Edited by Ulf Linnemann, R. Damian Nance, Petr Kraft, and Gernold Zulauf
-
THE GEOLOGICAL SOCIETY OF AMERICA®
Special Paper 423
The Evolution of the Rheic Ocean: From Avalonian-Cadomian Active Margin to Alleghenian-Variscan Collision
Edited by Ulf Linnemann Staatliche Naturhistorische Sammlungen Dresden Museum für Mineralogie und Geologie Königsbrücker Landstrasse 159 D-01109 Dresden Germany R. Damian Nance Department of Geological Sciences 316 Clippinger Laboratories Ohio University Athens, Ohio 45701 USA Petr Kraft Charles University Prague Institute of Geology and Paleontology Albertov 6 128 43 Prague 2 Czech Republic Gernold Zulauf Institut für Geowissenschaften Universität Frankfurt am Main Senckenberganlage 32-34 60054 Frankfurt am Main Germany
Special Paper 423 3300 Penrose Place, P.O. Box 9140
Boulder, Colorado 80301-9140 USA
2007
Copyright © 2007, The Geological Society of America, Inc. (GSA). All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact the Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editor: Marion E. Bickford Library of Congress Cataloging-in-Publication Data The evolution of the Rheic Ocean : from Avalonian-Cadomian active margin to Alleghenian-Variscan collision / edited by Ulf Linnemann ... [et al.]. p. cm. — (Special paper ; 423) Includes bibliographical references and index. ISBN 978-0-8137-2423-2 (pbk.) 1. Geology, Stratigraphic—Paleozoic. 2. Plate tectonics. 3. Rifts (Geology). 4. Continental drift. 5. Continental margins. 6. Formations (Geology). 7. Rheic Ocean. I. Linnemann, Ulf. QE654 .E96 2007 551.7′2—dc22 2007061018 Cover: Isoclinal folded turbidites of a Cadomian retroarc basin, Late Neoproterozoic (Ediacaran), Lausitz Group, Lausitz antiform, Saxo-Thuringian zone, Bohemian Massif, quarry Butterberg near the city of Kamenz, Germany. Deformation is Cadomian (earliest Cambrian). (Photograph by Ulf Linnemann.)
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Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .vii 1. The assembly of West Gondwana—The view from the Rio de la Plata craton . . . . . . . . . . . . . . . . 1 Kerstin Saalmann, Léo A. Hartmann, and Marcus V.D. Remus 2. Geodynamic evolution of the northwestern Paleo-Gondwanan margin in the Moroccan Atlas at the Precambrian-Cambrian boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 Andre Pouclet, Abdellatif Aarab, Abdelilah Fekkak, and Mohammed Benharref 3. The continuum between Cadomian orogenesis and opening of the Rheic Ocean: Constraints from LA-ICP-MS U-Pb zircon dating and analysis of plate-tectonic setting (Saxo-Thuringian zone, northeastern Bohemian Massif, Germany) . . . . . . . . . . . . . . . . . 61 Ulf Linnemann, Axel Gerdes, Kerstin Drost, and Bernd Buschmann 4. The Lausitz graywackes, Saxo-Thuringia, Germany—Witness to the Cadomian orogeny . . . . . . 97 Helga Kemnitz 5. Paleontological data from the Early Cambrian of Germany and paleobiogeographical implications for the configuration of central Perigondwana . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 Olaf Elicki 6. The Variscan orogeny in the Saxo-Thuringian zone—Heterogenous overprint of Cadomian/Paleozoic Peri-Gondwana crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 153 U. Kroner, T. Hahn, Rolf L. Romer, and Ulf Linnemann 7. Far Eastern Avalonia: Its chronostratigraphic structure revealed by SHRIMP zircon ages from Upper Carboniferous to Lower Permian volcanic rocks (drill cores from Germany, Poland, and Denmark). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 Christoph Breitkreuz, Allen Kennedy, Marion Geißler, Bodo-Carlo Ehling, Jürgen Kopp, Andrzej Muszynski, Aleksander Protas, and Svend Stouge 8. Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks in response to changing geotectonic regimes: A case study from the Barrandian area (Bohemian Massif, Czech Republic) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 191 Kerstin Drost, Rolf L. Romer, Ulf Linnemann, Oldřich Fatka, Petr Kraft, and Jaroslav Marek 9. The diversity and geodynamic significance of Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the northern part of the Bohemian Massif: A review based on Sm-Nd isotope and geochemical data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 209 Christian Pin, R. Kryza, T. Oberc-Dziedzic, S. Mazur, K. Turniak, and Jarmila Waldhausrová
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10. Sm-Nd isotope and trace element study of Late Proterozoic metabasalts (“spilites”) from the Central Barrandian domain (Bohemian Massif, Czech Republic) . . . . . . . . . . . . . . . . 231 Christian Pin and Jarmila Waldhausrová 11. Structural evolution of the Prague synform (Czech Republic) during Silurian times: An AMS, rock magnetism, and paleomagnetic study of the Svatý Jan pod Skalou dikes. Consequences for the nappes emplacement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249 Tahar Aïfa, Petr Pruner, Martin Chadima, and Petr Štorch 12. Cadomian and Variscan metamorphic events in the Léon domain (Armorican Massif, France): P-T data and EMP monazite dating . . . . . . . . . . . . . . . . . . . . . . . 267 Bernhard Schulz, Erwin Krenn, Fritz Finger, Helene Brätz, and Reiner Klemd 13. U-Pb depositional age for the upper Barrios Formation (Armorican Quartzite facies) in the Cantabrian zone of Iberia: Implications for stratigraphic correlation and paleogeography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 287 Gabriel Gutiérrez-Alonso, Javier Fernández-Suárez, Juan Carlos Gutiérrez-Marco, Fernando Corfu, J. Brendan Murphy, and Mercedes Suárez 14. Contrasting mantle sources and processes involved in a peri-Gondwanan terrane: A case study of pre-Variscan mafic intrusives from the autochthon of the Central Iberian Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 297 Miguel López-Plaza, Mercedes Peinado, Francisco-Javier López-Moro, M. Dolores Rodríguez-Alonso, Asunción Carnicero, M. Piedad Franco, Juan Carlos Gonzalo, and Marina Navidad 15. Tectonic evolution of the upper allochthon of the Órdenes complex (northwestern Iberian Massif): Structural constraints to a polyorogenic peri-Gondwanan terrane . . . . . . . . . 315 Juan Gómez Barreiro, José R. Martínez Catalán, Ricardo Arenas, Pedro Castiñeiras, Jacobo Abati, Florentino Díaz García, and Jan R. Wijbrans 16. Crustal growth and deformational processes in the northern Gondwana margin: Constraints from the Évora Massif (Ossa-Morena zone, southwest Iberia, Portugal) . . . . . . . . 333 M. Francisco Pereira, J. Brandão Silva, Martim Chichorro, Patrícia Moita, José F. Santos, Arturo Apraiz, and Cristina Ribeiro 17. The Lower–Middle Cambrian boundary in the Mediterranean subprovince. . . . . . . . . . . . . . . . 359 Rodolfo Gozalo, Eladio Liñán, María Eugenia Dies Álvarez, José Antonio Gámez Vintaned, and Eduardo Mayoral 18. Avalonian and Baltican terranes in the Moesian Platform (southern Europe, Romania, and Bulgaria) in the context of Caledonian terranes along the southwestern margin of the East European craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 375 Martin S. Oczlon, Antoneta Seghedi, and Charles W. Carrigan 19. Crete and the Minoan terranes: Age constraints from U-Pb dating of detrital zircons. . . . . . . . 401 G. Zulauf, S.S. Romano, W. Dörr, and J. Fiala 20. Geological evolution of middle to late Paleozoic rocks in the Avalon terrane of northern mainland Nova Scotia, Canadian Appalachians: A record of tectonothermal activity along the northern margin of the Rheic Ocean in the Appalachian-Caledonide orogen . . . . . . 413 J. Brendan Murphy
