie des Sciences, Paris D, 286, 1853– 1856. L E P AGE , A. 1983. Les grandes unite´s des Mauritanides, aux confins du Se´ne´gal et de la Mauritanie. L’e´volution structurale de la chaıˆne du Pre´cambrien supe´rieur au De´vonien. The`se d’Etat, Universite´ Aix–Marseille III. L E P AGE , A. 1988. Rock deformation associated with the displacement of allochthonous units in the central segment of the Caledono-Hercynian Mauritanide belt (Islamic Republic of Mauritania and eastern Senegal). Journal of African Earth Sciences, 7, 265–283. L ECORCHE´ , J. P., D ALLMEYER , R. D. & V ILLENEUVE , M. 1989. Definition of tectonostratigraphic terranes in the Mauritanides, Bassaride, and Rokelide orogens, West Africa. In: D ALLMEYER , R. D. (ed.) Terranes in the Circum-Atlantic Paleozoic Orogens. Geological Society of America, Special Papers, 230, 131–144. L ECORCHE´ , J. P., B RONNER , G., D ALLMEYER , R. D., R OCCI , G. & R OUSSEL , J. 1991. The Mauritanides orogen and its northern extensions (Western Sahara and Zemmour), West Africa. In: D ALLMEYER , R. D. &
TECTONIC EVOLUTION OF THE MAURITANIDES L ECORCHE´ , J. P. (eds) The West African Orogens and Circum-Atlantic Correlations. Springer, Berlin, 187–227. L ISLE , R. J. 1977. Clastic grain shape and orientation in relation to cleavage from the Aberystwyth Grits, Wales. Tectonophysics, 39, 381–385. L ISTER , G. S. & S NOKE , A. W. 1984. S –C mylonites. Journal of Structural Geology, 6, 617– 638. L ISTER , G. S. & W ILLIAMS , P. F. 1983. The partition of deformation in flowing rock masses. Tectonophysics, 92, 1– 33. L ITTLE , T. A., H OLCOMBE , R. J. & I LG , B. R. 2002a. Ductile fabrics in the zone of active oblique convergence near the Alpine Fault, New Zealand: Identifying the neotectonic overprint. Journal of Structural Geology, 24, 193–217. L ITTLE , T. A., H OLCOMBE , R. J. & I LG , B. R. 2002b. Kinematics of oblique continental collision inferred from ductile microstructures and strain in midcrustal Alpine Schist, central South Island, New Zealand. Journal of Structural Geology, 24, 219– 239. O LDOW , I. S., B ALLY , A. W. & A VE´ L ALLEMANT , H. G. 1990. Transpression, orogenic float, and lithospheric balance. Geology, 18, 991–994. P ASSCHIER , C. W. 2001. Flanking stuctures. Journal of Structural Geology, 23, 951–962. P ASSCHIER , C. W. & S IMPSON , C. 1986. Porphyroclast systems as kinematic indicator. Journal of Structural Geology, 8, 831–844. P ETKOVIC , M. 1971. Etude des ressources minie`res du Se´ne´gal oriental. Rapport sur la recherche du cuivre a` Gabou, de´partement de Bakel. Rapport ine´dit, Projet Se´ne´gal, 17. P LATT , J. P. 1984. Secondary cleavage in ductile shear zones. Journal of Structural Geology, 6, 439 –442. R AMSAY , J. G. 1967. Folding and Fracturing of Rocks. McGraw-Hill, New York.
497
R AMSAY , J. G. & H UBER , M. I. 1983. The Techniques of Modern Structural Geology, 1: Strain Analysis. Academic Press, London. R EMY , P. 1985. Caracte`res tholeiitiques et alcalins dans le mate´riel ophiolitique du groupe d’El Aouidla (Mauritanides centrales). Ofioliti, 10, 391– 409. S ANDERSON , D. J. & M ARCHINI , W. R. D. 1984. Transpression. Journal of Structural Geology, 6, 449– 458. S IMPSON , C. 1984. Borrego Springs–Santa Rosa mylonite zone: a late Cretaceous west-directed thrust in southern California. Geology, 12, 8 –11. T APPONNIER , P., P ELTZER , G., L EDAIN , A. Y. & A RMIJO , R. 1982. Propagating extrusion tectonics in Asia: new insights from simple experiments with plasticine. Geology, 10, 611– 616. T IKOFF , B. & DE S AINT B LANQUATT , M. 1997. Transpressional shearing and strike-slip partitioning in the late Cretaceous Sierra Nevada magmatic arc, California. Tectonics, 16, 442–459. T IKOFF , B. & G REENE , D. 1997. Stretching lineations in transpressional shear zones: an example from the Sierra Nevada Batholith, California. Journal of Structural Geology, 19, 29–39. V ASSALLO , J. J. & W ILSON , C. J. L. 2002. Palaeoproterozoic regional-scale non-coaxial deformation: an example from eastern Eyre Peninsula, South Australia. Journal of Structural Geology, 24, 1– 24. V ILLENEUVE , M. 1993. The West African fold belts: structure and evolution. Comptes Rendus de l’ Acade´mie des Sciences, Se´rie II, 316, 411– 417. W OODCOCK , N. H. 1986. The role of the strike-slip fault systems at plate boundaries. Philosophical Transactions of the Royal Society of London, Series A, 317, 13– 29.
Deep structure of the southeastern margin of the West African craton from seismic reflection data, offshore Ghana KODJOPA ATTOH & LARRY BROWN Department of Earth and Atmospheric Sciences, Cornell University, Ithaca, NY 14853, USA (e-mail:
[email protected]) Abstract: Seismic reflection data, acquired as part of an oil exploration survey offshore Ghana, have been reprocessed to reveal the deep structure of the Proterozoic continental crust on the continental shelf, off West Africa. The seismic profiles presented are located across the Romanche transform margin, which transects the West African craton (WAC) margin where it is bounded by the Pan-African Dahomeyide orogen. The seismic data reveal highly reflective middle– lower crust of the WAC but a nearly transparent upper crust of c. 10 km thickness. Bundles of discontinuous reflections are located at the base of the unreflective upper crust and locally define the trough of a synformal structure interpreted as the keel of a Palaeoproterozoic (Birimian) greenstone belt exposed onshore. The deepest crust imaged is characterized by more continuous, subhorizontal reflections that are very similar to those observed in the lower crust of cratonic areas such as the Superior Province. Near the WAC margin the subhorizontal reflections are truncated by east-dipping reflections in the Pan-African domain, which are correlated with cratonward directed fold-and-thrust structures formed during the c. 600 Ma Pan-African orogeny. These seismic observations represent, to date, the first deep seismic reflection images of the West African craton margin and the continental lithosphere in the region.
Seismic reflection surveys have amply demonstrated the capability of the active source technique to reveal the deep structure of Precambrian orogens (Lucas et al. 1993; Lewry et al. 1994; Calvert et al. 1995). Among the significant results of deep seismic reflection experiments are evidence for highly reflective lower crust underlying cratons and well-defined Moho beneath shield areas, and occasional observations of strong reflectors in the continental mantle lithosphere that have been interpreted as relict subduction zones (Green et al. 1990; Calvert et al. 1995). To date, however, only few deep seismic reflection experiments have been carried out in Africa, although the network of orogenic belts outlining cratons and the cratons themselves are potentially fertile targets for systematic seismic exploration (e.g. De Wit & Tinker 2004). The West African craton (WAC) with its surrounding Pan-African orogens is a prime example of regions yet to be explored. In this paper, we report the results of the reprocessing and interpretation of marine seismic reflection data from an oil exploration survey offshore Ghana, which image the Proterozoic crust across the southeastern margin of the WAC. The data allow correlation of some of the deep reflectors with structures projected from the surface geology and provide, for the first time, seismic images of the deep structure of a critical segment of the WAC boundary. The seismic reflection data presented here also demonstrate the potential for using offshore marine seismic survey
data to begin addressing the problem of the paucity of deep seismic data to investigate the continental lithosphere in Africa.
Tectonic setting The Romanche transform margin offshore Ghana formed as a result of the Mesozoic break-up of Gondwana, and preserves the unmistakable scar of its transform origin (Basile et al. 1993; Mascle et al. 1996; Attoh et al. 2004). This margin is a favourable setting to use marine seismic reflection techniques to investigate the structure of the adjoining continental crust because, unlike rifted margins, a normal thickness of continental crust is predicted to be preserved here. The preservation of minimally extended continental lithosphere in such settings is the result of their formation by wrench shear (Scrutton 1982) and is evident in the satellite-derived bathymetric map (Fig. 1), which suggests that continental crust of significant thickness extends offshore until it abruptly abuts the oceanic crust. This transform margin setting is ideal to image the deep structure of the WAC boundary as well as the bounding Pan-African Dahomeyide orogen using marine seismic reflection data. Figure 2 displays the relation between the Romanche fracture zone (RFZ) and onshore tectonic elements; it shows that the Ghana transform continental margin, bounded by the RFZ, is
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 499–508. DOI: 10.1144/SP297.24 0305-8719/08/$15.00 # The Geological Society of London 2008.
500
K. ATTOH & L. BROWN
Fig. 1. Satellite-derived topographic –bathymetric map of the southeastern margin of the WAC showing the Palaeoproterozoic and Archaean domains, the bounding Pan-African orogen (Pan-African front; PAF) and the bathymetric expression of the Ghana margin, which suggests that the shelf is underlain by relatively thick continental crust. The Ghana margin is bounded by the Romanche fracture zone (RFZ) and the Deep Ivorian basin (DIB). In the land area, the highest regions (.1000 m) are light brown and the lowest regions are green, and in the ocean basin, areas below 2000 m are deep blue and those above 200 m are light blue, indicating the steep slope along the RFZ.
Fig. 2. Principal tectonic elements of onshore geology in relation to the deep seismic sections: PF, Pan-African Front; PS, Pan-African Suture (HP mafic suture zone unit shown by striped pattern); TS, Trans-Saharan Shear Zone; DIB, Deep Ivorian basin; DIBF, Deep Ivorian basin fault. Palaeoproterozoic greenstone belts are shown (W, Winneba; D, Dixcove). Inset shows the RFZ in relation to other Atlantic fracture zones.
DEEP STRUCTURE, OFFSHORE GHANA
an oblique transect of the boundary between the WAC underlain by Palaeoproterozoic (Birimian) rocks and the Pan-African Dahomeyide orogenic belt (Attoh & Ekwueme 1997). Along this transect, the principal tectonic units of the onshore geology include prominent structures of the Pan-African orogen that are favourably oriented to be imaged by offshore seismic reflection lines. The major tectonic elements of the Dahomeyide orogen include the Pan-African front (PF), which represents the western limit of deformation in the external zone, and the Pan-African suture (PS), represented by a ductile shear zone at the base of high-pressure mafic granulites that define the eastern edge of the WAC (Castaing et al. 1993; Attoh et al. 1997; Attoh & Morgan 2004). Offshore seismic reflection profiles reveal the projections of the Pan-African structures beneath the shelf strata (Attoh et al. 2004) and indicate that the projection of the RFZ coincides with the Trans-Saharan shear zone (TS), a dextral shear zone, rather than the PS or PF as was previously assumed (Edwards et al. 1997; Attoh et al. 2005). The southwestern limit of the Ghana shelf is defined by the Deep Ivorian basin (DIB), which is bounded by a normal fault, the Deep Ivorian basin fault (DIBF) (Fig. 2). This fault formed as part of the lithospheric-scale rift system, which was kinematically coupled to the Romanche transform during the opening of the Atlantic (Attoh et al. 2004). To the west of the Pan-African front (PF, Fig. 2) is the Palaeoproterozoic (Birimian) granite– greenstone terrane of the WAC. It is composed of NNE–SSW-striking greenstone belts (D and W, Fig. 2) as well as intervening metasedimentary belts intruded by granitoids. These lithologies underlie the southeastern terrane of the extensive c. 2.1 Ga juvenile crust (Attoh & Ekwueme 1997). The Birimian granite– greenstone terrane of the WAC displays many similarities to Late Archaean greenstone belt terranes such as those of the Superior Province of Canada (Sylvester & Attoh 1992, and references therein). Because granite – greenstone belts are significant components of many cratons, the deep structure of such terranes and the nature of their boundaries are important to understanding the origin of the cratonic lithosphere worldwide. Figure 3 is schematic geological cross-section of the southeastern margin of the WAC and its bounding Pan-African Dahomeyide orogen. It shows the current interpretation of the structures within the Dahomeyide orogen such as the westward verging folds represented by Ataccora nappes and panels of thrust-bounded units. Granitoid rocks to the east of the suture zone are inferred to represent terrane accreted to the WAC margin. The Winneba greenstone belt (W) is schematically
501
shown as a synformal structure west of the Pan-African front.
Seismic reflection profiles Seismic data and processing The multi-channel seismic reflection data used in this study were acquired for oil exploration by Geophysical Services Incorporated (GSI) in the 1980s. Digital tapes containing stacked seismic sections for selected lines, including the two lines discussed in this paper, were made available to us by Western Geophysical Incorporation. Although the survey was designed to image the sedimentary section of the Ghana shelf and slope, the record lengths are long enough to image basement structures on the lines near the coast. The two lines discussed here were selected because they are located close to the shoreline and contain significant record lengths beneath the Phanerozoic sedimentary section. Our efforts focused on post-stack reprocessing, carried out using ProMax (TM, Landmark Corporation) software. Most of the post-stack improvement in the sections shown were the result of coherency enhancement and time variant scaling (e.g. Kong et al. 1985), which revealed some of the deep structures better than migrated sections (Attoh et al. 2004).