Contents
21. Vestige of the Rheic Ocean in North America: The Acatlán Complex of southern México . . . . 437 R. Damian Nance, Brent V. Miller, J. Duncan Keppie, J. Brendan Murphy, and Jaroslav Dostal 22. Provenance of the Granjeno Schist, Ciudad Victoria, México: Detrital zircon U-Pb age constraints and implications for the Paleozoic paleogeography of the Rheic Ocean . . . . . . . . . 453 R. Damian Nance, Javier Fernández-Suárez, J. Duncan Keppie, Craig Storey, and Teresa E. Jeffries 23. Ordovician calc-alkaline granitoids in the Acatlán Complex, southern México: Geochemical and geochronologic data and implications for the tectonics of the Gondwanan margin of the Rheic Ocean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 465 Brent V. Miller, Jaroslav Dostal, J. Duncan Keppie, R. Damian Nance, Amabel Ortega-Rivera, and James K.W. Lee 24. Ordovician–Devonian oceanic basalts in the Cosoltepec Formation, Acatlán Complex, southern México: Vestiges of the Rheic Ocean? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 477 J. Duncan Keppie, Jaroslav Dostal, and Mariano Elías-Herrera 25. P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlán Complex of southern México. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 489 Matt Middleton, J. Duncan Keppie, J. Brendan Murphy, Brent V. Miller, R. Damian Nance, Amabel Ortega-Rivera, and James K.W. Lee 26. Life and death of a Cambrian–Ordovician basin: An Andean three-act play featuring Gondwana and the Arequipa-Antofalla terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 511 Sven O. Egenhoff 27. A Late Ordovician ice sheet in South America: Evidence from the Cancañiri tillites, southern Bolivia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 525 Frank Schönian and Sven O. Egenhoff 28. Sedimentary basins in the southwestern Siberian craton: Late Neoproterozoic– Early Cambrian rifting and collisional events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 549 Julius Konstantinovich Sovetov, Anna Evgen’evna Kulikova, and Maxim Nikolaevich Medvedev 29. Aluminum phosphate in Proterozoic metaquartzites: Implications for the Precambrian oceanic P budget and development of life . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 579 Giulio Morteani, Dietrich Ackermand, and Jörg Trappe Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 593
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This book contains contributions from keynote, invited, and volunteered papers presented at several meetings of the International Geological Correlation Program (IGCP) Project 497, “The Rheic Ocean: Its origin, evolution and correlatives” (2004–2008, http://www.snsd.de/igcp497/). Most of the papers are based on presentations given at the following meetings organized by members of IGCP 497: • “Gondwanan margin of the Rheic Ocean in the Bohemian Massif,” opening meeting of IGCP (Prague, Czech Republic, 16–25 September 2004); • Special symposium “Neoproterozoic to Early Paleozoic orogenic processes at the northern margin of Gondwana,” thirty-second International Geological Congress (Florence, Italy, 28 August 2004); • Special symposium “Acatlan complex, southern Mexico: Part of the Iapetus, Rheic or paleo-Pacific Ocean?” IV Reunión Nacional De Ciencias De La Tierra (Queretero, Mexico, 31 October–5 November 2004) ; • “Devono-Carboniferous evolution of the northern margin of the Rheic Ocean,” first annual meeting of IGCP 497 (Portsmouth, UK, 5–11 July 2005); and • Special session “Assembling Avalon and other Peri-Gondwanan terranes,” annual meeting of the Geological Association of Canada (Halifax, 17 May 2005). The Rheic Ocean, which is the focus of IGCP Project 497, was one of the dominant oceans of the Paleozoic. Its origins can be traced back to the Late Neoproterozoic and the plate tectonic processes responsible for the development of the Avalonian-Cadomian orogenic belt that culminated around the Precambrian-Cambrian boundary. The opening of the Rheic Ocean between the continents of Baltica and Avalonia to the north and the supercontinent Gondwana to the south occurred in the Cambro-Ordovician. The ocean reached its widest extent during the Silurian, as its predecessor, the Iapetus Ocean, closed. Closure of the Rheic Ocean began in the Lower Devonian and ended with the formation of the Variscan-Appalachian-Ouachita orogenic belt during the Carboniferous assembly of the supercontinent Pangaea. The history of the Rheic Ocean involves North and South America, Africa, Baltica, and a number of peri-Gondwanan terranes (Fig. 1). This history documents a chain of global events and produced orogenic belts that extend discontinuously from México to easternmost Europe in the Dobrogea (Romania) and Turkey. The ocean’s evolution was responsible for the formation of a wide variety of sedimentary basins. It significantly affected the history of life and profoundly influenced contemporary paleoclimate and global environmental conditions. The fields of research involved in its study, therefore, range widely and, as this book illustrates, involve stratigraphy, sedimentology, paleontology, paleogeography, paleooceanography, igneous and metamorphic petrology, tectonics, structural geology, provenance analysis, geochemistry, geochronology, and paleomagnetism. Despite decades of research, however, aspects of the evolution of the Rheic Ocean remain controversial. With this book, we hope to answer several important questions and to encourage further research. Many geoscientists have been involved in the review process that made this book possible, and their helpful suggestions and criticisms of the original manuscripts greatly improved the quality of the papers this book contains. We are grateful to the following colleagues who kindly provided these reviews: Luis vii
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Silurian (ca. 440 Ma) Panthalassa Ocean Siberia Equator
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Figure 1. Paleogeography of the Iapetus and Rheic oceans in the lower Silurian (440 Ma). A— Armorica (Brittany, Normandy, Central Massif); B—Barrandian; C—Carolina; EWI—England, Wales, and southern Ireland; F—Florida; I—Iberia; M—Méxican terranes; NF—New Foundland; NS—Nova Scotia; PA—proto-Alps; RH—Rheno-Hercynian; SX—Saxo-Thuringian; TP—Turkish plate. Modified after C.R. Scotese (Paleomap Web site: www.scotese.com).