Interpretation of seismic sections Attoh et al. (2004) used available chronostratigraphic data from oil exploration wells located across the Ghana shelf to calibrate the seismic stratigraphy and this led to the identification of three principal stratigraphic sequences representing pre-, syn- and post-rift strata. The pre-rift sequence consists largely of Palaeozoic strata ranging in age from Devonian to Carboniferous deposited on Palaeoproterozoic basement. The synrift sequence consists of Aptian to Albian siliciclastic strata, which have a distinct continental facies. These strata are inferred to have been deposited in pull-apart basins that formed during early stages of continental break-up along intracontinental strike-slip fault zones (Mascle et al. 1996). Seismic images of folding and faulting of the pre-rift and synrift strata were also presented and interpreted by Attoh et al. (2004) as the products of transpressional deformation related to transform tectonics. The seismic sections shown in Figure 4 are along line 6, which is located primarily in the offshore projection of the WAC. They reveal strong reflections in the Palaeoproterozoic continental crust beneath the reflectors of the overlying
502
K. ATTOH & L. BROWN
Fig. 3. Schematic geological cross-section of southeastern margin of the WAC and its bounding Pan-African Dahomeyide orogen.
sedimentary strata consisting of Palaeozoic (pre-rift) and Mezozoic (synrift) sequences. Near the centre of this profile (Fig. 4a), the sedimentary strata are offset by up to 3.5 s along a steep, eastdipping normal fault; on the downthrown side of this fault Devonian– Carboniferous strata show evidence of thickening from east to west and are interpreted as growth strata deposited during a Palaeozoic rifting event. The overall structure is interpreted as a half-graben. In the western, upthrown block of the half-graben, the stratified rocks thin rapidly (to less than 0.5 s) allowing access to a significant thickness (c. 6.0 s) of Palaeoproterozoic WAC basement. The Proterozoic crust underlying the downthrown block, and especially east of it (Fig. 4a), also reveals prominent reflectors that provide important information on the structure of the WAC. Figure 4b is the detailed seismic section of the western end of line 6; it displays a non-reflective upper crust down to a depth of 3 s. This nonreflective upper crust (c. 10 km thick) contrasts sharply with highly reflective crust below 3.5 s. Discontinuous reflection bundles (B1, B2 and B3) with typical vertical extents of c. 0.3 s occur at 3 s (5900 CDP (common depth point)) and 3.4 s (5500 CDP) and 3.6 s (5300 CDP). The reflection bundle at B1 apparently dips shallowly eastward and projects to the B2 reflections, suggesting that B1 –B3 reflections, with minor offsets, may represent reflectors that define the base of the upper crust, which is inferred to be composed of the metasediments and paragneisses exposed onshore. This non-reflective upper crust may represent Palaeoproterozoic intrusions. In contrast, the bundles of strong, subhorizontal reflections between 4 s and 7 s (B4– B5) are characteristically more continuous and define a crustal zone of strong reflections that is similar to middle and
lower crustal reflections just above the Moho in the Superior Province (Calvert et al. 2004) and other cratons. Assuming a velocity of 6–7 km s21, the maximum depth estimate for these reflections is, however, only 20– 25 km, which would correspond to depths in the middle crust rather than the lower continental crust. The longest of these deep reflections (B5) extends for c. 15 km (6000– 5400 CDP) between 6.5 and 7 s (Fig. 4b, at the bottom of section) and is subhorizontal. In Figure 4c, between 1700 and 1300 CDP, the ‘bow-tie effect’ characteristic of opposite-dipping bands of reflectors is evident at 4.5 s (C1). It indicates a synformal structure, which is here interpreted as the trough of the Winneba greenstone belt (W, Fig. 3). This synformal trough correlates with the eastern projection of the inferred base of supracrustal rocks to the west (B1 – B3, Fig. 4b) and represents the first ever seismic image of the deep structure of a Palaeoproterozoic greenstone belt in the WAC. Assuming typical continental crust velocities of 6 – 7 km s21, the estimated preserved thickness of greenstone belt supracrustal succession is c. 12 km, a thickness that is close to the stratigraphic thickness of typical greenstone belt sequences of the WAC (e.g. Attoh & Ekwueme 1997, tables 5.4.1 and 5.4.2), in contrast to the thickness of 4 – 5 km estimated from gravity anomaly models. For example, Attoh (1982) presented gravity models to show that the maximum thickness of the Palaeoproterozoic Nongodi greenstone belt is less than 5 km whereas the estimated stratigraphic thickness is greater than 9 km. Similarly, in the Archaean Sula greenstone belt of Sierra Leone the total stratigraphic thickness is about 8 km (see Attoh & Ekwueme 1997, and references therein), which is greater than typical estimates from gravity anomalies.
DEEP STRUCTURE, OFFSHORE GHANA
503
Fig. 4. Seismic profiles along line 6. TWTT, two-way travel time. (a) Reflective Palaeoproterozoic crust in the upthrown block and beneath the half-graben filled with thick Palaeozoic growth-strata. (b) Reflections at the western end of line 6, showing nearly transparent upper crust beneath thin Palaeozoic strata underlain by highly reflective middle crust (below 4.0 s). (c) ‘Bow-tie’ pattern of opposite-dipping reflections (C1) indicating a synformal structure correlated with the Palaeoproterozoic Winneba greenstone belt and prominent, east-dipping, deep reflections (C2) near the WAC margin.
504
K. ATTOH & L. BROWN
Fig. 4. (Continued).
The deep reflections near the eastern end of line 6 (Fig. 4c; C2) dip to the east from 5.5 s to 7 s. This feature is consistent with the apparent easterly dip of the mid-crustal reflections in the WAC (Fig. 4a), which may indicate overall tilt of the crust underlying the Ghana shelf as a result of uplift along the lithospheric-scale normal fault (DIBF, Fig. 2) located to the west of the seismic sections. This uplift was probably in response to the formation of the DIB by rifting, a deformation that was kinematically linked to wrench displacements along the Romanche transform during the opening of the equatorial Atlantic (Basile et al. 1993; Attoh et al. 2004). Thus the western segment of the Ghana margin represents an uplifted block formed by flexural rebound as a result of
displacement along the DIBF where relatively deep, east-tilting Proterozoic crust is preserved by the tectonic processes linked to transform margin formation. The strong, east-dipping, deep reflections near the WAC margin (e.g. C2 in Fig. 4c) may, however, also be relict of the thinning of the continental margin during Neoproterozoic rifting leading to the formation of the eastern margin of the WAC. This earlier rifting was responsible for the creation of the WAC margin, which was subsequently sutured along the Pan-African Dahomeyide orogen. Line 8 is located mainly in the Pan-African domain (Fig. 2) and, as such, provides the setting to image the deformed rocks of the WAC margin. Near the western end of line 8, reflections below
DEEP STRUCTURE, OFFSHORE GHANA
3.5 s (Fig. 5a) are interpreted as structures related to the Pan-African front (PF). The east-dipping deep reflections contrast with the coherent reflections of the Palaeozoic (pre-rift) and Mesozoic (synrift) strata overlying them. The deep reflections (D1– D2) between 4 and 7 s project to the structures with similar dip in the PF zone of the onshore surface geology (Attoh et al. 1997). These deep reflections are characterized by bundles of strong reflections (,0.3 s thick), separated by broader zones (up to 1.0 s thick), and are especially well developed from 7300 to 7700 CDP. The apparent flattening of these structures (E1 –E2, Fig. 5a) between 3.5 and 4.5 s (at 8000–7800 CDP) is also consistent with the nappe structures of the Pan-African Dahomeyide external zone. The apparent changes in the amplitude of deep reflections are ascribed to the variable intensity of deformations in the Pan-African zone, especially in the deformed margin of WAC represented onshore by protomylonitic granitoid gneisses. Figure 5b displays the reflections in the Pan-African domain near the central part of line 8; they consist of deep, east-dipping reflection bundles (F1 –F2) which contrast with the subhorizontal reflections of the overlying pre-rift Palaeozoic strata. The western end of the prominent deep reflections (F1) coincides with the projection of the Pan-African suture zone (PS) mapped on the surface whereas a broad zone to the west of it is relatively unreflective. This broad, unreflective crust extends to the zone of deep reflections near the west end of line 8 (Fig. 5a); it is here interpreted as consisting of weakly foliated granitoids, as evident in the variable intensity of deformation in the surface rocks. The coincidence of the strong reflections (F1) with the projection of PS (Fig. 5b) is considered significant, as it may represent the elusive deep seismic image of the Pan-African Dahomeyide suture. Assuming average crustal rock velocities of 6– 7 km s21 and noting that the thickness of the Phanerezoic sedimentary section is over 2 s, the estimated thickness of Palaeo- and Neoproterozoic crustal sections imaged in Figure 5 is ,12 km. Thus the section of the WAC margin imaged here probably represents upper crustal levels, especially considering that the orogenic belt is assumed to be underlain by over-thickened crust. In comparing the reflections along line 6 with those along line 8 (Fig. 5) the transition from the WAC domain to the Pan-African domain becomes evident. For example, the reflections in the WAC domain (Fig. 4) are largely subhorizontal to gently dipping and apparently represent deeper crustal levels, whereas those from the Pan-African belt are characterized by significant dip and apparently represent higher crustal levels.
505
Discussion The southeastern margin of the WAC provides the setting to address several critical problems in this region, using deep seismic reflection data. These problems include (1) determining the deep structure of the Pan-African domain, which preserves the record of West Gondwana assembly from the WAC and adjoining cratons (Amazonian and Sa˜o Francisco –Congo), and (2) evaluating the inference from seismic tomography of the WAC margin (Ritsema & van Heijst 2000), which showed that a shear-wave velocity anomaly corresponds to this boundary. The anomaly has been interpreted as evidence for a cool, thick, cratonic lithosphere (.200 km) beneath the WAC in contrast to the thinner, warm lithosphere beneath the surrounding Pan-African domain. Thus this region is clearly a key transect to investigate lithospheric aggregations during Gondwana assembly (using controlled source seismic imaging techniques). The data presented in this paper, although limited, confirm the potential of using deep seismic reflection data to map the cratonic boundary and contribute to the understanding of some of these problems. The contrasting deep reflections of the WAC crust compared with the Pan-African domain are similar to those of other Proterozoic continental margins (e.g. Lucas et al. 1993; Hall et al. 1995) but these data represent the first such observations in West Africa. Deeper seismic reflection data are needed across this and other segments of the WAC boundary to further document lithospheric interactions during Gondwana assembly. In the WAC itself, a significant result of the interpretation of the available data is the estimate of the thickness of Palaeoproterozoic greenstone belts that is consistent with those from stratigraphic measurements. In contrast, the estimated thickness of greenstone belts based on gravity anomaly models tends to be significantly low (e.g. Attoh 1982), and this has been a perplexing problem of greenstone belt geology. In the Pan-African Dahomeyide domain, the east-dipping, deep reflections beneath the weakly deformed, subhorizontal, Phanerozoic reflectors are those predicted from onshore geological observations. The dipping structures in the transition zone between the WAC and the Pan-African domain may represent relict structures of the thinning of WAC lithosphere during rifting but could also be the result of Pan-African deformation extending deep into the craton. The strong reflectivity of the WAC deep crust is typical of cratonic areas, suggesting a common origin of the reflections (Klemperer 1987; Green et al. 1990; Calvert et al. 1995). The subhorizontal attitude and the amplitude of reflectors in the deep crust are, however, unexpected from surface
506
K. ATTOH & L. BROWN
Fig. 5. Seismic profiles along line 8 showing deep reflections: (a) correlated with Pan-African structures onshore (D1– D2, E1– E2) along the deformed margin of the WAC; (b) (F1–F2) underlying the projection of the Pan-African suture zone (PS).
DEEP STRUCTURE, OFFSHORE GHANA
geological considerations or predictions of physical properties of the deep crust. For example, whereas the surface geology displayed on maps and evident in the field relations on which they are based presents complex dipping structures, the deep reflections, in contrast, suggest simple, nearhorizontal structures. In that sense, there appears to be a disconnection between the surface geology projections and deep reflections, but such complex structures, including the presence of subhorizontal sills, are beyond the resolution of the seismic reflection technique. The occurrence of strong reflections in the lower crust below the brittle –ductile transition is also unexpected, although it has been suggested that the acoustic discontinuities may be related to the presence of fluids in the deep crust (e.g. Ross et al. 2004). Connolly & Podladchikov (2004) have shown that, in appropriate tectonic settings, fluid flow in the deep crust can be restricted to a zone of neutral buoyancy that can preserve optimally oriented porosity domains to account for lower crustal reflectors. Their model predicts an inverted pressure gradient, which acts as a barrier to upward fluid flow, resulting in the stagnation of deep-crust generated fluids that can account for mid-crustal seismic reflectivity. This project was supported by a grant from Petroleum Research Fund of the American Chemical Society (PRF 34908-AC8). Thanks go to Western Geophysical Incorporation for providing the offshore Ghana seismic data and to the Ghana National Petroleum Corporation for information on and permission to access the data. A. Calvert, C. M. Doucoure´ and M. De Wit provided helpful reviews of the paper.
References A TTOH , K. 1982. Structure, gravity models and stratigraphy of an Early Proterozoic volcanic– sedimentary belt in northeastern Ghana. Precambrian Research, 18, 275–290. A TTOH , K. & E KWUEME , B. N. 1997. The West African shield. In: DE W IT , M. J. & A SHWAL , L. (eds) Greenstone Belts. Oxford University Press, Oxford, 517–528. 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 the lower crust. Journal of African Earth Sciences, 39, 1201– 1208. A TTOH , K., D ALLMEYER , R. D. & A FFATON , P. 1997. Chonology of nappe assembly in the Pan-African Dahomeyide orogen, West Africa: evidence from 40 Ar/39Ar mineral ages. Precambrian Research, 82, 153–171. A TTOH , K., B ROWN , L., G UO , J. & H AENLEIN , J. 2004. Seismic stratigraphic record of transpression and uplift on the Romanche transform margin, offshore Ghana. Tectonophysics, 378, 1 –16.
507
A TTOH , K., B ROWN , L. & H AENLEIN , J. 2005. The role of Pan-African structures in intraplate seismicity near the termination of the Romanche Fracture Zone, West Africa. Journal of African Earth Sciences, 43, 549– 555. B ASILE , C., M ASCLE , J., P OPOFF , M., B OULLIN , J. P. & M ASCLE , G. 1993. The Ivory Coast– Ghana transform margin: a marginal ridge structure deduced from seismic data. Tectonophysics, 222, 1 –19. C ALVERT , A. J., S AWYER , E. W., D AVIS , W. J. & L UDDEN , J. N. 1995. Archaean subduction inferred from seismic images of a mantle suture in the Superior Province. Nature, 375, 670–674. C ALVERT , A. J., C RUDEN , A. R. & H YNES , A. J. 2004. Seismic evidence for preservation of the Archaean Uchi granite– greenstone belt by crustal extension. Tectonophysics, 388, 135– 143. C ASTAING , C., T RIBOULET , C., F EYBESSE , J.-L. & C HEVREMENT , P. 1993. Tectonometamorphic evolution of Ghana, Togo, and Benin in the light of the PanAfrican/Brasiliano orogeny. Tectonophysics, 18, 323– 342. C ONNOLLY , J. A. D. & P ODLADCHIKOV , Y. Y. 2004. Fluid flow in compressive tectonic settings: Implications for midcrustal seismic reflectors and downward fluid migration. Journal of Geophyical Research, 109, B04201, doi:10.1029/2003JB002822. D E W IT , M. J. & T INKER , J. 2004. Crustal structures across the central Kaapvaal craton from deep-seismic reflection data. South African Journal of Geology, 107, 185–206. E DWARDS , R. A., W HITMARSH , R. B. & S CRUTTON , R. A. 1997. The crustal structure across the transform continental margin off Ghana, eastern equatorial Atlantic. Journal of Geophysical Research, 102, 747– 772. G REEN , A. G., M ILKEREIT , B., M AYRAND , L. J. ET AL . 1990. Deep structure of an Archaean greenstone terrane. Nature, 344, 327 –330. H ALL , J., W ARDIE , R. J., G OWER , C. F., K ERR , A., C OFLIN , K., K EEN , C. E. & C AROLL , P. 1995. Proterozoic orogens of the northeastern Canadian shield: new information from the Lithoprobe ECSOOT crustal reflection seismic survey. Canadian Journal of Earth Sciences, 32, 1119– 1131. K LEMPERER , S. L. 1987. Reflectivity of the crystalline crust—Hypothesis and tests. Geophysical Journal of the Royal Astronomical Society, 89, 217– 222. K ONG , S. M., P HINNEY , R. A. & C HOUDHURY , K. R. 1985. A nonlinear signal detector for enhancement of noisy seismic records. Geophysics, 50, 539–550. L EWRY , J. F., H AJNAL , Z., G REEN , A. ET AL . 1994. Structure of a Palaeoproterozoic continent –continent collision zone: a Lithoprobe seismic reflection profile across the Trans-Hudson orogen, Canada. Tectonophysics, 232, 143–160. L UCAS , S. B., G REEN , A., H AJNAL , Z. ET AL . 1993. Deep seismic profile across a Proterozoic collision zone: surprises at depth. Nature, 363, 339–342. M ASCLE , J., L OHMANN , G. P., C LIFT , P. D. ET AL . (eds) 1996. Proceedings of the Ocean Drilling Program, Initial Report, 159. Ocean Drilling Program, College Station, TX.
508
K. ATTOH & L. BROWN
R ITSEMA , J. & VAN H EIJST , H. 2000. New seismic model of the upper mantle beneath Africa. Geology, 28, 63–66. R OSS , A. R., B ROWN , L. D., P ANANONT , P. ET AL . 2004. Deep reflection surveying in central Tibet: Lower crustal layering and crustal flow. Geophysical Journal International, 156, 115– 128. S CRUTTON , R. A. 1982. Passive continental margins: a review of observations and mechanisms. In:
S CRUTTON , R. A. (ed.) Dynamics of Passive Margins. American Geophysical Union, Geodynamics Series, 6, 5 –15. 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 Archaean– Proterozoic boundary. Journal of Geology, 100, 377–393.