Eguiluz Alarcón, Loren E. Babcock, Mervin J. Bartholomew, Patrick J. Brenchley, Dennis Brown, Luis A. Buatois, Bernd Buschmann, Quentin G. Crowley, Zoltan de Cserna, Richard S. D´Lemos, Jean-Bernhard Edel, Fritz Finger, Stanley Finney, Peter A. Floyd, Wolfgang Franke, Edward S. Grew, James Hibbard, Jindrich Hladil, Jana M. Horák, Rolet Joël, Susan C. Johnson, Wolfgang Kramer, Jean-Paul Liégeois, Eladio Liñán, Stanislaw Mazur, Franz Neubauer, Fernando Ortega Gutiérrez, Florentin Paris, Tim Pharaoh, Victor N. Puckkov, Cecilio Quesada, Victor Ramos, Scott D. Samson, Graham Shields, Petr Štorch, Don Tarling, Martin Timmermann, Petek Ayda Ustaomer, Jürgen F. von Raumer, John A. Winchester, Armin Zeh, Andrzej Żelaźniewicz, and Andrey Yu. Zhuravlev. Four of the reviewers decided to remain anonymous. Ulf Linnemann R. Damian Nance Petr Kraft Gernold Zulauf
Geological Society of America Special Paper 423 2007
The assembly of West Gondwana—The view from the Rio de la Plata craton Kerstin Saalmann* Geologisch-Paläontologisches Institut, J.W. Goethe-Universität Frankfurt, Senckenberganlage 32-34, D-60054 Frankfurt am Main, Germany Léo A. Hartmann* Marcus V.D. Remus* Instituto de Geociências, Universidade Federal do Rio Grande do Sul, Caixa Postal 15001, 91501-970 Porto Alegre, Rio Grande do Sul, Brazil
ABSTRACT The southern Brazilian Shield comprises a number of tectonostratigraphic blocks representing two terranes. The São Gabriel block consists of relics of two Brasiliano juvenile magmatic arcs; the Porongos belt located on the Encantadas block formed in a passive margin setting. Plate tectonic evolution started with opening of an oceanic basin to the east of the Rio de la Plata craton since at least 0.9–1.0 Ga. An intra-oceanic island arc formed due to eastward subduction and was subsequently accreted to the eastern margin of the Rio de la Plata craton. Westward subduction beneath the newly formed active continental margin occurred between ca. 850 and 700 Ma. At the same time, the Porongos basin formed on stretched continental crust of the Encantadas passive margin. Collision of the two terranes took place at ca. 700–660 Ma followed by left-lateral ductile shear along the Dorsal de Canguçu Shear Zone between 670 and 620 Ma and 630- to 610-Ma sinistral shearing in the Dom Feliciano belt farther east. The episodic character of orogenic evolution can be observed throughout Brazil. The Brasiliano belts cannot be directly linked with pan-African belts in southwestern Africa because main deformation in the latter occurred 50–70 Ma later. The assembly of Gondwana comprises a series of collisions of cratons and microcontinents over a time span of nearly 400 Ma; however, a number of orogenic episodes can be discriminated. Their synchroneity suggests that temporally equivalent episodes are coupled with the global plate tectonic framework, which, however, is far from resolved. Keywords: West Gondwana, Rio de la Plata craton, Brasiliano orogeny, Neoproterozoic, Gondwana assembly
*Present address, Saalmann: Geological Survey of Finland, P.O. Box 96, 02151 Espoo, Finland;
[email protected]. E-mail, Hartmann:
[email protected]. E-mail, Remus:
[email protected]. Saalmann, K., Hartmann, L.A., and Remus, M.V.D., 2007, The assembly of West Gondwana—The view from the Rio de la Plata craton, in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, p. 1–26, doi: 10.1130/2007.2423(01). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Figure 1. Map of Gondwana showing Neoproterozoic belts between cratons and major sutures. G—Gawler craton; P/Y—Pilbara/Yilgarn craton. Indo-Antarctica and East Gondwana events after Boger and Miller (2004); Peri-Gondwana after Linnemann et al. (2000), with age data compiled from various papers in Dörr et al. (2004); data for the Arabian-Nubian Shield from Abdelsalam and Stern (1996); west Gondwana data from Machado et al. (1996), Pimentel et al. (1999), Hartmann et al. (2000a), Alkmim et al. (2001), Pedrosa-Soares et al. (2001), and Saalmann et al. (2005c).
INTRODUCTION The reconstruction of the amalgamation of West Gondwana (Fig. 1), a mosaic made of Archean and Paleo- and Mesoproterozoic cratonic nucleii, still holds uncertainties in the timing and structural mechanisms of collisional events between different crustal blocks and cratons. Unknown occurrences and widths of former oceanic basins, lack of detailed studies on the structural evolution and kinematics of distinct tectonostratigraphic units, and a small number of precise ages of orogenic belts greatly limit the understanding of the assembly. Compared to other areas, little attention was paid to the South American Rio de la Plata craton and hence little is known about its role and position relative to other South American cratons and to the African cratons between Rodinia breakup and Gondwana assembly. Precambrian units in southernmost Brazil (in the state of Rio Grande do Sul) record a tectonometamorphic history beginning in the Archean. A number of major tectonostratigraphic units and Neoproterozoic belts that formed during the Brasiliano orogenic cycle can be distinguished. This article reviews the structural geometry and evolution of Neoproterozoic belts at the eastern margin of the Rio de la Plata craton, focusing mainly on the schist belts. We present a plate
tectonic model for the Brasiliano orogenic cycle in this region and compare the structural evolution and age data with other Neoproterozoic belts in Brazil, southwestern Africa, and other Gondwana belts to fit these events into a broader plate tectonic framework. Neoproterozoic Tectonostratigraphic Units in the Southern Brazilian Shield Based on lithostratigraphy, petrography, geophysical data, and geochemistry, a number of major tectonostratigraphic units (Fig. 2C) can be distinguished in southern Brazil (Jost and Hartmann, 1984; Soliani, 1986; Fragoso-César, 1991; Chemale et al., 1995; Hartmann et al., 1999; Heilbron et al., 2004). Parts of the Rio de la Plata craton are exposed in the southwestern and western portions of the Southern Brazilian Shield, including the Taquarembó block, which is part of the Rio de la Plata craton and consists of Archean to Paleoproterozoic granulites and gneisses (Hartmann, 1998; Hartmann et al., 1999). Major crustal accretion occurred during the 2.26- to 2.00-Ga Trans-Amazonian orogeny (Santos et al., 2003), which represents the most important orogenic and crust-building event in this region (Hartmann et al., 2000a; Hartmann and Delgado, 2001). The Taquarembó block (Rio de la Plata craton) is tectonically juxtaposed against
The assembly of West Gondwana—The view from the Rio de la Plata craton
300 600 km
LA LP PR DF F Por R SGB T
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Luís Alves block Rio de la Plata craton Paraná craton Dom Feliciano belt Florianópolis Batholith Porongos belt Ribeira belt São Gabriel block Tebicuary river area
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A Camaquã basin
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Figure 2. (A) Location of the study area in southernmost Brazil. (B) Map of southern Brazil showing the main cratons and Brasiliano tectonic units (compare also Fig. 10). (C) Geologic map of southern Brazil and its tectonostratigraphic units (modified from Fernandes et al., 1992 and Chemale et al., 1995). (D) Delineation of geophysical domains based on magnetic and gravimetric data (simplified after Costa, 1997, and Chemale, 2000). DCSZ—Dorsal de Canguçu Shear Zone; SB-ACPL—Sierra-Ballena/Alferez-Cordillera-Punta del Este lineament.
the São Gabriel block, which comprises Neoproterozoic juvenile calc-alkaline granites and gneisses (Cambaí Complex) (Babinski et al., 1996; Leite et al., 1998) intruding a metamorphic volcanosedimentary sequence (Palma Group). The lower Palma Group consists of ultramafic and mafic metavolcanic rocks interleaved with pelitic and quartzitic schists and paragneisses. The upper Palma Group comprises andesitic and dacitic low-grade metamorphic volcanic and volcanoclastic rocks, intermediate tuffs, and tuffaceous rocks, as well as psammitic and pelitic schists. The Cambaí Complex consists of voluminous juvenile deformed diorites, tonalites, and trondhjemites cut by different generations of dikes and veins of trondhjemitic, granitic, and pegmatitic
composition. Their calc-alkaline chemical composition indicates a magmatic arc environment (Silva-Filho and Soliani, 1987; Remus, 1990; Chemale et al., 1995; Babinski et al., 1996). Southwest–northeast-oriented, elongated lenticular bodies of synkinematic granites (Sanga do Jobim granite) with thicknesses varying from several meters to hundreds of meters across-strike intrude the lower Palma Group succession parallel to the main foliation. The Santa Zélia granite, exposed in the westernmost parts of the São Gabriel block (Fig. 3), shows a magmatic and subsolidus foliation that parallels the foliation of the country rocks. Its fabric indicates a synkinematic, but latetectonic, emplacement.
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C 562 Ma Caçapava granite
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Lavras 595 Ma granite
Taquarembó block (Rio de la Plata craton)
Jaguarí granite
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Palma region
2
Sanga do Jobim section
3
Vila Nova region, Bossoroca belt
4
Cerro Mantiqueiras ophiolite, Passínho metadiorite
cities: L C
Lavras do Sul Caçapava do Sul
Figure 3. Geological map of the São Gabriel block (modified from Remus et al., 2000b). Fm—formation.
The southwest–northeast-oriented Porongos belt is exposed to the east of the São Gabriel block (Fig. 2). It has a length of ~170 km and width of 15–30 km. The contact between the two schist belts is covered by unmetamorphosed, late Neoproterozoic to early Paleozoic sedimentary and volcanic rocks of the Camaquã basin. The Porongos belt consists of greenschist to amphibolite–facies metavolcanic and metasedimentary rocks of possible Meso- or Neoproterozoic age with exposures of Paleoproterozoic basement. The belt can be subdivided into western and eastern parts separated by narrow fault-bounded grabens filled with siliciclastic Camaquã sediments as well as by pre-Brasiliano gneissic rocks (Encantadas Complex; Porcher and Fernandes, 1990; Remus et al., 1990; Tommasi et al., 1994), which are exposed in the cores of large-scale antiforms (Jost and Bitencourt, 1980; Soliani, 1986; Fig. 4). The Encantadas Complex comprises 2.4- to 2.2-Ga (Hartmann et al., 2000b) dioritic, tonalitic, and granodioritic gneisses (Encantadas gneisses), mylonitized syenogranites and monzogranites, and lens-shaped amphibolites on a scale of tens to hundreds of meters. The Encantadas Complex formed during the Paleoproterozoic Trans-Amazonian orogenic cycle. In the central parts of the belt (Santana “dome” in Fig. 4), two sequences can be distinguished within the overlying volcanosedimentary Poron-
gos succession. The eastern sequence, exposed in the eastern part of the Porongos belt, contains pelitic schists, graphite schists, quartzites, and marble lenses as well as acid metavolcanic rocks. The western sequence consists of metapsammite, metapelite, and marble intercalated with felsic tuffaceous rocks and minor ultramafics (Jost and Bitencourt, 1980; Remus et al., 1987, 1991; Porcher and Fernandes, 1990; Saalmann et al., 2005c). The Porongos belt is limited to the east by the Dorsal de Canguçu Shear Zone (Tommasi et al., 1994; Koester et al., 1997; Fernandes and Koester, 1999; Figs. 2C and 4). This major southwest–northeast-oriented, left-lateral ductile shear zone separates the Porongos belt from the Dom Feliciano belt in the east. It is intruded by a number of synkinematic granites. The Dom Feliciano belt (Pelotas Batholith), representing the easternmost tectonic unit in Rio Grande do Sul, is characterized by extensive Neoproterozoic crustal reworking of Trans-Amazonian basement gneisses (Babinski et al., 1996, 1997; da Silva et al., 1999, 2000b; Hartmann et al., 2000a). Comparable granitegneiss belts occur farther north in the State Santa Catarina (Florianópolis Batholith, Fig. 2B) and in Uruguay (Fig. 2C). However, the basement of both the Florianópolis Batholith and the Pelotas Batholith comprises Archean and Palaeoproterozoic units with
The assembly of West Gondwana—The view from the Rio de la Plata craton
5
Phanerozoic cover sediments
Late to post-Brasiliano sedimentary rocks Camaquã basin sediments and volcanics a
PB = Piquiri Basin; ABB = Arroio Boicí Basin; GB = Guaritas Basin
30°30´S
Granitoids Dom Feliciano event
52°45´W
sheared granites of DCSZ
Capané gneiss
PB
deformed granites
Porongos Complex quartzite lenses
marble lenses GB
acid + intermediate metavolcanics b
metapelites, -tuff(ite)s, -volcanics
Trans-Amazonian basement (Encantadas Complex)
31°00´S
mylonitic syeno-/ monzogranites Encantadas gneisses c
N DCSZ
ABB
10 km
DCSZ = Dorsal de Canguçu Shear Zone
antiform axis a Capané antiform b Santana “dome” 31°30´S
c Serra do Godinho antiform
53°15´W
mylonitic rocks fault 53°30´W
Figure 4. Geological map of the Porongos belt (modified after Chemale, 2000).
6
Saalmann et al.
clear absence of younger rocks (Babinski et al., 1997; da Silva et al., 2000b), whereas ca. 1000-Ma U-Pb zircon ages have been reported from the Punta del Este Terrane in Uruguay (Preciozzi et al., 1999; Basei et al., 2000, 2005). Hence, bearing in mind the tectonostratigraphic complexity, poorly constrained age data for some units (e.g., eastern Uruguay), and unresolved basement relationships, the Dom Feliciano belt sensu lato does not represent a coherent unit. Therefore, in this article, the term Dom Feliciano belt is restricted to the belt in eastern Rio Grande do Sul (Pelotas Batholith), and its description is likewise restricted to this area. The Porongos and Dom Feliciano belts in the southern Brazilian Shield share the same basement, Encantadas Complex (Leite et al., 2000), and thus are part of the same craton or microcontinent. This block is named the Encantadas block or Encantadas microcontinent (Chemale, 2000). The origin of this block is not resolved. It could represent a continental fragment, which either split off from the Rio de la Plata craton in late Palaeo- to Meso- or early Neoproterozoic times or could originally have been part of the Congo or Kalahari cratons. Three geophysical domains (Fig. 2D) have been distinguished in Rio Grande do Sul based on magnetic and gravimetric data (Costa, 1998). They are separated by magnetic and/or gravimetric anomalies. The Caçapava do Sul magnetic anomaly represents the border between the São Gabriel block and the Porongos belt and thus separates two distinct terranes of different structural and geochemical affinities (see below). The Porto Alegre magnetic anomaly has been interpreted as a suture zone (Costa, 1998). However, because the basement units are similar on both sides of the anomaly, this structure represents an intracontinental feature rather than a suture between two terranes. The Brasiliano orogenic cycle comprises three major tectonic events (Hartmann et al., 1999, 2000a): (1) onset of subduction activity is marked by the ca. 880-Ma Passinho diorite (Leite et al., 1998), representing the oldest Neoproterozoic tectonic event in southern Brazil (Passinho event); (2) the 750- to 700Ma São Gabriel event represents the development of a juvenile magmatic arc in the São Gabriel block (Cambaí Complex); and
(3) the ca. 630- to 600-Ma Dom Feliciano event represents extensive melting of older crust in the Dom Feliciano belt followed by widespread intrusion of voluminous post-tectonic granites. AGES, GEOCHEMISTRY, AND TECTONIC SETTING OF THE TECTONIC BLOCKS São Gabriel Block The calc-alkaline, deformed diorites, tonalites, and trondhjemites of the Cambaí Complex have zircon U-Pb conventional and sensitive high-resolution ion microprobe (SHRIMP) ages of ~750–700 Ma (Babinski et al., 1996; Leite et al., 1998). They show positive εNd(t) values (Babinski et al., 1996; Chemale, 2000; Saalmann et al., 2005a) and mark the development of a juvenile magmatic arc (São Gabriel event). The upper Palma Group, consisting mainly of andesitic to dacitic metavolcanic and volcanoclastic rocks, has been interpreted as the volcanic part of this magmatic arc (Koppe and Hartmann, 1988; Chemale et al., 1995; Hartmann et al., 1999). This interpretation is corroborated by 753 ± 2-Ma and 757 ± 17-Ma zircon U-Pb crystallization ages of metadacites (Machado et al., 1990; Remus et al., 1999). The age determination of the metasedimentary and (ultra-) mafic metavolcanic rocks of the lower Palma Group has long been ambiguous. They may be either relics of Paleoproterozoic greenstone belts (Hartmann and Nardi, 1983; Jost and Hartmann, 1984; Koppe and Hartmann, 1988; Remus et al., 1993, 1999; Hartmann et al., 1999; Hartmann and Remus, 2000) or Neoproterozoic oceanic crust (Wildner, 1990; Fragoso-César, 1991; Fernandes et al., 1992; Strieder et al., 2000). However, zircon ages (Machado et al., 1990) and recent Sm-Nd analyses that yield depleted mantle (TDM) model ages ranging between ca. 1.3 and 0.6 Ga (Chemale, 2000; Saalmann et al., 2005a) clearly indicate a Neoproterozoic age for these successions (Fig. 5). Juvenile (Meso- to) Neoproterozoic rocks in the São Gabriel block, therefore, comprise both calc-alkaline magmatic arc plutonic rocks (Cambaí Complex) as well as associated mafic metavolcanic and interleaved
Figure 5. (A) Nd data and TDM model ages for the Porongos belt and São Gabriel block (data from Saalmann et al., 2005a, 2006). Left: εNd(t) vs. time diagram. Nd data of basement units are added (Santa Maria Chico granulites: Mantovani et al., 1987; Encantadas gneiss recalculated from da Silva et al., 1999). The Santa Maria Chico granulites are exposed in the Taquarembó block (see Fig. 3) of the Rio de la Plata craton. Right: Frequency histogram of TDM model ages. The two age groups in the Porongos belt distinguish an eastern and a western sequence (Encantadas Complex TDM model ages from Chemale, 2000; Dom Feliciano belt TDM model ages from Frantz et al., 1999). DF—Dom Feliciano belt; G+MG—gneiss + metagabbro (Encantatas Complex). (B) Rock units in the Porongos belt. The Porongos Group can be subdivided into a western and an eastern sequence. EC—Encantadas Complex (basement of the Porongos and Dom Feliciano belts). (C) εNd(t) vs. (87Sr/86Sr)0 isotope correlation diagram (data from Saalmann et al., 2005a, 2006). Nearly all samples from the São Gabriel block plot in the upper-left quadrant demonstrating their juvenile signature; only the Santa Zélia granite (in the western São Gabriel block) plots in the “enriched” quadrant showing contribution of continental crust. In contrast, samples from the Porongos belt show high (87Sr/86Sr)0 and negative εNd(t) values and thus clearly plot in the field for continental crust. (D) Stratigraphic scheme of the São Gabriel block. The metasedimentary and metavolcanic rocks of the lower Palma Group are intruded by the 879-Ma Passinho dorite as well as by the 750- to 700-Ma Cambaí Complex (see Fig. 3). The upper Palma Group probably represents the volcanic counterpart of the Cambaí magmatic arc, so that stratigraphic boundaries are diachronous. The Santa Zélia granite intruded the western part of the Palma region during the late stages of D3. The lower successions of the Camaquã (molasses) basin are deformed in contrast to the upper parts. The Lavras granite and the Caçapava granite are post-tectonic intrusions with ages of 595 and 560–540 Ma, respectively. For locations of geographic areas and stratigraphic units, see Figures 3 (São Gabriel block) and 4 (Porongos belt).