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) JULIEN BERGER1,2,3, NASSER ENNIH4, JEAN-PAUL LIE´GEOIS1, COLLIN NKONO2, JEAN-CLAUDE C. MERCIER3 & DANIEL DEMAIFFE2 1
Section de Ge´ologie Isotopique, Muse´e Royal de l’Afrique Centrale, 3080 Tervuren, Belgium (e-mail:
[email protected]) 2
Laboratoire de Ge´ochimie Isotopique et Ge´odynamique Chimique, CP160/02, Universite´ Libre de Bruxelles (ULB), 1050 Brussels, Belgium
3
Centre Littoral de Ge´ophysique, UMR-CNRS 6250 “LIENSs”, Universite´ de La Rochelle, 17402 La Rochelle Cedex-1, France
4
Laboratoire de Ge´odynamique, Universite´ d’El Jadida, BP20, 24000 El Jadida, Morocco Abstract: We used quantitative textural measurement, electron microprobe microanalysis and thermobarometry on clinopyroxene from a Cenozoic pyroxene-nephelinite flow located along the northern boundary of the West African craton to decipher magma differentiation processes in underlying magma chambers. The crystal size distributions of clinopyroxene phenocrysts show straight but also curved and kinked patterns and the clinopyroxene show large compositional variations in a single flow (Mg-number 48–88). These observations are strong evidence for magma mixing between a nephelinite magma and a more differentiated phonolitic melt at depth. Detailed thermobarometry on these clinopyroxene shows that at least three magma chambers are present below Saghro and that they are emplaced at the main physical interface within the lithosphere: (1) at the crust– mantle boundary, where the mantle-derived nephelinite has been mixed with a pre-existing phonolitic magma chamber; (2) at the lower–upper crust boundary; (3) close to the surface in a sub-volcanic magma chamber. Some high-pressure phenocrysts (up to 14 kbar) have also probably crystallized within the upper lithospheric mantle. The high clinopyroxene proportion in samples from the base of the flow is thought to reflect crystal settling during cooling of the nephelinite flow at the surface.
Lava flows at the surface are generally considered as simple, rather homogeneous geochemical systems. Nevertheless, complex crystallization histories before eruption can sometimes be recognized by the presence of complex zoning of phenocrysts, implying complex chemical evolution of the magma (fractional crystallization, assimilation, local disequilibrium) and/or changes in the physical conditions (pressure and temperature) during magma rise and solidification (Duda & Schmincke 1985; Dobosi 1989; Bachmann & Dungan 2002). Phenocryst resorption is also strong evidence of solid – liquid disequilibrium during magmatic evolution (O’Brien et al. 1988; Streck et al. 2002). Understanding the crystallization conditions of a magma before its eruption has many petrological consequences: the P – T conditions have a direct effect on the composition of the fractionating phases and thus on the geochemical evolution
trend of the volcanic series (Grove & Baker 1984; Scoates et al. 2006). The physical conditions that prevailed during the crystallization steps of the magma development are thus needed to model the magmatic evolution. In addition, the pressure of crystallization of the phenocrysts is directly linked to the depth of the magma chamber. A detailed thermobarometric approach to the phenocryst assemblage could then be used to infer the depth of the magma chambers beneath active or extinct volcanoes and, by extension, the depth of the main physical interfaces within the lithosphere (crust – mantle boundary, elastic – plastic intracrustal boundary, etc.). We show in this paper that, by a combination of quantitative textural (crystal size distribution; CSD) and chemical (electron microprobe analysis; EMPA) studies on clinopyroxene phenocrysts coupled with thermobarometric investigations on a
From: ENNIH , N. & LIE´ GEOIS , J.-P. (eds) The Boundaries of the West African Craton. Geological Society, London, Special Publications, 297, 509–524. DOI: 10.1144/SP297.25 0305-8719/08/$15.00 # The Geological Society of London 2008.
510
J. BERGER ET AL.
single Cenozoic nephelinite flow from the Djbel Saghro field (Morocco), the complete and complex evolution of the magma can be deciphered. We demonstrate that polybaric differentiation combined with magma mixing has a strong imprint on the compositional zoning of phenocrysts, and on the texture of the nephelinite rock itself. Also, latestage cooling of the magma at the surface has an effect on the distribution of phenocrysts within the flow.
Geological setting The Saghro volcanic field is located at the eastern edge of the Saghro inlier in the Anti-Atlas of Morocco (Fig. 1). The volcanic activity occurred in two periods during the Cenozoic (Berrahma et al. 1993): the Late Miocene phase (9.6– 7.6 Ma) was concentrated in the south and the Late Pliocene (c. 2.9 Ma) in the north. The Saghro volcanic field is as large (c. 1500 km2) as the contemporaneous huge Sirwa stratovolcano (25 km in diameter) located further to the west but in the Saghro inlier. All Cenozoic volcanic exposures (lava flows, necks and minor pyroclastic deposits) are of small extent. The chemistry of the erupted lavas is typically bimodal, consisting of silica-undersaturated nephelinites and phonolites. A single exposure of intermediate lava with phono-tephrite composition is
known (Ibhi 2000, and our unpublished data). The Saghro suite of lavas is peralkaline in overall composition and is associated with a carbonatitic magmatism expressed as carbonatite xenoliths in the nephelinites (Ibhi et al. 2002). The geodynamical event responsible for the onset of this alkaline magmatism is still debated. Considering that the Saghro volcanic field, the Sirwa stratovolcano and the Canary Islands are located on the same structural line, some workers (Hoernle & Schmincke 1993) have related this volcanism to a mantle plume rising from the base of the upper mantle. In contrast, others (Lie´geois et al. 2005), taking into account the fact that the Cenozoic alkaline magmatism of Morocco, and of West Africa as a whole, occurred along Pan-African or Variscan structures reactivated during the Alpine orogeny, have related the volcanism to the paroxysmal events of the Africa –Europe convergence. The Saghro and Sirwa volcanic fields are located along the northern boundary of the West African craton, just to the south of the current High Atlas mountain range. We focused our investigations on a 12 m thick flow of pyroxene-nephelinite from Foum el Kous (Fig. 1). This lava flow rests upon a contemporary flow of olivine-nephelinite itself covering continental Cretaceous sediments. The eruption of the pyroxene-nephelinite flow has been dated at c. 2.9 Ma (whole-rock K –Ar, Berrahma et al. 1993).
Fig. 1. (a) Location of the Saghro Cenozoic volcanic field within Morocco and (b) a simplified geological map of the eastern edge of the Saghro inlier (modified after Ibhi et al. 2002).
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM
Sample description and whole-rock composition Under the microscope the nephelinite lava flow consists of phenocrysts of Ti-augite, olivine and minor Ti-magnetite in a fine-grained groundmass dominated by clinopyroxene, nepheline and Ti-magnetite, with accessory apatite, biotite and perovskite. Olivine phenocrysts (,1 mm) are generally resorbed with corrosion gulfs and partial iddingsitization of their borders. The phenocrysts are unzoned with the notable exception of clinopyroxene, which displays complex zoning patterns. Four types of clinopyroxene have been recognized on a textural basis. Type 1 clinopyroxenes are zoned euhedral crystals (.700 mm) with dark greenish brown cores surrounded by light brown mantles and thin dark brown rims (Fig. 2a and b). The central dark core is always anhedral to subhedral and exhibits complex zonations. It contains mineral (biotite, amphibole) and abundant fluid inclusions. Groundmass minerals (nepheline, Ti-magnetite and apatite) are present as inclusions at the contact zone between the dark core and the light mantle zone. The latter is unzoned, but in some samples it evolves gradually towards dark rim, which also contains numerous inclusions of the groundmass minerals. Type 2 clinopyroxenes are small (,500 mm), euhedral, complexly zoned crystals with green cores (Fig. 2c and d). The complete succession of zones is: in the centre, an anhedral green core with numerous inclusions of the groundmass minerals, rimmed by an anhedral dark green part, itself surrounded by a light brown mantle, and finally a thin rim of brown clinopyroxene with numerous inclusions of nepheline, Ti-magnetite and apatite. Type 3 clinopyroxenes are small (,500 mm) light brown phenocrysts with tabular shapes (Fig. 2e). They rarely exhibit sector zoning but have a distinct core and rim. In contrast to the core, which is free of inclusions, the dark brown augitic rim has numerous inclusions of the matrix minerals. The border of the crystals sometimes shows resorption features marked by corrosion gulfs filled with matrix minerals (clinopyroxene, nepheline, Ti-magnetite). Type 4 clinopyroxenes (50 and 200 mm) are groundmass phases with granular or acicular habits. They are in textural equilibrium with the other matrix minerals and sometimes contain inclusions of Ti-magnetite, apatite, nepheline and perovskite. One of the most striking features of this lava flow is the presence of numerous megacrysts. Amphibole and olivine have been observed (Ibhi 2000) but clinopyroxene is by far the most abundant. These clinopyroxene megacrysts are euhedral
511
to subhedral, but occasionally only the relicts of larger broken crystals are preserved (Fig. 2f). They are fractured and rich in fluid inclusions but they are not zoned, nor do they have mineral inclusions, except for a thin rim of darker clinopyroxene containing inclusions of the groundmass minerals. Xenoliths are abundant; they are generally rounded and do not exceed 5 cm. They are mainly monomineralic clinopyroxenite with accessory biotite, apatite and olivine, but subordinate mantle peridotites and carbonatites have been found (Ibhi et al. 2002). The xenoliths generally show equilibrated medium-grained granular texture. The clinopyroxene is zoned and has a dark core and a light brown rim. The clinopyroxene from the xenoliths contains numerous small vitreous inclusions that probably formed during the strong temperature increase undergone by the pyroxenites during their ascent in the host lava. The whole-rock composition of the flow is rather homogeneous (Fig. 3). It falls in the foidite field of the total alkalis–silica (TAS) diagram (Le Maitre 2002), and the presence of modal and normative (14 wt%) nepheline together with the dominance of clinopyroxene as phenocrysts define this rock as a pyroxene-nephelinite. The total alkali content is high (4–5 wt%) for a relatively low SiO2 content (39 wt%). The high Mg-number (63–64) of the nephelinite confirms its primitive character compared with the whole series of the Saghro volcanic field (Ibhi et al. 2002, and our unpublished data).
Quantitative textural analysis: crystal size distribution of clinopyroxene Methods The crystal size distribution (CSD) analysis was made on thin-section photomicrographs obtained with an optical microscope. The crystals were outlined using computer-aided design software and the drawing was exported as a black and white bitmap file. The size of the phenocrysts was measured with the ‘SPO’ software (Launeau 2004; Launeau & Robin 2005). With this method, each grain of the digitalized thin section is positioned, outlined and represented by an ellipse and a box (inertia tensor method). The measure used for the CSD analysis is the length of the ellipse major axis for each grain (200 –500 grains per sample; the smallest measured grains are in the range 0.15 –0.2 mm). The CSD 2D size data from the binary image were converted into volume data by applying stereological corrections based on the Schwartz –Saltikov algorithm (De Hoff & Rhines
512
J. BERGER ET AL.
Fig. 2. Photomicrographs of clinopyroxenes. (a, b) Type 1: Ti-augite with brown Al-rich core: (a) transmitted light; (b) backscattered electron image. (c, d) Type 2: Ti-augite phenocryst with Fe-rich green core (c) transmitted light; (d) backscattered electron image. (e) Type 3: Ti-augite in plane-polarized light. (f) Transmitted light photomicrograph of a broken megacryst. Ne, nepheline; Ap, apatite; Ttm, titanomagnetite; Prv, perovskite.
1972). This procedure (Hammer et al. 1999) does not cause artefacts such as the pronounced reduction of the scatter of data observed for the method proposed by Peterson (1996). Stereological corrections based on the Schartz–Saltikov algorithm, as adopted by Armienti et al. (1994) and Higgins (2000), have been carefully tested against
real and theoretical cases; they represent a satisfactory solution in most cases of petrological investigation. The CSD parameters used in this work were obtained using the software ‘CSDcorrection1.36’ (Higgins 2000, 2002). The data were plotted, following general convention (Marsh 1988), as linear crystal
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM 18 Phonolite
16
Total alkalis
14 12 Tephriphonolite
10
Trachyte Phonotephrite Basaltic Trachydacite Tephrite trachyBasanite Trachy- andesite basalt
Foidite
8 6 4 2 0 35
40
45
SiO2
50
55
60
Fig. 3. TAS diagrams (fields from Le Maitre 2002) showing the composition of two analysed samples from the pyroxene-nephelinite flow. Grey fields indicate the composition of the whole Saghro lava suite (Ibhi 2000, and our unpublished data).
size v. population density (ln(n) (cm24) v. L (cm)); a log10 based scale with 4– 8 bins per decade was used. Ten samples from the core and the lower border of the flow were selected for this purpose. The main parameters of the CSD analysis on these samples are shown in Table 1.
Results Most of the plotted CSDs for the clinopyroxenes are straight or slightly curved but, in detail, three types of distribution can be recognized: (1) six samples have straight to slightly concave-upward curved CSD patterns (Fig. 4a); (2) three samples have a more pronounced curvature (Fig. 4b): the CSD is straight for the small crystals (,0.1 cm, corrected grain size) and shows a pronounced concaveupward pattern for larger grains (.0.1 cm); (3) one sample has a ‘kinked’ CSD pattern (Fig. 4c)
513
characterized by a steep slope for smaller grains (,0.1 cm) and a more gentle slope for larger sizes (.0.1 cm), with a gap between the two populations. The curved CSDs are best interpreted as mixing of two populations of crystals with contrasted crystallization histories. Higgins (1996) and Peterson (1996) have shown that mixing of two populations of crystals, each characterized by contrasting slopes in CSD patterns, can produce a slightly curved CSD. The kinked CSD pattern presented in Figure 4c can be divided into two parts: a slightly concave-upward curve similar to the curved CSD presented above and a population of large crystals identified as broken megacrysts. The first concave population, with a steep negative slope in the CSD diagram, is rich in small grains. It corresponds to the tabular euhedral unzoned small phenocrysts (type 3) mixed with small amounts of type 1 and 2 phenocrysts. The regressed CSD line of the slightly concave-upward curve has a slope between 243 and 296 cm21 and an intercept between 12.5 and 16.2 cm24. The population of large clinopyroxenes, with a more gentle negative slope has a lower intercept, corresponds to a few large clinopyroxenes interpreted as megacrysts that show the same composition as type 1 phenocrysts. The CSD analysis of the few large crystals with kinked CSD patterns gives a slope around 230 cm21. The regressed slope and intercept of all CSDs are negatively correlated (Table 1). In a classic CSD diagram the regressed lines define a fan (not shown here; Resmini 1993; Zieg & Marsh 2002). However, following Higgins (2002), this correlation is not necessarily meaningful, because of the closure limit effect, and a diagram of characteristic length v. clinopyroxene modal proportion would give more information on geological processes. CSD regression parameters calculated on straight CSDs and small (usually ,0.1 cm) clinopyroxene populations of curved and kinked
Table 1. Results of CSD analysis on clinopyroxenes from the 10 analysed samples from the pyroxene-nephelinite flow of Foum el Kous Sample Fn3 Fs1 Fs6 Fs10 Fs14 Fs15 Fs17 Fs29 Fs41 Fs43
Type
Intercept (cm24)
1s
Slope (cm21)
1s
Cpx vol% measured
Cpx vol% regressed
Curved Straight Straight Kinked Straight Curved Straight Straight Curved Straight
14.42 14.42 12.40 13.48 15.17 16.27 14.40 13.22 15.81 16.20
0.17 0.18 0.22 0.20 0.14 0.24 0.18 0.20 0.15 0.17
258.6 265.5 243.3 248.9 267.0 292.3 263.5 247.5 286.1 295.7
2.9 3.6 3.2 3.3 3.0 6.6 3.9 3.1 4.0 5.2
18.0 14.9 11.4 17.0 18.1 17.8 18.3 16.1 21.1 21.0
24.9 20.5 14.2 10.0 27.1 21.2 22.8 21.5 29.2 26.5
514
J. BERGER ET AL.
a
14
Ln population density (cm–4)
c
b
Straight CSDs
Curved CSDs
Kinked CSD (Fs10)
12
(7)
10
8
6
4
2
Fs1
Fs14
Fs29
Fs6
Fs17
Fs43
(7) (5)
(5)
(6)
(7)
Fn3
Fs15
(8)
(7)
Fs41 (6)
(7)
0 0
0.05
0.1
0.15
0.2
0.25
0.3
0
0.05
0.1
Size (cm)
0.15
0.2
Size (cm)
0.25
0.3
0
0.05
0.1
0.15
0.2
0.25
0.3
0.35
Size (cm)
Fig. 4. The three types of CSD patterns observed for the clinopyroxenes of the pyroxene-nephelinite lava flow from Foum el Kous.