The assembly of West Gondwana—The view from the Rio de la Plata craton +10
7
A
number
8
DEPL ETED MA
NTL E
7
CHUR
0
ic
–10
ss ei gn
Ch
an
uli
te
6
5
nt Sa
Porongos Belt 4
São Gabriel Block
εNd(t)
ca En
as ad nt
ia
ar aM
r og
3
Porongos belt São Gabriel block
E
2
W
–20 1
São Gabriel block
Porongos Porongos West East
0.5
1.0
1.5
2.0
2.5
TDM
3.0
DF
TDM model ages –30
EC
T (Ga)
Encantadas G+MG Complex
3
B
metapelite quartzite graphite schist acid-intermediate metavolcanics marble metapelite quartzite, metapsammite marble tuffitic metasedimentary rocks ultramafic rocks, serpentinite
EG
C
30 20 10
São Gabriel block
Santa Zélia granite Porongos east
0
–10 Porongos west
–20 –30 0.695
(tonalitic) gneisses amphibolite
age (Ma)
0.700
0.705
0.710
0.715
0.720
0.725
0.730
0.735
(87Sr/86Sr)t
Camaquã basin
500
L
600
Z Z C
C
D Caçapava granite (CG) post tectonic Lavras granite (L) late tectonic Santa Zélia granite (Z)
CG
700
2
εNd(t)
1
PORONGOS Group west east
0
C
São Gabriel uP event (D ) 3
lPv 900 Passínho diorite
lPs
Passínho event
undeformed deposits
Camaquã basin
deformed deposits
Cambaí gneisses (C)
upper Palma group (uP)
dioritic, tonalitic, granodioritic, trondhjemitic, gneisses
metandesite, metavolcanoclastic rocks, subordinate metapelite, metarenite
intrusive contact
lower Palma group lPv 1300
(metased. rocks lPs; metavolcanic rocks lPv)
serpentinite, amphibolite, magnesian schist; paragneiss, various metapelites, lenses of quartzite and marble
8
Saalmann et al.
metasedimentary rocks. Trace element concentrations, relative enrichment in light rare earth elements (LREE), low contents of Nb and other high-field-strength elements, and enrichment in large ion lithophile elements (LILE) of most igneous samples from both the Palma Group and the Cambaí Complex indicate an origin in a subduction zone environment (Silva-Filho and Soliani, 1987; Koppe and Hartmann, 1988; Remus, 1990; Strieder et al., 2000; Saalmann et al., 2005b). The data indicate the possible existence of two suites, an oceanic island arc and a continental arc or active continental margin. However, some ultramafic samples indicate the existence of another volcanic suite of intraplate character, possibly representing relics of oceanic island basalts (OIB). The metasedimentary rocks record slightly older TDM model ages and lower εNd(t) values, although εNd(t) values are still positive and (87Sr/86Sr)t values are low, suggesting that they were derived mainly from a young, juvenile source with only minor input from old crust. The metasedimentary rocks were derived from andesitic to mixed felsic and basic arc sources. Whereas ca. 697-Ma (Pb-Pb zircon, evaporation; Remus et al., 2001) synkinematic granites intruding the lower Palma Group show positive εNd(t) values of +5.2 (Babinski et al., 1996), the late tectonic Santa Zélia granite marks the first significant contribution of old continental crust in the São Gabriel block, which had been previously characterized by juvenile crust. The granite displays slightly negative εNd(t) values as well as high (87Sr/86Sr)t, which are higher than the values of the metasedimentary rocks (Saalmann et al., 2005a) and hence did not originate from melting of the metasedimentary rocks or by differentiation from a basaltic mantle source, but rather suggest mixing of mantle-derived melts with partial melts of old cratonic crust.
support the distinction of an eastern and a western sequence. The Sm-Nd and Sr data indicate sediment supply from Archean to Paleoproterozoic basement units and only minor contribution from younger sources. Various tectonic settings have been suggested for the Porongos belt, ranging from a passive margin (Jost and Bitencourt, 1980), a back-arc basin (Fernandes et al., 1992, 1995; Babinski et al., 1997; Hartmann et al., 1999, 2000a; Chemale, 2000) to a forearc setting (Issler, 1983). Trace element concentrations as well as isotope data of the metavolcanic and metasedimentary rocks demonstrating reworking of the pre-Brasiliano basement favor deposition on stretched continental crust. A rift setting associated either with a passive margin environment or an ensialic back-arc basin would be compatible with both the lithology and the geochemical data. Interbedded alkali-rich tholeiites (Marques et al., 1996) have an age of ca. 880 Ma (Rb-Sr; Soliani, 1986) and are interpreted to represent rift-related rocks (Frantz and Botelho, 2000; Frantz et al., 2000). Given that the 780-Ma age of volcanism dates the approximate depositional age of the Porongos sequence, the absence of zircons younger than 1998 Ma (Hartmann et al., 2004), and the lack of evidence for significant contribution of Neoproterozoic juvenile rocks to the metasedimentary and metavolcanic rocks, a passive margin or continental rift environment best fits both the deposition of shallow marine to deep shelf sediments and the stretching of continental crust leading to volcanism due to high heat flow in the thinned lithosphere, which is characterized by significant contamination by old continental crust.