CSDs are plotted in Figure 5. The characteristic length of augite phenocrysts is negatively correlated with the modal proportion. The samples from the base of the flow have the highest clinopyroxene content (17–22 vol%) but generally low characteristic length (,0.013 cm).
Mineral chemistry: focus on clinopyroxene The data have been acquired on a Cameca SX100 electron microprobe at the CAMPARIS section of the Paris VI University. The operating conditions were 10 nA for the beam current, 15 kV for the accelerating voltage and 10 s counting time per
Fig. 5. Characteristic length v. clinopyroxene content (vol%) for straight CSDs and the left, straight, part of curved and kinked CSDs.
element. Corrections of the raw data were made using the PAP method of Pouchou & Pichoir (1984). Typical accuracy is close to 1% for oxides with concentration .1 wt% and around 10 wt% for elements with contents ,1%. Unzoned olivine phenocrysts have Mg-numbers between 82 and 87 for the entire flow. Groundmass nepheline shows a restricted compositional range of Ne68 – 73Ks26 – 31Q0 – 3 and the Ti-magnetite has 11 –16 wt% of TiO2 with rather high MnO and MgO contents (1–2 and 3– 6 wt%, respectively). Minor amounts of Nb– Zr–REE-rich perovskite and fluorapatite are present in the groundmass but they have not been quantitatively analysed.
The clinopyroxene chemistry The chemistry of the Saghro clinopyroxenes (Table 2; Fig. 6) is directly linked to their aspect and texture. The dark brown cores (type 1) are Tiand Al-rich augite that can be distinguished by their high Al and Ti content, high VIAl/IVAl ratio and rather low Mg-number (65–82). This compositional range is the same as those of the megacrysts and the clinopyroxene in pyroxenite xenoliths. Their high Al (0.3–0.5 p.f.u.) content and VIAl/IVAl ratio point to a high pressure of crystallization (Aoki & Shiba 1973). These pyroxenes also have low Cr2O3 contents (,0.33 wt% Cr2O3), similar to high-pressure pyroxenes from alkaline lavas (Schulze 1987). The small unzoned euhedral phenocrysts (type 3) and the light brown mantle surrounding green and brown cores have the same composition with low Al and Ti contents but with a higher mean Mg-number (range 75–83) than the brown Ti–Al augite cores (type 1); their VIAl/IVAl ratio is low (,0.3), indicating a lower crystallization pressure. The compositional differences between the brown cores and their
Table 2. Representative analyses of clinopyroxene phases from the Foum el Kous pyroxene-nephelinite flow Fs10 Type 1 Core
Fs17 Type 1 Core
Fs29 Type 1 Core
Fn3 Type 2 Core
Fs5 Type 2 Core
Fs10 Type 2 Core
Fs5 Type 3 Core
Fs10 Type 3 Core
Fs17 Type 3 Core
Fs5 Type 4 Core
Fs10 Type 4 Core
Fs29 Type 4 Core
wt% SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O Sum
46.11 2.55 8.68 0.00 8.12 0.10 11.29 22.74 0.53 100.13
44.26 2.89 10.91 0.06 7.93 0.03 10.53 22.86 0.47 99.96
43.74 3.35 10.59 0.01 7.88 0.13 10.66 22.65 0.47 99.49
39.45 3.98 14.95 0.01 12.26 0.11 6.29 22.50 0.97 100.52
42.82 3.36 11.58 0.00 10.11 0.18 9.27 22.61 0.93 100.86
43.80 2.96 9.27 0.00 8.92 0.12 10.58 24.22 0.48 100.35
47.28 2.09 7.74 0.09 5.40 0.07 13.59 22.87 0.63 99.75
44.84 2.93 8.59 0.37 6.17 0.14 12.30 23.04 0.85 99.24
47.67 2.33 6.02 0.12 6.26 0.07 13.76 23.92 0.56 100.71
50.91 2.06 2.64 0.00 5.75 0.19 14.77 24.14 0.50 100.96
49.99 2.19 3.36 0.00 5.77 0.17 14.26 23.98 0.55 100.28
47.20 3.17 4.83 0.00 6.00 0.16 13.36 23.54 0.79 99.05
1.74 0.26 0.08 0.11 0.06 0.00 0.75 0.06 0.00 0.90 0.04 0.82
1.67 0.33 0.05 0.17 0.08 0.01 0.68 0.02 0.00 0.92 0.06 0.78
1.75 0.25 0.01 0.15 0.06 0.00 0.75 0.04 0.00 0.94 0.04 0.80
1.86 0.11 0.00 0.06 0.06 0.00 0.80 0.09 0.01 0.94 0.04 0.84
1.84 0.15 0.00 0.08 0.06 0.00 0.78 0.08 0.01 0.95 0.04 0.83
1.76 0.21 0.00 0.12 0.09 0.00 0.74 0.04 0.00 0.94 0.06 0.82
Structural formulae based on 6 oxygens, Fe 3þ calculated by local charge balance Si 1.72 1.65 1.64 1.49 1.59 AlIV 0.28 0.35 0.36 0.51 0.41 0.10 0.13 0.11 0.16 0.10 AlVI Fe3þ 0.08 0.08 0.09 0.19 0.18 Ti 0.07 0.08 0.09 0.11 0.09 Cr 0.00 0.00 0.00 0.00 0.00 Mg 0.63 0.59 0.60 0.35 0.51 Fe2þ 0.17 0.16 0.16 0.19 0.13 Mn 0.00 0.00 0.00 0.00 0.01 Ca 0.91 0.91 0.91 0.91 0.90 Na 0.04 0.03 0.03 0.07 0.07 Mg-no. 0.71 0.70 0.71 0.48 0.62
1.63 0.37 0.04 0.19 0.08 0.00 0.59 0.09 0.00 0.97 0.03 0.68
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM
Sample: Type: Core/rim:
(Continued)
515
516
Table 2. Continued Sample: Type: Core/rim:
Fs10 Light mantle Rim
Fs17 Light mantle Rim
Fs10 Rim Rim
Fs17 Rim Rim
Fs29 Rim Rim
Fs17 Megacryst Core
Fs17 Megacryst Core
Fs17 Megacryst Core
Fs3 Xenolith Core
Fs3 Xenolith Core
Fs3 Xenolith Core
48.11 2.10 5.98 0.15 6.04 0.08 14.01 23.32 0.53 100.33
48.44 1.82 5.40 0.00 6.82 0.11 13.27 23.19 0.61 99.67
46.97 2.58 5.82 0.00 6.15 0.18 13.59 24.34 0.38 100.02
53.66 0.89 0.46 0.02 4.36 0.14 15.83 24.30 0.56 100.24
50.12 2.21 3.47 0.08 5.55 0.12 14.38 24.41 0.43 100.77
50.41 2.09 2.73 0.01 5.30 0.15 14.61 24.36 0.59 100.25
47.27 2.11 8.01 0.39 4.43 0.01 13.37 23.11 0.57 99.27
43.53 3.41 9.10 0.05 7.14 0.12 11.71 23.41 0.52 98.99
48.27 1.75 7.04 0.41 4.27 0.09 14.12 23.01 0.50 99.48
44.96 2.65 9.92 0.05 6.71 0.04 11.72 22.47 0.74 99.27
46.71 1.78 8.08 0.20 6.05 0.04 12.89 23.16 0.51 99.42
48.61 1.85 6.34 0.22 5.42 0.08 13.77 22.82 0.61 99.74
1.85 0.12 0.00 0.07 0.06 0.00 0.80 0.06 0.00 0.96 0.04 0.86
1.75 0.25 0.10 0.06 0.06 0.01 0.74 0.07 0.00 0.92 0.04 0.84
1.64 0.36 0.04 0.17 0.10 0.00 0.66 0.05 0.00 0.94 0.04 0.75
1.78 0.22 0.08 0.06 0.05 0.01 0.78 0.07 0.00 0.91 0.04 0.85
1.68 0.32 0.11 0.12 0.07 0.00 0.65 0.09 0.00 0.90 0.05 0.76
1.73 0.27 0.09 0.11 0.05 0.01 0.71 0.08 0.00 0.92 0.04 0.79
1.79 0.21 0.07 0.07 0.05 0.01 0.76 0.09 0.00 0.90 0.04 0.82
Structural formulae based on 6 oxygens, Fe 3þ calculated by local charge Si 1.77 1.80 1.74 1.96 AlIV 0.23 0.20 0.25 0.02 AlVI 0.03 0.03 0.00 0.00 Fe3þ 0.12 0.11 0.15 0.03 Ti 0.06 0.05 0.07 0.02 Cr 0.00 0.00 0.00 0.00 Mg 0.77 0.73 0.75 0.86 0.06 0.10 0.03 0.09 Fe2þ Mn 0.00 0.00 0.01 0.00 Ca 0.92 0.92 0.96 0.95 Na 0.04 0.04 0.03 0.04 Mg-no. 0.81 0.78 0.81 0.88
balance 1.84 0.15 0.00 0.07 0.06 0.00 0.79 0.09 0.00 0.96 0.03 0.83
J. BERGER ET AL.
wt% SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O Sum
Fn3 Light mantle Rim
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM
517 0.70
0.18 0.16
0.60
0.14 0.50
VI
0.10
0.40
0.08
0.30
Al
Al
0.12
0.06
0.20
0.04 0.10
0.02 0.00 0.00
0.10
0.20
0.30
0.40
0.50
0.60
0.70
IV
Mg#
Al
0.16
0.16
0.14
0.14
0.12
0.12
0.10
0.10
0.08
0.08
0.06
0.06
0.04
0.04
0.02
0.02
0.00 0.00
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.00 0.45 0.50 0.55 0.60 0.65 0.70 0.75 0.80 0.85 0.90
Mg#
Al
0.18
0.18 0.16 0.14 0.12
Ti
Ti
0.00 0.45 0.50 0.55 0.60 0.65 0.70 0.75 0.80 0.85 0.90
Type 4 Type 1 Dark-brown core Groundmass Type 2 Small phenocrysts Type 3 Common to all types Phenocryst rim Green core Light Mantle Megacrysts
0.16 0.14 0.12
0.10
0.10
0.08
0.08
0.06
0.06
0.04
0.04
0.02
0.02
0.00 0.00
0.10
0.20
0.30
0.40
0.50
0.60
0.70
Al
Na
Na
Xenoliths
0.00 0.45 0.50 0.55 0.60 0.65 0.70 0.75 0.80 0.85 0.90
Mg#
Fig. 6. Composition of the clinopyroxenes from the pyroxene-nephelinite flow of Foum el Kous.
surrounding light mantles for type 1 clinopyroxene are illustrated by a microprobe traverse (Fig. 7a): there is a drastic increase of Mg content and a concomitant drop in Al and Fe contents from the core to the mantle zone. Na and Ti contents do not vary significantly in the two parts of the phenocrysts. The zoned clinopyroxene with green core (type 3) has a slightly higher Fe content than the other pyroxenes; it corresponds to an Al–Fe augite. The
Mg-number is low (45 –65) whereas the Na, Al and Ti contents are high (.0.3, 0.03 –0.1 and 0.06 –0.12 p.f.u., respectively). Here also, the high VI Al/IVAl ratio is indicative of a high pressure of crystallization. A close examination of the microprobe traverse (Fig. 7b) shows that four zones are present: a green core with 0.3–0.4 p.f.u. Fe, a green external core characterized by much higher Fe and Al contents (both up to 0.5 p.f.u.), a light
518
J. BERGER ET AL.
(a)
Dark-brown core
Light mantle 0.80
0.80
0.70
0.70 0.60
Ti Fe Mg
0.40
Na
0.30
0.40 0.30 0.20
0.10
0.10
0
50
100
350
400
0.00 450
0.80
Light mantle 0.70
0.50 0.40 0.30
0.60 0.50 0.40 0.30
0.20
0.20
0.10
0.10
0.00
p.f.u.
Thin external rim
0.60
150 200 250 300 Distance from core (µm)
Extgreen core
Internal green core
0.80 0.70
p.f.u.
0.50
0.20
0.00
(b)
0.60
p.f.u.
p.f.u.
0.50
Thin external rim
Al
0.00 0
50
100 150 200 Distance from core (µm)
250
300
Fig. 7. Microprobe traverses within the clinopyroxene phenocrysts with brown core (type 1: a) and green core (type 2: b).
large brown mantle with lower Al but higher Mg contents, and a thin (,20 mm) rim with extremely low Al but high Mg. The composition of the groundmass clinopyroxene overlaps those of the extreme rims of the phenocrysts. They have the lowest Al, Ti and VIAl contents but the highest Mg-number (78–87).
Thermobarometry Two methods have been used here to quantitatively estimate the crystallization pressure. The first method is the structural barometer of Nimis & Ulmer (1998); this has the advantage of being composition independent and avoids the problems of cpx–melt disequilibrium. Exact
knowledge of the equilibrium melt composition is not required but the type of melt has to be known (anhydrous v. hydrous, tholeiitic v. alkaline, etc.). The presence of hydrated minerals (amphibole and biotite) in the megacrysts and in the xenoliths together with the 2.5 wt% value of the whole-rock loss on ignition (LOI) argue for a significant water content of the parental magma. In consequence, the barometer calibrated for hydrous basaltic compositions has been used (the anhydrous calibration was tested but the results were significantly too low, ,5 kbar). The second method is the barometer based on clinopyroxene– liquid equilibrium (Putirka et al. 2003); it was not applied to phenocrysts with Mg-number ,70 because of evident disequilibrium
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM
with the nephelinite melt composition. In this barometer, the relation between the composition and the pressure was established by minimizing the Gibbs free energy at equilibrium, assuming an ideal behaviour of the components. This barometer covers a wide compositional range, from silica-undersaturated to silicic melts, including alkali-rich lavas. However, it requires an independent estimation of the temperature. For that purpose, the pyroxene solvus thermometer of Lindsley (1983) was used. As the content of nonquadrilateral components (Al, Ti, Na) is high in the Saghro clinopyroxenes, the error on the temperature calculation is .50 8C (Lindsley 1983), approaching 100 8C (the error increase is about 5 8C per mol% of non-quadrilateral components). For the pyroxene megacrysts, the pyroxenes of the xenoliths and the dark brown cores of phenocrysts (type 1), the data cluster between 1000 and 1200 8C, with most values between 1100 and 1200 8C. This result is consistent with those of previous studies on clinopyroxenes from alkaline lavas (Bondi et al. 2002; Damasceno et al. 2002). The light mantle overgrowths and the small phenocrysts (types 2 and 3) have crystallization temperatures in the same range, or slightly lower, around 1100 8C. The groundmass pyroxenes (type 4) and the rims of phenocrysts have crystallized at significantly lower temperatures, below 1000 8C, with a mean value close to 950 8C. Concerning pressure, to handle the large number of values calculated and taking into account the large errors inherent in the geobarometers used (+2.6 kbar, according to Nimis & Ulmer 1998) the results have been treated statistically. Histograms of crystallization pressures are presented in Figure 8 for the two barometers. Pressure estimates from the structural barometer for the dark brown cores of type 1 clinopyroxene, green cores of type 2, megacrysts and xenoliths range between 7 and 14 kbar, with a peak between 9 and 10 kbar (mean 9.9 +1.6 kbar; the presented error is the standard deviation on the population of measured pressures). The light mantle surrounding the core of phenocrysts and the small phenocrysts (type 3) have crystallization pressures between 3 and 9 kbar, with a peak between 5 and 6 kbar (mean 5.9 +1.9 kbar). The cpx –liquid barometer (Putirka et al. 2003) gives a slightly lower mean value (mean 4.6 + 1.5 kbar), but this is similar within error limits to type 3 clinopyroxenes, although the distribution appears to be bimodal with one cluster between 3 and 4 kbar and another between 6 and 7 kbar. Groundmass clinopyroxene and the phenocryst rims (type 4) have significantly lower crystallization pressure (0–4 kbar) with a mean of 1+1.1 kbar. In detail, the crystallization pressure
519
Fig. 8. Histograms of pressure estimates for the different types of clinopyroxene from the nephelinite flow.
of the type 4 pyroxenes also seems to be bimodal, with one cluster around 0 kbar (surface conditions) and another between 2 and 3 kbar. The calibration of Putirka et al. (2003) either gives low values (,2 kbar) for groundmass clinopyroxene (type 4) or is unable to calculate a pressure. This is due to the very low content in the jadeite end-member (close or equal to zero) in the groundmass pyroxenes.