Porongos Belt
The Dom Feliciano belt consists predominantly of 630- to 600-Ma granitoids. Six intrusive suites can be distinguished (Philipp and Machado, 2001, 2005), which are related to shear zone activity. Negative εNd(t) values indicate significant contribution of old crust; variations of TDM model ages between 1.5 and 2.3 Ga suggest different mixing proportions between ancient crust and mantle material (Babinski et al., 1997; Frantz et al., 1999). The granitoids contain many enclaves of mainly tonalitic composition, and tonalitic gneiss xenoliths, decimeters to several meters in scale, occur in the oldest intrusive bodies (i.e., the Pinheiro Machado suite; da Silva et al., 1999). Based on zircon U-Pb SHRIMP analyses, da Silva et al. (1999) identified two episodes of crustal partial melting in the Pinheiro Machado suite, at ca. 800 Ma and ca. 610 Ma. A tonalitic gneiss xenolith has a magmatic age of ca. 781 Ma but also shows negative εNd(t) values and a 2.24-Ga TDM model age (da Silva et al., 1999) and thus supports the occurrence of ca. 800-Ma remelting of ancient (Paleoproterozoic) crust. It has been noted that the 780-Ma melting would correspond to the emplacement of the Cambaí Complex in the São Gabriel block (da Silva et al., 1999), and Leite et al. (2000) suggest a 800-Ma collisional event with low-angle shear zones was associated with the São Gabriel event farther west. However, the Encantadas block was not connected to the São
In contrast to the rock units within São Gabriel block, the age of the Porongos sequence is not well established. The youngest zircons within quartzites of the Porongos belt, recently published by Hartmann et al. (2004), have ages of ca. 1998 Ma, providing a maximum age for the basin fill, which, therefore, post-dated the Trans-Amazonian orogeny. Approximately 780-Ma U-Pb zircon ages for metarhyolitic rocks (Chemale, 2000; Porcher et al., unpubl. in Hartmann et al., 2000b) are interpreted as magmatic ages. Basei et al. (2000), in contrast, propose that the age of the volcanism is not representative of the time of deposition of the metasedimentary units, but instead dates the metamorphic climax leading to anatexis of deeply buried sedimentary rocks. However, the volcanic rocks are intercalated with the metasedimentary rocks and show the same deformation. Moreover, rocks of probable tuffaceous origin alternate with pelitic and quartzitic schists. Hence, the ages of magmatism date syndepositional volcanic activity and thus, the approximate age of basin development. The metavolcanic and metasedimentary rocks of the Porongos sequence, especially the succession in the western part (Fig. 5), show very evolved negative initial εNd(t) values and high TDM model ages (Saalmann et al., 2006). The data
Dom Feliciano Belt
The assembly of West Gondwana—The view from the Rio de la Plata craton Gabriel block prior to 700 Ma (see below). Because the Encantadas block contains both the Dom Feliciano belt and the Porongos belt, this partial melting event could be correlated with extension and basin development in the Porongos belt at ca. 800–750 Ma. Deposition in the Porongos belt was accompanied by volcanism derived from melting of ancient crust, so that both melting events could be linked to stretching and thinning of continental crust. Hence, both belts show the same crustal reworking event but represent different levels of exposure, with deeper crustal levels being exposed in the Dom Feliciano belt while subsurface levels are preserved in the Porongos belt. However, it cannot be excluded that structures in the Dom Feliciano belt were genetically linked to events occurring in areas farther east. These regions, however, would be located on the present-day shelf regions in the Atlantic and thus, are unknown. The voluminous 630- to 600-Ma magmatism in the Dom Feliciano belt has been attributed to a magmatic arc setting above either a west-dipping subduction zone of the Adamastor Ocean (Fernandes et al., 1992, 1995) or an east-dipping subduction zone of an ocean located to the west of the Dom Feliciano belt (Chemale, 2000). The belt lacks juvenile Neoproterozoic rocks, and isotopic studies indicate extensive reworking of Paleoproterozoic crustal material in the Dom Feliciano belt (Babinski et al., 1996, 1997; da Silva et al., 1999, 2000b; Hartmann et al., 2000a), so that a magmatic arc position for this belt during the Brasiliano orogeny has been questioned (Hartmann et al., 1999, 2000a). If not a magmatic arc, then the widespread remobilization of Trans-Amazonian basement in the Dom Feliciano belt was not due to subduction but induced by other processes (e.g., plume ascent), and deformation occurred along intracontinental strike-slip shear zones. STRUCTURAL EVOLUTION Structural Evolution of the São Gabriel Block In the São Gabriel block, four deformational events related to the Brasiliano orogenic cycle, D1 to D4, can be distinguished (Fig. 6). D1 and D2 only occur in the metasediments and (ultra-) mafic metavolcanic rocks of the lower Palma Group, whereas the first deformation in the upper Palma Group, in the calc-alkaline gneisses of Cambaí Complex and synkinematic granites, correlates with D3. The layering and banding (S1) of the lower Palma Group rocks was formed during the first deformational phase D1. Isoclinal folding of S1 took place under amphibolite-facies metamorphic peak conditions during D2 associated with southeast-directed thrusting, folding, and formation of the foliation S2. The third deformation phase (D3) represents the last major ductile deformation in the São Gabriel block. In the Palma Group schists, D3 took place mainly under greenschist to lower amphibolite–facies metamorphic conditions. In these rocks, D3 led to refolding of F2 folds and of S2 on a scale of centimeters to meters. The F3 folds probably represent subordinate folds to regional fold structures.
9
The L3 lineation is a crenulation lineation (L3cr) that formed parallel to F3 fold axes. In most cases S2 still represents the dominant foliation, and the earlier L2 lineation is still preserved except in narrow noncoaxial shear zones, centimeters–decimeters to several meters in scale and oriented subparallel to S2, which overprinted the earlier fabrics and display a well developed S-planes (S3) as well as a stretching lineation (L3str). D3 is the first deformation of the upper Palma Group. It is characterized by southeastvergent folding. Cleavage-bedding intersection lineations and fold axes strike southwest-northeast and correspond to F3 and L3 in the lower Palma Group. The deformed diorites, tonalities, and trondhjemites of the Cambaí Complex intruded subparallel to the foliation planes of the country rocks (lower Palma Group), indicating a syn-D3 emplacement. This emplacement is also the case for synkinematic granites, which intrude the lower Palma Group. Syn-D3 emplacement of the synkinematic plutonic rocks took place in a southwest–northeast-oriented right-lateral shear regime. Fabrics and structures of the rocks reflect a deformational evolution, starting with magmatic flow fabrics followed by subsolidus and solid-state high- to low-temperature deformation characterized by right-lateral noncoaxial progressive deformation. The Santa Zélia granite intruded during the late stages of D3 and displays a nonpenetrative solid-state deformational overprint, which occurred at decreasing temperatures until upper greenschist-facies temperature conditions. However, the magmatic fabric is preserved to a large degree. D3 right-lateral shear characterizes the deformation during ascent, emplacement, and solidification of juvenile plutonic rocks, whereas the wallrocks were deformed predominantly by northwest-southeast contraction. Such a partitioning of the deformation in noncoaxial transcurrent shear zones with folding in coaxial strain parts suggests an overall deformational regime of dextral transpression during D3. Right-lateral ductile transpression led to formation of a southeast-verging stack of slices (Fig. 7). The last deformation phase (D4) took place under retrograde conditions and is characterized by localized semi-brittle southeast-directed thrust faulting, leading to imbrication, kinking, and thrust-related folding of the rocks. D4 fault zones locally reactivated D3 shear zones. Structural Evolution of the Porongos Belt The deformation of the Porongos sequence comprises several phases of folding (Fig. 6). The layering of the metapelites contains S-parallel quartz veins and segregations as well as isoclinal folds on a millimeter to centimeter scale and thus comprises alreadyfolded layers. The layering therefore represents the first foliation (S1), associated with a first folding episode (F1). Isoclinal folding of the layering (S1) and the quartz mobilizates are attributed to a second deformational phase (D2) and development of a foliation (S2) related to folding F2. S2 is parallel to S1. A NNE–SSWtrending mineral and stretching lineation (L2) is locally preserved
DOM FELICIANO BELT
metam. peak
retrograde metam.
F3, L3
synkinematic magmatism
S3, L3, F3 brittle-ductile
D4 Open folding, top-SE/E thrusting
D3
DCSZ 670–630 Ma
retrograde metam.