Discussion Polybaric differentiation and the level of magma chambers The high-Al augite (cores of type 1 and 2) of the nephelinite flow is characterized by high VIAl/IVAl ratio and has generally the same composition as the pyroxenes from the xenoliths and the megacrysts. Pressure estimates for these clinopyroxenes are around 10 kbar, but a higher pressure has been obtained for one sample (up to 14 kbar). The geophysical modelling of the Anti-Atlas lithosphere (Teixell et al. 2005) has revealed a crustal thickness of 36 km. Considering a density of 2750 kg m23 for the upper crust and 2930 kg m23 for the lower crust and respective thicknesses of 23 and 13 km (Teixell
520
J. BERGER ET AL.
et al. 2005), the calculated pressure is 6.1 kbar at the base of the upper crust and 9.8 kbar at the crust– mantle boundary. The high-Al augite thus represents fractionation products of the nephelinite magma in a magma chamber located close to the Moho. Some megacrysts may even have crystallized at higher pressure (c. 14 kbar) in the mantle at a depth of about 50 km. Considering that the base of the lithospheric mantle is at c. 60 –70 km (Teixell et al. 2005), this depth could correspond to the depth at which the lithospheric mantle becomes ductile; that is, at the transition between the mechanical boundary layer and the thermal boundary layer (MBL–TBL transition). Crystallization pressure of the small phenocrysts (type 3) and of the light mantle surrounding high-Al cores are around 6 kbar, a pressure that closely corresponds to the depth of the upper– lower crust boundary. We can then ascribe a second magma chamber to this physical boundary (Fig. 9). The low-Al augites from the groundmass and the thin rims of the phenocrysts have crystallized at very low pressure (0–1 kbar), in a subvolcanic magma chamber or at the surface, after eruption of the nephelinite flow. The P estimates summarized in Figure 8 seem to suggest a continuous range of pressure from 14 to 0 kbar for clinopyroxene crystallization instead of the three discrete
stages as discussed above. However, a close examination of the phenocryst zonations shows that the contacts between the zones are sharp from both the chemical and optical points of view. Pressure computed from the composition along a microprobe traverse in a type 1 phenocryst also shows that the core has crystallized at high pressure (10 kbar) and the mantle at about 5–6 kbar with no intermediate value. The apparent continuous pressure range is thus most probably an effect of the uncertainties inherent in the structural barometer (+2.6– 3.0 kbar; Nimis & Ulmer 1998) and in analytical errors. We can thus infer the presence of magma chambers at three or maybe four structural levels beneath the Djbel Saghro volcanic system: just below the surface, at the upper –lower crust boundary, at the crust –mantle boundary, and possibly within the lithospheric mantle, at the MBL –TBL boundary. However, whether there are several magma chambers or only one at each structural level for the Saghro volcanic field as a whole (1500 km2) has not yet been assessed. Evidence for multi-chamber magmatic differentiation is common in alkaline volcanic series (Woodland & Jugo 2007). Beneath the Sirwa volcano (Bondi et al. 2002), the high-Al diopside is thought to have crystallized in a deep magma chamber, at the crust –mantle boundary. In the
Fig. 9. Schematic sketch representing the location of the magma chambers in the lithospheric cross-section, the textural evolution of the clinopyroxene phenocrysts during their ascent to the surface and the related CSD patterns.
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM
Kerguelen archipelago, a detailed study of clinopyroxenes has revealed three magma chambers emplaced at the levels of the physical boundaries in the crust (Damasceno et al. 2002). This is also the case for the Saghro lavas. The interaction between the Mg-rich mantlederived nephelinite and a phonolitic melt at the base of the crust probably enhanced the rapid ascent of the mafic melt towards the magma chamber situated at the Conrad discontinuity (Woodland & Jugo 2007). The mixing between two magmas of contrasted compositions presumably led to the uptake of the differentiated Al- and Fe-rich phenocrysts and megacrysts from the phonolitic magma chamber in the nephelinite magma. In areas for which geophysical data are not available, this barometric approach on clinopyroxenes could be a powerful tool to estimate the depth of the most important interfaces within the lithosphere, in active or recently extinct volcanic fields.
CSD analysis: evidence for magma mixing and crystal settling The CSD of clinopyroxene phenocrysts shows three types of patterns. Samples with straight CSDs are devoid of high pressure (HP; c. 10 kbar) crystals. Curved and kinked CSDs are a mixture between two populations of clinopyroxene: (1) types 1 and 2 pyroxenes with Fe- and Al-rich cores that crystallized at around 10 kbar and (2) type 3 Mg-rich brown augite forming either light mantles around high-pressure cores or unzoned phenocrysts (or megacrysts) that crystallized at medium pressure (MP; c. 0.6 kbar). The population of MP phenocrysts (population A, Fig. 9) has a steep slope in CSD diagrams whereas pyroxenes with HP cores (population B) have CSDs with more gentle slopes. In CSD diagrams (Figs 4 and 9), the transition between the two populations corresponds to a size around 0.1 cm. The difference between the curved and the kinked CSDs probably results from the different percentages of the two populations in a given sample and from the contrast of size between the two CSD populations. In fact, the transition between the two CSD populations is progressive; there is a continuous range of grain sizes. Following Higgins (1996) and Peterson (1996), a curved CSD would correspond to a mixture of population A grains (type 1 pyroxene) with only a small amount (1–5 vol%) of population B grains (type 1 and 2 phenocrysts). The kinked CSD represents samples in which only a few very large crystals (broken megacrysts, up to 0.4 cm) with HP cores are present in a matrix of smaller MP phenocrysts. The gap between the two parts
521
of the kinked CSD represents the lack of crystals with intermediate sizes (0.1 –0.25 cm). Similar conclusions have already been drawn on textural analysis of megacryst-bearing granitoids (Higgins 1999). Even if only three large crystals are present, they account for a significant part of the CSD because they occupy a large modal volume (up to 7 vol%), forming their own CSD with a gentle slope. The CSDs in our samples are the witness of an important process of crystal mixing of two populations with contrasted composition and pressure of crystallization (Fig. 9). The steeper CSD (measured on clinopyroxene populations ,0.1 cm) of the three samples collected at the base of the flow cannot be explained as the result of faster cooling rates at the base of the flow (Zieg & Marsh 2002) because clinopyroxene phenocrysts were already present in the magma before eruption (the small phenocrysts have crystallized at around 6 kbar). Moreover, the Bagnold effect cannot account for the steeper slope and high modal proportions observed in the lowermost samples, because this will displace most of the crystals toward the core of the flow, where magma velocities are higher (Nkono et al. 2006, and references therein). As a result, the base of the flow will be crystal poor and this does not fit with our observations. We propose that the higher clinopyroxene modal proportions towards the base of the flow are the result of crystal settling during cooling of the nephelinite flow at the surface; such observations have already been made at the base of thick andesitic flows (Higgins 2002).
Evidence for mixing between phonolite and nephelinite A characteristic of the studied nephelinite flow is the wide compositional range of its pyroxenes. The high variability of clinopyroxene Mg-number (48 –88) suggests that most of them are in disequilibrium with the host nephelinite liquid. Fe–Mg exchange coefficients between clinopyroxene and liquid (Kd) vary from 0.1 to 0.49 in most crystallization experiments (0.01 –110 kbar; 800–1800 8C in basic to silicic melts; Putirka et al. 2003). An average Kd value of 0.27 + 0.06 is in agreement with most experiments (Putirka et al. 2003; see references given by Scoates et al. 2006). The pyroxenes in equilibrium with the Saghro nephelinite liquids (Mg-number 62 –70) should have an Mg-number between 80 and 92. This implies that the pyroxenes with Mg-number , 80 have crystallized from more evolved Fe-rich melts. Only the groundmass pyroxene, the rim of the phenocrysts and the small euhedral unzoned phenocrysts
522
J. BERGER ET AL.
(including the light mantle overgrowths surrounding the HP cores) seem to be in equilibrium with the host rock. The green augite forming the core of some clinopyroxene phenocrysts (types 1 and 2) is also too Fe-rich to be in equilibrium with the nephelinite. As no other peralkaline magmatic activity is documented in the area, the low Mg but Fe- and Al-rich green augite cannot be a xenocryst from the surrounding rocks and must be contemporaneous with the nephelinite magmatism. Its low Mg-number suggests that it was probably in equilibrium with a phonolitic liquid (Fig. 10). The brown high-Al Ti-augite forming the core of some phenocrysts as well as pyroxene megacrysts and those from pyroxenite xenoliths have a composition intermediate between the augite from the phonolite and that from the nephelinite. We can conclude that these phenocrysts have crystallized from a melt that probably formed by mixing of a phonolite and a nephelinite. Indeed, only one exposure of intermediate magma has been found in the Saghro area and field evidence (presence of mafic volcanic enclaves in a more felsic matrix) suggests that this lava was formed through the mechanical mixing between two different magmas. Such a model is reasonable, as magma mixing is also thought to play an important role in the genesis of the nephelinite–phonolite bimodal associations (Duda & Schmincke 1985; Gwalani et al. 2000). The nephelinite magma has been mixed with a more evolved melt of phonolitic composition at the
Fig. 10. Histogram showing the distribution of the clinopyroxene Mg-number in the pyroxene-nephelinite flow. The ranges labelled phonolites and nephelinites are the theoretical ranges of composition of clinopyroxenes in equilibrium with phonolites and nephelinites from the Saghro, considering a range for Fe– Mg exchange coefficient between 0.21 and 0.33. Legend as in Figure 8.
level of the Moho discontinuity. There is thus a need for a magma chamber of phonolitic composition at the Moho. The first clinopyroxene in equilibrium with the nephelinite (type 3 pyroxene represented by population A in CSDs) crystallized only at 6 kbar, at the upper –lower crust interface. The interaction between the nephelinite magma and the phonolitic magma probably produced rapid ascent of the nephelinite melt towards the lower –upper crust boundary, where it began to crystallize MP, Mg-rich clinopyroxene phenocrysts. Magma mixing is indeed one of the proposed factors to explain rapid ascent of magma from a deeper level; it thus allows the uptake of xenocrysts, megacrysts and xenoliths to the surface (Woodland & Jugo 2007).
Conclusion The crystal size distribution (CSD) of clinopyroxene phenocrysts in a single nephelinite flow of the Saghro volcanic field (Morocco) suggests that at least two populations of augite phenocrysts coexist. The curved and kinked CSD patterns can be modelled as the mixture of two straight CSDs: a population of small phenocrysts of brown clinopyroxene (,0.1 cm) and a population of large augites (.0.1 cm) characterized by dark brown to green cores. The thermobarometric investigations show that the two populations have not crystallized at the same depth. The first population comes from a magma chamber at the upper –lower crust boundary (6 kbar) and the second population corresponds to the relicts of the high-pressure fractionation of the magma, at the Moho discontinuity (10 kbar). In addition, a few crystals give a still higher pressure of 14 kbar, suggesting crystallization within the lithospheric mantle, at the MBL –TBL boundary. Groundmass clinopyroxene probably crystallized during flow emplacement but the existence of a subvolcanic chamber cannot be excluded. The various clinopyroxenes observed in a single nephelinite flow indicate the presence of four magma chambers, each located at rheological discontinuities of the local lithosphere: (1) at the brittle –ductile transition within the lithospheric mantle; (2) at the mantle –crust transition; (3) at the brittle–ductile transition within the crust; (4) at the lithosphere –atmosphere transition. The pressure determined for these four structural levels (respectively 14, 10, 6 kbar and c. 1 kbar) is in agreement with geophysical modelling. In addition, these magma chambers are close to a vertical lithospheric structure corresponding to the transition between the northern boundary of the West African craton (Ennih & Lie´geois 2001) and the Jurassic palaeorift of the High Atlas mountains.
COMPLEX MULTI-CHAMBER MAGMATIC SYSTEM
The wide range of clinopyroxene compositions in a single nephelinite flow (Mg-number 48 –88) and the mixing of the two crystal populations suggested by the CSD analysis demonstrate that magma mixing between a phonolitic melt and a nephelinitic melt plays an important role in the genesis of the Saghro bimodal peralkaline lava suite. The detailed approach that we use in this study is potentially applicable to many active volcanoes. It could be used to monitor active volcanic fields by investigating the depth of the subsurface and deep-seated magma reservoirs. J.B. and N.E. thank the local Moroccan people from the Djbel Saghro area for their hospitality and help in the field. J.B. and D.D. also acknowledge the ULB ‘Fonds Cambier’ for financing field trips to Morocco. The helpful comments of M. D. Higgins and R. G. Resmini improved the quality of the paper.
References A OKI , K. & S HIBA , I. 1973. Pyroxenes from lherzolite inclusion of Itinomegata, Japan. Lithos, 6, 41–51. A RMIENTI , P., P ARESCHI , M. T., I NNOCENTI , F. & P OMPILIO , M. 1994. Effects of magma storage and ascent on the kinetics of crystal-growth—the case of the 1991– 93 Mt Etna eruption. Contributions to Mineralogy and Petrology, 115, 402–414. B ACHMANN , O. & D UNGAN , M. A. 2002. Temperatureinduced Al-zoning in hornblendes of the Fish Canyon magma, Colorado. American Mineralogist, 87, 1062– 1076. B ERRAHMA , M., D ELALOYE , M., F AURE -M URET , A. & R ACHDI , H. 1993. Premie`res donne´es ge´ochronologiques sur le volcanisme alcalin du Jbel Saghro, AntiAtlas, Maroc. Journal of African Earth Sciences, 17, 333–341. B ONDI , M., M ORTEN , L., N IMIS , P., R OSSI , P. L. & T RANNE , C. A. 2002. Megacrysts and mafic– ultramafic xenolith-bearing ignimbrites from Sirwa Volcano, Morocco: phase petrology and thermobarometry. Mineralogy and Petrology, 75, 203–221. D AMASCENO , D., S COATES , J. S., W EIS , D., F REY , F. A. & G IRET , A. 2002. Mineral chemistry of mildly alkalic basalts from the 25 Ma Mont Crozier section, Kerguelen Archipelago; constraints on phenocryst crystallization environments. Journal of Petrology, 43, 1389– 1413. D OBOSI , G. 1989. Clinopyroxene zoning patterns in the young alkali basalts of Hungary and their petrogenetic significance. Contributions to Mineralogy and Petrology, 101, 112– 121. D UDA , A. & S CHMINCKE , H.-U. 1985. Polybaric differentiation of alkali basaltic magmas: evidence from green-core clinopyroxenes (Eifel, FRG). Contributions to Mineralogy and Petrology, 91, 340–353. 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.
523
G ROVE , T. L. & B AKER , M. B. 1984. Phase equilibrium controls on the tholeiitic versus calc-alkaline differentiation trends. Journal of Geophysical Research, 89, 3253– 3274. G WALANI , L. G., R OCK , N. M. S., R AMASAMY , R., G RIFFIN , B. J. & M ULAI , B. P. 2000. Complexly zoned Ti-rich melanite– schorlomite garnets from Ambadungar carbonatite –alkalic complex, Deccan Igneous Province, Gujarat State, Western India. Journal of Asian Earth Sciences, 18, 163– 176. H AMMER , J. E., C ASHMAN , K. V., H OBLITT , R. P. & N EWMAN , S. 1999. Degassing and microlite crystallization during pre-climactic events of the 1991 eruption of Mt. Pinatubo, Philippines. Bulletin of Volcanology, 60, 355– 380. H IGGINS , M. D. 1996. Magma dynamics beneath Kameni volcano, Thera, Greece, as revealed by crystal size and shape measurements. Journal of Volcanology and Geothermal Research, 70, 37–48. H IGGINS , M. D. 1999. Origin of megacrysts in granitoids by textural coarsening: a crystal size distribution (CSD) study of microcline in the Cathedral Peak granodiorite, Sierra Nevada, California. In: F ERNANDEZ , C. & C ASTRO , A. (eds) Understanding Granites: Integrating Modern and Classical Techniques. Geological Society, London, Special Publications, 158, 207–219. H IGGINS , M. D. 2000. Measurement of crystal size distributions. American Mineralogist, 85, 1105–1116. H IGGINS , M. D. 2002. Closure in crystal size distributions (CSD), verification of CSD calculations, and the significance of CSD fans. American Mineralogist, 87, 171– 175. H OERNLE , K. & S CHMINCKE , H. U. 1993. The role of partial melting in the 15-Ma geochemical evolution of Gran Canaria—a blob model for the Canary hotspot. Journal of Petrology, 34, 599– 626. I BHI , A. 2000. Le volcanisme Plio-Quaternaire de Saghro (Anti-Atlas, Maroc) et les enclaves basiques et ultrabasiques associe´es. PhD thesis, University of Agadir. I BHI , A., N ACHIT , H., A BIA , E. H. & H ERNANDEZ , J. 2002. Intervention of carbonate components in the petrogenesis of the pyroxene nephelinites from the Jbel Saghro (Anti-Atlas, Morocco). Bulletin de la Socie´te´ Ge´ologique de France, 173, 37– 43. L AUNEAU , P. 2004. Mise en e´vidence des e´coulements magmatiques par analyse d’images 2-D des distributions 3-D d’orientations pre´fe´rentielles de formes. Bulletin de la Socie´te´ Ge´ologique de France, 175, 331– 350. L AUNEAU , P. & R OBIN , P.-Y. F. 2005. Determination of fabric and strain ellipsoids from measured sectional ellipses—implementation and applications. Journal of Structural Geology, 27, 2223–2233. L E M AITRE , R. W. 2002. Igneous Rocks. A Classification and Glossary of Terms, 2nd edn. Cambridge University Press, Cambridge. 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 Papers, 388, 379–400.
524
J. BERGER ET AL.
L INDSLEY , D. H. 1983. Pyroxene thermometry. American Mineralogist, 68, 477– 493. M ARSH , B. D. 1988. Crystal Size Distribution (CSD) in rocks and the kinetics and dynamics of crystallization. 1. Theory. Contributions to Mineralogy and Petrology, 99, 277– 291. N IMIS , P. & U LMER , P. 1998. Clinopyroxene geobarometry of magmatic rocks Part 1: An expanded structural geobarometer for anhydrous and hydrous, basic and ultrabasic systems. Contributions to Mineralogy and Petrology, 133, 122– 135. N KONO , C., F EMENIAS , O., D IOT , H., B ERZA , T. & D EMAIFFE , D. 2006. Flowage differentiation in an andesitic dyke of the Motru Dyke Swarm (Southern Carpathians, Romania) inferred from AMS, CSD and geochemistry. Journal of Volcanology and Geothermal Research, 154, 201– 221. O’B RIEN , H. E., I RVING , A. J. & M C C ALLUM , I. S. 1988. Complex zoning and resorption of phenocrysts in mixed potassic mafic magmas of the Highwood Mountains, Montana. American Mineralogist, 73, 1007–1024. P ETERSON , T. D. 1996. A refined technique for measuring crystal size distributions in thin section. Contributions to Mineralogy and Petrology, 124, 395–405. P OUCHOU , J. L. & P ICHOIR , F. 1984. Un nouveau mode`le de calcul pour la microanalyse quantitative par spectrome´trie de rayons X—Partie I: application a` l’analyse d’e´chantillons homoge`nes. Recherche Ae´rospatiale, 3, 167–192. P UTIRKA , K. D., M IKAELIAN , H., R YERSON , F. & S HAW , H. 2003. New clinopyroxene–liquid thermobarometers for mafic, evolved, and volatile-bearing
lava compositions, with applications to lavas from Tibet and the Snake River Plain, Idaho. American Mineralogist, 88, 1542– 1554. R ESMINI , R. G. 1993. Dynamics of Magma Within the Crust: A Study Using Crystal Size Distributions. PhD dissertation, Johns Hopkins University, Baltimore, MD. S CHULZE , D. J. 1987. Megacrysts from alkalic volcanic rocks. In: N IXON , P. H. (ed.) Mantle Xenoliths. Wiley, New York, 433–451. S COATES , J. S., L O C ASCIO , M., W EIS , D. & L INDSLEY , D. H. 2006. Experimental constraints on the origin and evolution of mildly alkalic basalts from the Kerguelen Archipelago, Southeast Indian Ocean. Contributions to Mineralogy and Petrology, 151, 582– 599. S TRECK , M. J., D UNGAN , M. A., M ALAVASSI , E., R EAGAN , M. K. & B USSY , F. 2002. The role of basalt replenishment in the generation of basaltic andesites of the ongoing activity at Arenal volcano, Costa Rica: evidence from clinopyroxene and spinel. Bulletin of Volcanology, 64, 316– 327. T EIXELL , A., A YARZA , P., Z EYEN , H., F ERNANDEZ , M. & A RBOLEYA , M. L. 2005. Effects of mantle upwelling in a compressional setting: the Atlas Mountains of Morocco. Terra Nova, 17, 456– 461. W OODLAND , A. B. & J UGO , P. J. 2007. A complex magmatic system beneath the Deve`s volcanic field, Massif Central, France: evidence from clinopyroxene megacrysts. Contributions to Mineralogy and Petrology, 153, 719– 731. Z IEG , M. J. & M ARSH , B. D. 2002. Crystal size distributions and scaling laws in the quantification of igneous textures. Journal of Petrology, 43, 85–101.
Index Note: Page numbers denoted in italics refer to figures and those in bold refer to tables. accretionary complex 236, 240, 241– 242, 241 Adoudou Formation 285– 302, 329, 330–331, 333, 341, 436 bimodal volcanism 285 –302 carbon isotopes 294–295 chemostratigraphic control 294– 295 oxygen isotopes 294 –295 aeromagnetic anomalies 84– 86, 85 Anna structure 84 and Riedel fracture model 86 tectonic control 85– 86 Affela N’ouassif pluton 265, 266–267, 268–269, 270, 271, 274, 279 Affole´ high 173 Aftout granitoids 80 Afyon Basement Complex 411–412, 412 Cadomian arc related magmatism 414–415 felsic magmatism 409–431 fractional melting and crystallization model 421– 422, 423, 424, 425, 427 geodynamics 425– 427 schists 411, 413 TTG source 422, 423, 424, 425, 427 see also Afyon granites Afyon granites 409, 411, 413, 415, 420, 427 crystallization 418– 421 geochemistry 409, 416–418, 417, 418, 420, 424–425, 425 magmatic differentiation 416– 418 magmatic temperatures 418– 421 source 409, 418– 421, 422, 425–427 Agualilet unit 181 Ahmeyim greenstone Belt 40 Aı¨t Daoui granite 4, 9, 11–12, 13, 14 Ait Nebdas pluton 266, 268, 271, 274, 278 Ait Saoun area 299, 333 Calcaires infe´rieures 289–293, 290– 291, 292, 295 Ait Saoun area volcanics Basal Series 296– 297, 296 geochemistry 295– 299, 297, 298 Harker diagram 297, 297 ignimbritic rhyolite 296, 297 provenance 297 –298, 300 texture 296, 297 Ait-Atman, Na-amphiboles 235 Akjoujt area 54, 56, 57, 72 deformation events 54–55 Fe-Mg carbonate 68–71, 70, 71, 73 Fe-Mg clinoamphibole-chlorite schist 60–64, 64, 64, 65, 66– 67, 72, 73 geological setting 53–75, 55 meta-carbonate 60, 64, 64, 65, 65, 66– 67, 68–71, 69, 72 metamorphism 54–56, 56, 74 structural features 54– 56, 56
Akjoujt meta-basalt 54–57, 56, 58– 59, 68, 72, 73 amphibolite 58, 63, 73 biotite-actinolite schist 59, 63, 72 chlorite schist 59, 63 hydrothermal overprint 71, 71 source 59, 68, 71 tectonic discrimination 63 trace element data 53, 58–60, 61– 62 Algeria 147– 167 Alougoum volcanic complex 285, 288, 332, 441 Alpine orogeny 8, 303, 435 Alpine Tauride-Anatolide Platform 409 alumina saturation index 10, 12 Amalaoulaou massif 203, 205, 211– 212 garnet rich pyriclasites 211 granulitic metamorphism 211 LT-HP overprint 211–212 metagabbros 203, 205, 211– 212 Pan-African overprint 211–212 Amazonia, cratonic fragments 349 Amlouggi tonalites 267 amphibole chemical composition 237, 240 thermobarometric calculations 206 Anfeg batholith 113, 144 Anietir unit 182 Anna structure garnet analysis 99–101, 103, 104 kimberlite indicator minerals (KIM) 99 Ansongo eclogites 205, 212, 213, 215 Anti-Atlas 1– 17, 234–235, 234, 266, 333, 433–452, 453 autochthonous 433, 444–445, 448 basement uplift 448, 462 basement–cover relationships 445– 447 Cambrian rocks 330, 332–337, 336, 337, 338, 339 –340, 435 Cenozoic volcanism 506 Devonian extension 453 –465 Devonian fault pattern 458, 459–461, 460, 461 Devonian geology 455, 467 –482, 468 Eburian basement 4– 14 folds 438–439, 441–443 geological setting 5, 250, 434, 435– 437, 454, 455 gold mineralization 249–264 Hercynian facies 467, 476 lateral variations 467 me´lange complex 235 Meseta correlations 461 Mesozoic– Cenozoic cover 437 metacratonic 1–17, 447, 448 Neoproterozoic rocks 332, 334–337, 334–335, 336, 337, 338 Palaeozoic cover 435, 436– 437, 454 Palaeozoic palaeofaults 460– 461 Palaeozoic palaeogeography 446, 455–459 Palaeozoic stratigraphy 455– 459
526 Anti-Atlas (Continued) para-autochthonous 433, 444– 445, 448 Precambrian basement 436–437, 445, 455 Proterozoic basement 435–436 Proterozoic extension 437, 445 Rhenish facies 467, 476 tectonic evolution 234, 265– 283, 448 tectonostratigraphy 435–437, 436 Variscan deformation 437– 444, 447– 448 Variscan reactivation 445, 446 Anti-Atlas Major Fault (AAMF) 4, 265, 329 Aoue´oua Formation 48, 49 Aoutitilt Suite 45 apatite saturation thermometry 418– 421, 420, 422– 424, 424 Areigat Lemha, kimberlite indicator minerals (KIM) 99 Armorican Massif 358 Arraiolos biotite granite 385, 386–394, 401–403 zircon data 398, 400, 401– 403, 401, 403, 404– 405 Askaoun pluton 265, 266, 267, 269, 270, 271, 278 Assa Formation 468 Assarag Suite, magmatism evolution 280–281 Assersa granite 4, 9, 11, 12, 13, 14 Atacora nappes 218 Atlas Palaeozoic Transform Fault 312, 448, 453 augen granite gneiss 46–47, 48 Avalonia 350– 354 accretion to Laurentia 354, 369 basement isotopic signature 350–351 and Cadomia 403 Cambrian 350 –351, 354 and Carolinia 369 evolution of 350, 353, 353, 368– 369 and Ganderia 369 metamorphism 351 ophiolitic rocks 353 palaeomagnetic data 354 subduction evidence 351, 353 tectonostratigraphy 351 transform activity 353–354, 353 zircon analysis 352, 353 Avalonian terranes 345, 349, 364, 369 Azegour formation 310 folding 309 geological map 309 granitoids 307, 310 lavas 316–319, 317–319, 320–322, 320 lithology 306, 307– 310 Tizgui dacite 313, 314, 315, 317 Azegour-Wirgane region 333, 334 Azguemerzi granodiorite 4, 9, 10, 11, 12, 13, 14 Bakel area 483, 484– 490, 485, 494 metamorphism 484, 493 oblique transpressional orogen 494, 494, 495 shear zones 483, 487, 488, 488, 493, 494 sinistral transpression 495 strain partitioning model 484, 492–494, 494, 495 structural geology 484, 492– 494 thrust planes 484, 494, 495 Bakel Thrust Fault 488 banded iron formations 42, 53, 71 Bani area, Variscan deformation 438– 439, 439
INDEX Bas Draˆa area 438, 444–445 Variscan deformation 437, 437– 439, 439 Bassaride belt 3, 169– 201, 177, 191 basalt geochemistry 187 chronological data 187 geodynamic model 193, 196– 197 geophysical data 187 lithostratigraphy 185, 185, 192 metamorphism 187 sedimentary cover 186– 187 Batapa Group 187 Bir Igueni Suite 42, 47–49, 48 Birmian terranes crustal genesis 19– 32 emplacement 19 granite-greenstone 501 Bissau-Kidira-Kayes fault zone (BKKF) 173 Bled El Mass diamond deposit 77– 79 Bleida area geochemistry 254 geological setting 250– 255 lithology 251– 253, 252, 253 mineralization 256–259 ophiolite complex 249–264, 265 polymictic breccia zones 255, 256 stratigraphy 251– 253, 252 structural geology 255 Bleida ore zones alteration 249, 260 –262 cobalt concentrations 249 cross sections 259 deformation events 260 deposit hypothesis 249– 250 epithermal system 249, 262 Fe-rich lithofacies 259, 259, 260 gold 250, 259, 259 Bohemian Massif 361 –362, 361 Bou Azzer-El Graara inlier 250, 250, 251, 436, 443 Variscan deformation 441–443, 443 Bou Salada Group 6 Bou-Azzer ophiolite 4– 6, 233 –247, 249, 332 blueschist facies 235, 237, 237, 329 cross sections 242 emplacement model 233– 234, 243– 244, 243 equilibrium phase diagram 238, 239 geological setting 234– 236, 242 greenschist facies 235, 237, 238 me´lange 242– 244 metamorphism 233, 235, 237– 238, 244 mineral assemblages 238, 239 Na-amphiboles 237, 237, 238, 240 Bou-Jerif area 437 Bouadil region, Palaeozoic cover 454 Bourre´ domain 204 Bove´ Basin 173, 178–179, 179, 186, 190 Cadomia 349, 358–359 Cadomian terranes 345, 364, 369 Cadomian-Avalonian orogeny 385, 386, 427 Calcaires infe´rieures 285 –302 cycles and controls 290–293, 299 depositional environment 285– 286, 289, 293, 294, 299 facies associations 289–293, 290–291, 293– 294, 300