D5b
SSW-NNE to WSW-ENE sinistral brittle shear
D3
brittle strikeslip faulting, Camaquã pull-apart basins, associated normal and top-SE and top-NW oblique brittle reverse faulting; folding is associated with and restricted to faults
localized strike-slip faulting
Camaquã lower deformed sequences 630–580 Ma
brittle/ductile
open regional-scale F5 folding, uplift of basement units in the cores of antiforms leading to folding and rotation of the previous structures; semibrittle normal faulting adjacent to the uplifting gneisses
NW-verging F4 folding, SW-NE axes, dm to hundrets of m-scale, NW-directed thrusting, crenulation cleavage S4, SW-NE trending L4 lineation parallel to F4 fold axes; late-D4 semiductile C4-shear bands
Partitioning of strain in (1) localized major dextral SW-NE strike-slip shear zones and (2) contractional zones displaying folding; intrafolial folds in orthogneiss, up-right oblique folds in dm-m-scale, possibly regional in scale; L3 = SW-NE; S3//S2
Syn- to late tectonic granites SGB 690–?660 Ma
prograde metamorphism
D5a
D4
transitional
SW-NE striking, ductile sinistral strikeslip shear zones
“tangential” shear zones, associated with ?thrusting?
?
D2
D1
Figure 6. Overview and correlation of deformational phases in the Dom Feliciano belt, Porongos belt, and São Gabriel block. Time markers such as synkinematic intrusions provide constraints on the age of distinct deformational phases.
prograde metam.
L2
D2
Top-SE to topESE directed shearing,F2 isoclinal folding, S2 is parallel to S1, L2=NW -SE;
D1
F2
F3 closed to locally isoclinal folding (cm- to dm-scale), refolding of F1 and F2 ; SW-NE striking S 3 foliation parallel to F 3 axial planes top-SW directed dextral sense of shear
F2 isoclinal folding (mm- to cm-scale) of S1 and the quartz veins, S2 foliation is parallel to S1, SSW-NNE trending mineral and stretching lineation L2, top-NNE directed thrusting
main foliation and layering S1; quartz veins (mobilizations) parallel to the first cleavage S1 (only in metapelites)
L2
D3
D2
D1
Cambaí gneisses 750–700 Ma
Shearing, F1 folding ?, quartz veins/ mobilizations in schists, S1 is parallel to S0
Passinho diorite 880 Ma
PORONGOS BELT
SÃO GABRIEL BLOCK
10 Saalmann et al.
The assembly of West Gondwana—The view from the Rio de la Plata craton
11
S4
W
E
underthrust basement of the Encantadas block ?
Encantadas block São Gabriel block
Taquarembó block (Rio de la Plata craton)
Guaritas subbasin (Camaquã basin)
Porongos belt
Dorsal de Canguçu Shear Zone
Caçapava granite
Rio de la Plata craton
Palma Group
Encantadas complex
Dom Feliciano belt
Piquiri and Aroio Boicí subbasins (Camaquã basin) Neoproterozoic-Cambrian sedimentary rocks
Granite
Figure 7. Schematic cross-section across the Southern Brazilian Shield from the eastern margin of the Rio de la Plata craton in the west to the Dom Feliciano belt in the east.
on the foliation planes. Mylonitic layers suggest occurrence of localized ductile shear zones during this phase. The sense of shear cannot be deduced with certainty; however, relics of kinematic indicators suggest a top-to-the-NNE directed sense of shear. D3 led to close to isoclinal refolding of F2 folds on a centimeter to decimeter scale around gently southwest-plunging fold axes. Folding (F3) was accompanied by a northeast-southwest directed dextral sense of shear, which seems to have been localized in narrow shear zones. Low to middle greenschist–facies peak metamorphic conditions were reached during D2 and D3. In the west, the metamorphic grade increases to upper greenschist–facies metamorphic conditions with temperatures exceeding 400 °C. Metamorphic zoning within the Porongos belt from east to west from the chlorite zone to the garnet–staurolite zone and findings of kyanite (Jost, 1982; Remus et al., 1991) show that conditions reached amphibolite-facies conditions and higher pressures. The fourth deformation (D4) took place under retrograde conditions and is represented by open to close chevron-style folding on a scale of decimeters to hundreds of meters. Northwestvergent F4 folds in outcrop scale are subordinate folds to major folds on a scale of hundreds of meters to kilometers. Folding was associated with northwest-directed thrusting, leading to nappe emplacement and thrust stacking (Fig. 7) and northwestward transport of the southeastern units of the Porongos sequence onto the northwestern parts. At least two major thrust units can be inferred, characterizing the overall geometry of the Porongos
belt. Recumbent and northwest-vergent folds on a scale of several hundred meters, related to possible nappe transport, have also been reported by Remus et al. (1987). The thrust-stack was cut by semi-brittle to brittle faults during D5 that developed within a sinistral transcurrent shear regime. Fault-bounded pull-apart basins with a narrow, elongate northeast–southwest-oriented shape formed in transtensional segments (e.g., Piquiri basin; Fig. 4). They are filled with relatively thick unmetamorphosed sediments, which represent the first deposits of the Camaquã basin. These sequences are also affected by faulting and folding. Structural Evolution of the Dom Feliciano Belt Three deformational events (D1 to D3), the first two characterized by ductile and the last by brittle tectonics, can be distinguished in the Dom Feliciano belt (Fernandes et al., 1992; Philipp et al., 1993; Fig. 6). D1 is restricted to the Pinheiro Machado Suite, which represents the oldest intrusive suite in the Dom Feliciano belt (ca. 625–605 Ma; Babinski et al., 1997). So-called “tangential” shear zones (i.e., flat-lying ductile shear zones with oblique, subhorizontal, west–east- to northwest–southeast-striking lineations) characterize D1. Gently west-plunging stretching lineations indicate ESE-directed tectonic transport (Fernandes et al., 1992; Frantz et al., 1999; Philipp and Machado, 2001); however, the
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Saalmann et al.