INDEX fenestral silty-cherty dolostone 289 geochemistry 294, 295 grainy dolostone 289, 292 offshore dolomite-shale rhythmites 290 palaeotopography 299 peritidal mixed facies 289– 293, 299 shale 289, 292 slide deposits 293 –294, 293, 299, 300 stratigraphic log 290–291, 293–294 stromatic dolostone 289, 292, 293, 293 Cambrian High Atlas-Meseta rift 460 carbonatite intrusion 217– 231 age 223– 224 and continental rifting 221– 224 geochemistry 221– 223, 222 magmatic origin 222 petrography 221 REE concentrations 222– 223, 223 Carolinia terrane 357– 358 accretion to Laurentia 358, 369 and Avalonia 369 platform strata 358 tectonic evolution 357– 358, 357, 368–369 Central Atlantic magmatic province 3, 455 Central Hoggar, Algeria 111–146 Chami Greenstone Belt 40, 41, 42, 49 Chegga granite 79, 93 Chenachane shear zone 86–88 chlorite index 261– 262, 261 chloritoid, thermobarometric calculations 206 Chortis terrane 362, 363 Choum-Rag el Abiod Terrane 33, 34, 35–38, 36, 45, 50 Archaean development 49 granite emplacement 36 metamorphic lithologies 35– 36, 40–41 migmatitic orthogneiss 35, 37, 49 partial melting 38 post TISZ plutonic rocks 36– 37 Precambrian events 37–38 shear fabrics 36, 37, 37–38, 39 thermobarometric calculations 38 clinopyroxene 210, 512, 515–516, 517, 518 crystal size distribution 509, 511– 514, 513, 514, 520–521, 520, 522 mineral chemistry 514– 518, 515– 516 texture 511–517, 520–521, 520 thermobarometry 206, 514, 518– 519, 519 clinopyroxene-liquid equilibrium 518–519 copper, gold association 252, 257, 260, 263 cratonic terranes 345, 349, 364, 370 cratons 1, 83, 112 see also metacratonic evolution crystal size distribution (CSD) 505 –520 Dahomeyide orogen 217, 218, 499, 501, 502 Dahomeyide suture zone 212, 218– 221, 218, 229–230 deformed alkaline rocks and carbonatite (DARC) 217, 221– 223 garnet megacrysts 219 granitoid gneiss 219 high pressure granulites and eclogites (HIPGE) 217, 218, 219 nepheline-bearing gneiss 220, 223 rutile exsolution rods 219
subduction 228–229 tectonic stacking 218 –219, 219 Wilson tectonic cycle 229–230, 229 Dargol granodiorite 19, 27– 28 zircon data 24–27, 25, 26, 27 Dargol pluton 20 Dayet Lawda unit 181 Deep Ivorian basin 500, 501, 504 Deep Ivorian basin fault 501, 504 Diagorou-Darbani greenstone belt 19–32, 21, 22, 23–24, 27, 28– 29 geochemistry 20–24, 23, 23, 28, 30 Diagorou-Darbani metaconglomerates Nd–Sm isotopic data 24, 24 protolith 19, 23, 27 REE patterns 21– 23, 23, 23 zircon data 24, 25–27, 25, 26, 27 Diagorou-Darbani micaschists Nd–Sm isotopic data 24, 24, 27 protolith 19, 23–24, 27 REE patterns 21– 23, 23, 23 zircon data 24–27, 25, 26, 27, 28 diamond exploration 79 geophysical survey methods 83 komatiitic origin 105 lamproites 105 primary sources 77– 109 Djebel Drissa ring complex 80, 82, 85, 89 Draˆa Basin depocentre shifts 467, 477 Devonian stratigraphy 468 eustatic controls 475, 476 lateral facies change 480 Rich group 467–482 sediment supply 467, 480 subsidence variations 467, 473, 476–477, 477, 480 tectonic controls 477– 478 tectonic interpretation 478– 480, 478 Draˆa channel, principle fault directions 478–480 Eburnean metapelites 111 –146 Eburnean orogeny 1, 2, 8, 111–112, 144, 165, 431 Eglab shield 77–109, 80 aeromagnetic anomalies 84– 86, 85, 86, 105 brittle deformation 82, 83 diamond exploration 86 dolerite dykes 87, 88, 93, 97– 98, 102 ductile deformation 81–82 Eburnean 80–81, 83 garnet analysis 99–101, 103, 104 geophysical structures 83–86 igneous ring complexes 79 kimberlite indicator minerals 98–101 lamprophyric dykes 87, 88, 97–98, 102 lithospheric faults 77, 79 regional geology 79–81 Riedel model 82 structural features 81–83, 86– 98, 105 tectonic model 105 El Annsar Formation 468 El Graara massif 286, 299–300 electrum 249, 260 E´vora Massif 385, 386, 388, 389, 403
527
528 Fadrat al Garod unit 181 Fale´me´ Series 484 Fale´me´ trough 175– 179, 178– 179 Florida terrane 358 Forecariah Group 189– 190 Foum el Kous 510, 511–518, 521, 522 clinopyroxene 511– 514, 512, 513, 514, 515–516, 517, 518 magma mixing 513 mineral chemistry 514– 518 nephelinite 513, 513, 515– 516 thermobarometry 522 total alkalis silica (TAS) diagram 511, 513 fractional crystallization 336, 337 –338, 409, 422 Franciscan Complex 241 Gabou Shear Zone 486, 489 Ganderia 354– 356 accretion to Laurentia 356, 369 arc magmatism 355, 356 and Avalonia 369 tectonic evolution 354–356, 355, 368– 369 zircon analysis 355 Ganderian terranes 345, 349, 364, 368, 370 Gao rift 3 garnet 99–101, 103, 104, 207 composition 99–101, 103, 104, 209, 209, 210, 226, 226 rutile exsolution 224 –226, 225 stable reaction 226 thermobarometric calculations 206 UHP metamorphism 217 zoning 155, 157, 210 garnet-biotite geothermobarometry 56 Geyik Dag unit 409, 410, 411– 412 magmatism 422– 425, 426, 427 quartz porphyry 419, 420, 422, 424, 424, 425 Gezmayet unit 181 Ghana 217–231 Ghana, offshore 499– 508 half graben 502, 503– 504 onshore correlations 499, 500, 500, 506 Proterozoic 499, 502 transform margin formation 504 Gibi Mountains Group 189 glaucophane 203, 207, 211, 238, 239 gold 249–264 copper association 252, 257, 260, 263 depositional environments 262–263 disseminations 249–250, 256, 257, 258, 260, 263 grain chemistry 259–260, 260 grain textures 259– 260, 263 hydrothermal alteration 260– 262 palladium association 260, 262– 263 types of 249 veins 249– 250, 256, 256, 257– 259, 257, 258, 260, 263 see also electrum; isomerite Gondwana 303–327 assembly 228– 229, 505 Laurentia collision 429 supercontinent 570Ma 402 West Gondwana 345, 346, 364, 369
INDEX Goulmime area 433 Gourma belt 1, 2, 203 –216 amphibolite facies 208–209, 210 cross-section 214 eclogites 203–216, 210 garnet glaucophanite 203– 216, 210 geological setting 204– 205, 204 glaucophane bearing eclogite 208– 209, 209, 215 Internal domain 204 –205 kyanite eclogites 212 mafic rocks 207–210 metamorphism 205, 212, 213– 215 micaschists 205–207 mineral chemistry 205 –211 paragenetic analysis 205–211 structural data 204–205, 213 subduction 203 –216 Grenvillian orogeny 169 Guelb Moghrein IOCG 53– 75, 56 brecciated Fe–Mg carbonate 65, 66–67, 68, 72, 73 composition 72– 73 geochemistry 65– 68 geological setting 54–57, 56 hydrothermal alteration 54, 72, 73 lithology 65–68 mineralization age 53– 54 ore breccia 65, 68 ore fluid 54, 73 protolith geochemistry 68– 72 shear zones 65, 68, 72, 73 Guinguan Group 185, 186 Hadeibt Lebtheiniye´ Greenstone Belt 40, 42 Hajar Dekhem Massif 181 Hercynian orogeny 169–171, 196, 479 High Atlas 1, 3, 303– 327, 304, 333 Cambrian lithology 303– 312, 330, 332–337, 336, 337, 338, 339–340 Cambrian volcanics 320–321, 320, 321– 322, 321, 322, 322 Late Neoproterozoic rocks 332, 334–337, 334– 335, 336, 337, 338 stress direction 311– 312 structural patterns 311, 312 tectonics 303–312 High Atlas fault 324 Ho gneiss 218 Hoggar shield 147–167, 461 geological map 112, 148 Hudeibt Agheyane Greenstone Belt 40 hydrothermal alteration 68, 249, 260–262 hydrothermal mineralization 3– 4, 53, 56–57, 65–68, 72– 73 I-type granite melts 409, 418, 421, 427 Iapetus Ocean 368, 369 Iberia terrane 346, 359–361 Cambrian fauna 360 Cambrian rifting 361 isotopic compositions 360– 361 sedimentary source 360– 361 tectonostratigraphy 360 zircon analysis 360–361
INDEX
529
Iberian Massif biotite granites 385– 408 paragneiss 385–408 Id Boukhtir 299, 300 Calcaires infe´rieures 290–291, 293– 294, 293 Ida ou Illoun pluton 265, 266, 267, 268, 269, 270, 271, 274, 279 Ifni inlier 439– 440 Ifouachguel pluton 265, 266, 268, 269, 270, 271, 278 Ifri formation 305–307 stratigraphy 305, 306 structural features 307, 308 tectonics 306 –307 Ifri greywacke 313, 314, 315 Ifri lavas 315–317, 316– 319, 320– 322, 320 Ifri-Azegour formation 303 –310, 305 deposition 323 –324 geochemistry 313– 322, 323 –324 geochronology 312– 313, 314, 315 magmatic signature 323, 324 petrography 313– 322 zircons 313, 314, 315 Iguilid metagabbro 37 Imdghar pluton 265, 266– 267, 269, 271, 274, 279 In Edem eclogite 212 In Edem schists 215 In Ouzzal metacraton 147–167 geochronology 149 geological setting 147–149, 148 UHT metamorphism 165 Ioulguend granite 35 Irherm-Tata area 441, 442 iron oxide copper gold deposits (IOCG) 53–75, 56, 66–67, 68, 262 isomertieite 249, 260
Khanfous complex 147– 167, 148 Khzama oceanic basin 265, 266, 280 kimberlite 83, 86 kimberlite indicator minerals (KIM) 77, 84, 98–101 Kleouat Massif 181 Kolente´ basin 178 –179, 188 komatiite-picrite 79, 93– 97, 94 AFM diagram 96–97, 101 phlogopite biotite 92, 96 pyroxene 90, 90, 94– 95, 95 REE pattern 102 spinel 95, 96– 99, 100 Komba basin 178– 179 Koughany Shear Zone 486 Kpong complex 219, 219 Kreidat Greenstone Belt 40 Ku¨tahya-Bolkar Dagi unit 409, 410, 411–412, 416, 418, 426 geochemistry 416–418, 417, 418 granite 414, 418, 422– 425, 425 magma source 418, 427 Na-metasomatism 416 zircon U–Pb ages 414 –415
Jbel Boho region 329, 332 Jbel Kerkar region 332–334
magma differentiation 329, 336, 337 –338, 509, 520–521 Maider basin 456, 457– 458, 459, 460, 461, 462 malachite 259 Mali Group 185–186, 187 Man rise 2, 19– 20, 20 Marampa Group 189 Marampa thrust belt 195 Marsa Thrust Fault 484, 488, 489 Matallah unit 181 Mauritanide belt 3, 33– 52, 53–75, 169– 201, 180, 196, 453, 461 Adrar-Souttouf section 181 Akjoujt section 181 Anti-Atlas section 179 Aouker-Kidira section 182 Central domain 486– 489, 491, 492, 494 cross section 191, 485 deformed pebbles 490– 492 deformed quartz grains 490– 492 Dhlou-Sekkem section 179–181 finite strain analysis 490– 492, 492, 491 Fry analyses 491 geochronological data 183 geodynamic model 183 –185, 193, 196 –197, 483–484 geological setting 174– 175, 177, 483–484, 484 geophysical data 183, 184 Koulountou-Bove´ section 182 lithostratigraphy 192
Kasila Group 190 Kasila thrust belt 195 Kenieba inlier 480 Kerdous region 329, 331–334, 331, 439– 440 Khanfous Al–Mg granulites 147– 167 biotite 159, 161 composition 149– 150, 149 cordierite 155, 159, 162– 164 cordierite-garnet-quartz 152– 153, 153– 155, 154 feldspars 160– 161 garnet 155, 156, 157 metamorphic stages 150– 155, 162, 163 mineralogy 155– 161 orthopyroxene 155, 158, 164 orthopyroxene-sillimanite-garnet 150–151, 150 P-T evolution 161–165, 162, 165 P-T pseudosections 147, 162–165, 163, 164 petrogenetic grid 161– 165, 162, 163 quartz-garnet-biotite 150 quartzitic 149, 162 reaction textures 147, 150, 153, 154, 162 sapphirine 159, 160 sapphirine-spinel-quartz-orthopyroxene 151– 153, 152 spinel 159, 160 UHT metamorphism 151, 152, 164, 165
Lakhsass Plateau area 439– 440, 440 lamproite 83, 86 Laouni terrane 111, 113 LATEA metacratonic evolution 111 –146 Laurentia 354, 356, 358, 364, 369 Laurentia-Gondwana-Baltica 366, 367, 368 Lembeitih Formation 53, 54, 71 Leo shield 33 Lie-de-vin Formation 288, 329, 330– 331, 333, 341, 443 Liptako, Niger 19– 32, 29 Lower Reguibat Complex 80
530
INDEX
Mauritanide belt (Continued) metamorphic characteristics 183 Niame´-Gue´tie´ subdomain 487, 488 Northern domain 485–486, 491, 493 ophiolitic assemblage 183 para-autochthonous 484 Samba Kontaye-Gabou subdomain 487, 488, 489, 489 Southern domain 489–490, 491 strain partitioning 483–497, 492 structural elements 485–490, 487, 489 tectonic interpretation 183, 485, 486, 492–494 transpression 483–497 U –Pb zircon data 43 Mauritanide frontal thrust 169, 171 Mauritanides orogeny 484 Maya terrane 362, 363 M’Bout unit 182 Meguma terrane 356– 357, 369 me´langes 233– 247 and ophiolites 239 –242 see also accretionary complex; ophiolitic me´lange Menderes Massif, augen gneiss 414 Merroucha annular structure 87, 88, 92– 93, 93 Merza Akhsai Formation 468 Meseta domain 447– 448, 453, 461, 463 metacratonic evolution 1, 3, 4, 14, 111 –146 metamorphism blueschist index minerals 237 high pressure 203, 217, 228 ultra high pressure 203, 205, 217– 231, 236– 237 Middle American terranes 362–363 Middle Western High Atlas fault 303, 312 migmatitic gneiss phlebitic 44 phlebitic/stromatic 40, 42 Minas fault zone 356, 357 Moho, shield areas 499 Moldanubian zone 361 monazite 53, 73, 74 saturation thermometry 418 –421, 420 Mont Lato eclogite 227 Montemor-o-Novo shear zone 386 Moravo-Silesian zone 362 Morocco 1– 17, 233–247, 249– 264, 265– 283, 285–302, 312, 433 –452, 453–465, 467–482 Cambrian rift geodynamics 303, 323–324, 323 Cenozoic volcanic field 509– 524 Moroccan outboard magmatic evolution 329–343 Mount Wa-Wa unit 182 Moussaya complex 189 Ndaouas Suite 42, 47 Nebdas pluton 265 Neoproterozoic peri-Gondwanan terranes 345 –383, 365 subduction 212– 213 nephelinite flow 509– 524 crystal settling 521 high-Al augite 519–520, 522 magma mixing 521– 522, 523 phonolite 521– 522, 522, 523 pyroxene compositions 521– 522, 522 Niokolo Koba Group 185, 186
Nkheila Formation 468 North Africa, diamondiferous province 77–109 Oaxaquia terrane 362, 363 ophiolites 233– 247 and me´langes 239 –242 obduction 241, 241 ophiolitic melange 240– 241, 241 ore breccia 53, 69, 70, 72, 73 ore fluid 53, 54, 72, 73 orogen-parallel tectonic transport 483–497 Ossa-Morena zone 359– 360, 361, 387, 403 Ediacaran basins 385–408 sediment provenance 385– 386, 403 Ouaankifondi complex 189 Ouarzazate Supergroup 1, 3, 4, 6, 235, 266, 329, 330, 433 –437, 441 hydrothermal event 12, 14 Ouarzazate-Agdz region 329, 333, 338 geochemistry 334–337 petrography 331 –337 stratigraphy 331– 334, 331 Ougarta domain 445 Ougnat culmination 461, 462 Ougnat Massif 453, 455 Ougnat-Ouzina Ridge 455, 457, 460 Ouin-Mesdour Formation 469 Oukhit region, Palaeozoic cover 454 Oulad Dlim area 448 Oum Jerane-Taouz fault 455, 459 Ouzaga formation 303 Ouzaga-Tizzirt formation 305, 310–311 deposition 323–324 geochemistry 313–322, 323–324 geochronology 312–313, 314, 315 magmatic signature 323, 324 petrography 313 –322 stratigraphy 310– 311, 311 Ouzaga-Tizzirt lavas 316–319, 319–322, 320 Palaeozoic, peri-Gondwanan terranes 345– 383, 365, 368 –369 palladium 249, 258, 260, 262–263 Pan-African belt 3, 4, 8, 33, 251, 322, 501, 505 eclogites 212–213 metacratonic evolution 10–14 re-activated fractures 448 Pan-African ocean, subduction 233, 243, 243 Pan-African orogeny 3, 4, 15, 54, 111– 112, 144, 169 –171, 285, 287, 329, 433, 436, 499, 501 I 195, 196 II 190, 195 –196 me´langes 233 –247 ophiolites 233–247, 441 re-mobilization 12, 14 Pan-African suture zone 203, 205, 212, 217– 231, 501, 505, 506 Pan-African-Cadomian granitic magmatism 409 paragonite, thermobarometric calculations 206 peri-Gondwanan terranes 3, 15, 346– 349, 347, 349, 404, 409 accretion to Laurentia 368, 370 Appalachian orogen 347, 347 arc magmatism 347, 364
INDEX basement composition 345, 349, 369 Brazil 363 Cambrian fauna 346, 370 Europe 347– 349, 348, 363 evolution 369 faunal affinities 349, 368 geodynamic models 364, 370, 427 geological setting 346–350 magmatic history 350–363 Middle America 347, 348 Middle East 363 Neoproterozoic 364–368, 365 North America 347, 347 palaeogeography 345– 383, 366, 367, 368, 369 Palaeozoic 365, 368–369 passive margin sequences 347, 368 sedimentary provenance 3, 346, 349, 370, 385 strike slip tectonics 364– 368 tectonostratigraphy 345– 383 zircon analysis 345–346, 349, 402 phengite, thermobarometric calculations 206 Pita group 179 platinum group element (PGE) 250 pyroxene solvus thermometer 515 Reggane placer deposit 77, 101 Reguibat rise 19, 77, 173 Reguibat shield 2, 33– 52, 34, 43, 49–50, 101, 102– 104 Rheic Ocean 368 Rich group 468, 469– 472, 479 basal limy facies 469–471, 471, 474– 475 clastic facies 472, 474– 475 deformation 477, 478 depositional environments 469– 472 depositional sequences 467, 474– 475, 476– 477, 476, 477, 480 lateral variations 468, 471, 476 sea level changes 467, 473, 480 sediment supply 476– 477 sedimentary structures 469– 472, 470 sequence boundaries 472–474 stratigraphy 467–469, 470, 472 –474, 473, 476, 477 and systems tracts 474–475, 476, 480 tectonic control 477–478 riebeckite 238, 239 Rodinia breakup 228, 230 Rodinia reconstruction 366 Rokel River basin 178–179, 188– 189 Rokelide belt 169–201, 177 cross section 191 foreland 188– 189 geochemical studies 190 geodynamic model 193, 196– 197 gravimetric data 190 lithostratigraphy 192 metamorphic basement 189 radiometric data 190– 195 tectonothermal event 190–195 Rokelide frontal thrust 188 Rokelide orogenic cycle 190 Romanche fracture zone 500 Romanche transform margin 499–508
531
bathymetric map 499, 500 tectonic setting 495–501 rutile exsolution rods 224–226, 225, 226, 227 S-type granite melts 409, 421 Saghro massif 266, 285–302, 286, 441, 455 Saghro volcanic field 510, 510 carbonatitic magmatism 510 clinopyroxene 514–517 magma chamber depths 519 –521, 520, 522– 523 Saghro-Ougnat area, Variscan deformation 443– 444, 444 Saghro-Ougnat axis 457, 458, 462 Sainte Barbe volcanic unit 53, 54– 55, 56, 57–58, 72– 73 biotite-garnet-quartz schist 58 quartz-sericite schist 57, 63, 68 trace element data 53, 58–60, 61– 62 Sandikli area 409, 411–412 Saxo-Thuringian zone 362, 385 sea level changes 467, 473, 474– 475, 480 Sebkhet Nich Greenstone Belt 40 seismic reflection profiles 499– 508 half graben 502, 503–504 Palaeoproterozoic intrusions 502 subduction zones 499 Winneba greenstone belt 502, 503– 504 Senegal 483–497 sericite index 261 –262, 261 Se´rie Negra Formation greywackes 404 paragneiss zircon analysis 385, 386–401, 394–395, 396 –397, 399, 400, 401, 403– 405 rift sequence 403 sediment sources 404, 405 Seyna Bela eclogite 210, 211 Seyna Bela garnet glaucophane 203, 207–210, 208–209 Shai Hills garnet mafic granulites 224, 227 garnet composition 226, 226, 227 high pressure metamorphism 224 –228, 227 thermobarometry 224, 227, 227 Titanium concentrations 226–227, 227 ultra high pressure metamorphism 224– 227, 225 Shai Hills gneiss HIPGE 218, 220–221 Sid El Houssein ring complex 6 –8 Sirba greenstone belt 27 Siroua massif 265–283, 266, 267 amphibole 270, 270 basic and intermediate rocks 274–276, 277 biotite composition 270 calc-alkaline group 274–276, 276 felsic rocks 276–279 geodynamic setting 274 –280 magma affinities 274–279, 277, 278, 279 magmatism evolution 265, 280– 281 mineral composition 269 –271, 269, 270 petrology 265– 283 pyroxene 269– 270, 269 tholeiitic group 274– 276, 276 whole rock chemistry 271–274, 272–274, 275 Sirwa complex 4, 510 Souss Basin 286–289 South Atlas Fault 4, 265, 329, 338– 341, 448, 453 Souttoufide event 169, 196 subduction settings 215
532
INDEX
subduction zones metamorphism 203–216, 236 metamorphism pressures 217 seismic reflection data 499 thermal gradients 234, 236–237, 236 Suwannee terrane 358 Taban Group 189 Tabia Member 294– 295 Tabouna complex 189 Tac¸arat-Inemmaudene Shear Zone (TISZ) 33–35, 36, 36, 44–46 augen granite gneiss 44, 45, 45, 46–47 cross-section 46, 46 flower structure 33, 46, 50 Nd isotopes 47 porphyritic granite 45 shearing age 47, 50 transpression 50 xenolithic granite 44, 45 Tafilalt basin 457–458, 460, 461 Tafilalt platform 456, 457, 462 Taghdout Group 4 Taka pluton 20 Tamanrasset, metamorphic evolution 137 Tamarouft granite 4, 9, 11–12, 13, 14 Tamwirine rhyolitic unit 6 Taoudeni basin 1, 2, 3, 33, 169, 172, 173–179 lithostratigraphy 173– 175, 176 Taganet sub-basin 173–175 Tambaoura sub-basin 173 Taroudant Group 285, 435 Basal Series 288 stratigraphy 287 see also Adoudou Formation Tasiast-Tijirit terrane 33, 38–44, 41, 42, 50 Archaean development 49 banded iron formations 42 biotite tonalites 41–42, 44 geochronology 47 greenstone belts 40 magma source 49–50 migmatitic gneiss 33, 40, 42, 49 Nd isotopes 48 pegmatitic muscovite granite 44 shear fabrics 39 tectonothermal events 44 zircon analysis 47– 49, 47, 48 Tata fault 441 Tata Group 4 Tauride-Anatolide Belt 410, 415 Tazenakht granite 4, 9– 11, 12, 13, 14 tectonics thick skinned tectonics 1, 3, 444 –445, 447, 453, 463 thin skinned tectonics 1, 3, 433, 484 Telime´le´ group 179 Te´ra pluton 20 Termesse Group 185, 186 –187 thermobarometry 203, 206, 210– 211, 224, 509–524 clinopyroxene-liquid equilibrium 518– 519 crystallization pressure 518– 519 structural barometer 518– 519 Tichka pluton 307, 310, 311 Tiddiline Formation 235, 441
Tidjenouine granulites 113 –116, 143 gedrite bearing 119, 129, 131, 133, 135 metamorphism 113, 138–142, 144 migmatitic biotite-garnet-sillimanite metapelites 113 P-T paths 113, 115, 133, 135, 143 zircon analysis 113, 138– 143, 139– 140, 141, 142 Tidjenouine migmatitic gneiss 113–115, 114, 115 Tidjenouine migmatitic metapelites 115–116 amphibolite facies 138, 143 biotite 121, 122 composition 115 cordierite 121, 123 decompressional metamorphism 135, 137–138, 143 garnet 121, 124, 129 geobarometers 133, 136 geochemical data 114 geochronology 111–146 geothermometers 133, 136 granulite metamorphism 133, 137, 143 mineral assemblages 116 –121, 116, 143 mineral chemistry 121 –133 orthoamphibole 126, 129, 130, 131 orthopyroxene 121, 125, 129, 129 orthopyroxene-bearing sillimanite-free 119– 121, 120 orthopyroxene-free corundum-bearing 117 orthopyroxene-free quartz-bearing 116–117, 118– 119, 132– 133 P-T evolution 111–146, 132–133, 134, 135, 136, 137, 137 P-T pseudosection 133, 135 peak metamorphism 135– 137, 143 petrogenetic grid 132 –133, 133, 134, 135 petrology 133–135 plagioclase 128, 130–133 quartz-bearing 133 quartz-free 134 reaction sequences 133, 134 reaction textures 116– 117, 118, 120, 133 retrograde metamorphism 135, 138, 143 secondary orthopyroxene-bearing 117– 119, 118– 119, 129 spinel 127, 129 Tifnout Member 294–295 Tijirit Greenstone Belt 40 Tikirt Member 286, 288, 299, 443 Timrhanrhart Formation 469 Tin Begane, metamorphic evolution 137 Tin Hama eclogite 203, 207, 208–209, 210, 211 Tindouf basin 2, 169, 171–173, 467 titanium solubility 225– 227, 227 Tizi n-Taghatine Group 4 Tizi-n’Test fault 303 tonalite trondhjemite granodiorite (TTG) 19, 271, 278 Touijenert-Modreı¨gue Granite 34, 36 Touyerma Granite 36 Trans-Saharan shear zone 3, 501 Tuareg shield 3, 77, 111, 112, 112, 144, 212 Turkey, Neoproterozoic magmatism 409–431 Upper Reguibat Complex 80 Variscan deformation 433– 452, 453– 465 Devonian extension 461– 463 fault kinematics 438, 461–462, 462
INDEX
533
folds 441– 443, 445, 446 greenschist facies metamorphism 447 palaeofault inversion 453, 463 Precambrian basement 435 shortening 445–447, 462 Variscan orogeny 1, 3, 8, 15, 54, 72, 303, 311, 324, 345, 433, 445, 453, 483 Voltas basin 3
Yetti domain 79, 89 Yetti malignite 87, 88, 88, 89–92 nepheline 90, 92 phlogopite biotite 90, 91, 92 pyroxene composition 89, 90 Yetti-Eglab Junction 77, 79, 81, 82, 84, 88, 105 Youkounkoun basin 178– 179, 179 Youkounkoun Group 186
West Africa, Amazonian connections 345–383 West African Craton 2 –3, 2, 20, 33, 78, 79, 169, 170, 196, 228, 478–480, 501 bathymetric map 500 collision 228– 229, 230, 244 deep structure 499–508 Devonian extension 453–465 geodynamic evolution 54, 194–195, 196– 197, 322–324 lithospheric faults 3 –4, 84 Neoproterozoic-Palaeozoic 1, 329–343 Ossa-Morena Ediacaran basins 385– 408 West African Craton boundaries 2, 3 –4, 15, 346 eastern 217–218 northern 79, 233– 247, 303 –327, 433–452, 453– 465, 467, 478 –480 southeastern 499–508, 501, 502 western 53–75, 169 –201, 194 –195, 196–197 West African orogenic belts 172, 195– 197 Winneba greenstone belt 502, 503–504 wrench style tectonics 483, 486
Zemmour belt 448 Zenaga plutons 4– 8, 6, 9 –10, 13 composition 7, 8– 9, 10– 14, 10–11, 12, 13, 13, 14 fluid influence 11, 14 peraluminous 10, 12, 14 petrography 9– 10 zircon analysis 24– 27, 25, 26, 27, 28, 28, 46, 138, 345–346, 386–395 Hf signature 47, 47 morphology 414– 415, 414, 415 spot analysis 141, 142 Tera-Wasserburg diagrams 395, 399, 400, 402 textures 141, 394–395, 395–397, 396–397, 398 U– Pb dating 46, 111, 385– 404, 396– 397, 398, 399, 400, 414–415, 414, 415 zoning 394– 395, 395, 396–397, 398, 401 –403 zircon saturation thermometry 418– 421, 420, 422–424, 424 zircon solubility model 418
The boundaries of rigid cratons can be affected by subsequent orogenic events, leading to ‘metacratonic’ characteristics not often properly recognized and still poorly understood. Major lithospheric thickening is absent and early events such as ophiolites are preserved; however, metacratonic boundaries are affected by major shear zones, abundant magmatism and mineralizations, and local high-pressure metamorphism. West Africa, marked by the large Eburnian (c. 2 Ga) West African craton, the absence of Mesoproterozoic events, the major Pan-African (0.9–0.55 Ga) mobile belts that generated the Peri-Gondawanan terranes, and the weaker but enlightening Variscan and Alpine orogenies, is an excellent place for tackling this promising concept of metacratonization. The papers in this book consider most of the West African craton boundaries, from the reworking of the Palaeoproterozoic terranes, through the Pan-African encircling terranes, the late Neoproterozoic–early Palaeozoic extension period and the Peri-Gondwanan terranes, the Variscan imprint to the current situation.