original dip and plunge of the tectonic elements might have been modified during subsequent deformation. D2 can be observed in all intrusive suites and is represented by large WSW–ENE- to southwest–northeast-trending, ductile leftlateral strike-slip shear zones, which extend over tens to hundreds of kilometers and have a width of up to hundreds of meters (Fernandes et al., 1992; Nardi and Frantz, 1995; Frantz et al., 1999; Philipp and Machado, 2001). This event affected various granitic suites, including ca. 595-Ma granites (Babinski et al., 1997). D3 is characterized by SSW–NNE- to WSW–ENE-trending sinistral brittle-ductile shear zones, which cut and displace D2 structures (Philipp and Machado, 2001). Correlation of Brasiliano Deformational Events in the Southern Brazilian Shield The age of events and comparison of the structural evolution of the various tectonostratigraphic blocks can be used to reconstruct the timing of their juxtaposition. The age of D3 in the São Gabriel block, representing the São Gabriel event, is well constrained by 750- to 700-Ma U-Pb SHRIMP ages for the synkinematic plutons of the Cambaí Complex (Babinski et al., 1996; Leite et al., 1998). The age of D1 and D2 in the São Gabriel block can only be deduced from the fact that the Palma Group rocks, which were affected by these events, formed in a Neoproterozoic arc environment and that D1 and D2 predate the 750- to 700-Ma São Gabriel event. According to the TDM model ages of the (ultra-)mafic metavolcanic rocks of the lower Palma Group, D1 and D2 in the São Gabriel block occurred between 900 and 750 Ma. This range is compatible with the 880-Ma Passinho diorite recording the first subduction activity and accretion in this area. The age of D4 in the São Gabriel block can only be estimated. Steep dips of D4 thrust faults exposed in metasedimentary rocks at the western margin of the Caçapava granite can be explained by rotation of these faults during emplacement of the granite in a dextral shear environment. Therefore, D4 occurred prior to granite intrusion, which has an age of ca. 562–540 Ma (Remus et al., 2000a). Deformation in the Porongos belt post-dates the ca. 780Ma age of deposition. Amphibolite-facies metamorphism during D2 and D3 can very likely be attributed to subduction and collision. D4 northwest-directed thrusting and nappe stacking could, however, be linked to southwest–northeast-oriented sinistral shear at the Dorsal de Canguçu Shear Zone and therefore postdates subduction and the main collision. The northwest-vergent thrust-sheets form part of a flower structure or a transpressive thrust-stack formed when ongoing convergence was accommodated by orogen-parallel transcurrent shearing during the final stages of collision. Emplacement of synkinematic granites occurred in transtensional segments of the shear zone (Fernandes and Koester, 1999), transpressional parts having been affected by northwest-southeast shortening, which would be compatible with northwest-vergent folding and thrusting during D4 in
the neighboring Porongos belt. In this case, the age of Dorsal de Canguçu Shear Zone activity and D4 in the Porongos belt is given by the synkinematic granites, which display ages between 670 and 620 Ma (Koester et al., 1997). Therefore, D1 to D3 deformation occurred between deposition and activation of the Dorsal de Canguçu Shear Zone, and thus very likely at ca. 750–700 Ma. D5 brittle block tectonics and strike-slip faulting affected also the lower sequences of the Camaquã basin. Syntectonic deposition occurred between 630 and 600 Ma within the eastern subbasins, whereas the basin fill of a sub-basin in the northwestern Porongos belt has an age of ca. 592–580 Ma (Paim et al., 2000). Consequently, D5 brittle strike-slip in the Porongos belt occurred from 630 Ma until at least 580 Ma. Left-lateral strike-slip shear zones of D2 in the Dom Feliciano belt can be linked to D5 sinistral strike-slip in the Porongos belt. Ductile shear in the Dom Feliciano belt, in contrast to brittle faulting in the Porongos belt farther west, is attributed to the deeper crustal levels exposed in the Dom Feliciano belt. The scheme in Figure 6 gives an overview of the deformation events within the tectonostratigraphic units and possible correlation of distinct phases. D1 and D2 in the São Gabriel block represent the oldest Brasiliano deformational events in the Southern Brazilian Shield. D1 to D3 in the Porongos belt are contemporaneous to D3 in the São Gabriel block (750–700 Ma). The kinematics of the deformations are compatible as well. Southwest–northeast-trending F3 fold axes and indications for southwest–northeast-directed dextral shear in the Porongos belt are compatible to southwestnortheast dextral transpression and southeast-directed oblique thrusting and folding in the São Gabriel block at this time. The final stage of São Gabriel event represents the collision of the Rio de la Plata craton with the Encantadas block occurring at, or shortly after, 700 Ma. D4 in the São Gabriel block either immediately followed D3 ductile transpression and thrust-stacking and represents the final stages of southeast-vergent nappe stacking, or it occurred some 20 m.y. later in response to D4-related northwest-directed thrusting and folding in the Porongos belt, which began at ca. 670 Ma. D5 strike-slip in the Porongos belt occurred from 630 Ma until at least 580 Ma. The correlative D2 sinistral strike-slip event in the Dom Feliciano belt affected most of the granites in this belt and thus must have occurred at ca. 620–590 Ma. Brittle D3 sinistral faults in the Dom Feliciano belt represent the waning stages of collision and deformation during uplift to lower crustal levels. They may be correlative to continued faulting further west (Porongos belt, Camaquã basins). PLATE TECTONIC MODEL FOR SOUTHERN BRAZIL Previous Models A number of tectonic models have been suggested for the Brasiliano orogen in southern Brazil. Most models propose west-dipping subduction of the Adamastor Ocean (Hartnady et al., 1985), located between the Kalahari and Rio de la Plata
The assembly of West Gondwana—The view from the Rio de la Plata craton cratons, beneath the Dom Feliciano belt, with the latter representing a Neoproterozoic magmatic arc (Soliani, 1986; Tommasi and Fernandes, 1990; Fernandes et al., 1992). Fragoso-César (1991) proposed the existence of two oceanic basins, the Charrua Ocean in the west and the Adamastor Ocean in the east. Fernandes et al. (1992) suggested that west-dipping subduction generated, in sequence, two magmatic arcs, first the Dom Felciano belt (800– 750 Ma), followed by a 750- to 650-Ma active continental margin forming the present-day São Gabriel block. Age determinations, however, show that magmatism in the western parts (São Gabriel block) started earlier than in the east (Dom Feliciano belt). This scenario is considered by another model (Chemale, 2000), starting with subduction of the Charrua Ocean and formation of an intraoceanic arc (Vila Nova belt). Closure of the Charrua Ocean by eastward subduction below the Encantadas microcontinent was associated with folding and thrusting while the Adamastor Ocean opened on eastern side. The Dom Feliciano magmatic arc formed as result of subduction of the Adamastor Ocean toward the west, leading finally to the collision of the Kalahari craton and the Encantadas microcontinent. Isotopic studies show extensive reworking of Paleoproterozoic crustal material in the Dom Feliciano belt (Babinski et al., 1996, 1997; da Silva et al., 1999, 2000b; Hartmann et al., 2000a), and a magmatic arc position for this belt during the Brasiliano orogeny has been debated. Juvenile rocks have been found only in the São Gabriel block (Babinski et al., 1996; Leite et al., 1998). Based on these studies, Hartmann et al. (1999) suggested a longlived eastward subduction beneath a continental mass comprising both the Rio de la Plata craton and the Encantadas gneisses, comparable with the modern Andes. A similar model has been proposed by Ramos (1988). These models, however, fail to explain the ca. ~620- to 600-Ma Dom Feliciano event in the east, far away from the São Gabriel subduction zone environment. They also do not take into account eastward-decreasing ages in magmatism and deformation, which cannot be explained by westward accretion caused by subduction to the east. Since the first model was proposed, many new isotopic and geochemical data have been published, providing an improved basis for modeling the tectonic evolution of southernmost Brazil. A tectonic model has to take into account the following important findings: • Existence of two 880-Ma and 780- to 700-Ma juvenile volcanic arcs (Passinho and Vila Nova) in the São Gabriel block; • Only subordinate contribution from older rocks to the magmatic and metasedimentary rocks in the São Gabriel block; • A ca. 800- to 700-Ma depositional age of the Porongos Group (based on the 780-Ma age of volcanism); • Lack of any signs of input of Neoproterozoic juvenile rocks to Paleoproterozoic and Archean source rocks for the Porongos succession (hence, the Porongos belt and São Gabriel block were separated and did not form a volcanic arc/back-arc basin pair);
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• Absence of reliable evidence for Neoproterozoic volcanic arc assemblages and other subduction-related rocks in the Porongos belt (as would be implied by eastward subduction below the Encantadas microcontinent); • Tectonic settings of the tectonometamorphic blocks (subduction environment versus passive margin); • Onset of tectonic events in the western areas, already starting by 880 Ma (Passinho arc), much earlier than in the east; • Migration of the deformation and tectonometamorphic events toward the east with time, indicating eastward progressive accretion and collision; • Structural evolution and style of deformation of each block; and • Sharing of the same basement (Encantadas Complex), characterized by Paleoproterozoic units, by the Porongos and Dom Feliciano belts. Proposed Model Two Brasiliano tectonostratigraphic terranes are exposed in the schist belts to the east of the Rio de la Plata craton in southernmost Brazil. The Porongos belt is located on the passive margin of the Encantadas microcontinent, whereas two magmatic arc assemblages, an intraoceanic arc (Passinho arc), and an active continental margin setting (Vila Nova arc) are preserved in the São Gabriel block. The Brasiliano plate tectonic evolution of the Southern Brazilian Shield starts with development of the ca. 879-Ma intraoceanic Passinho arc (Leite et al., 1998) in response to subduction of oceanic crust. The precise age of the oceanic basin, which opened to the east of the Rio de la Plata craton, is unresolved; however, the oceanic lithosphere that was consumed must have been generated some 100 m.y. earlier. This sequence is in accordance with 0.9- to 1.2-Ga TDM model ages of ultramafic rocks of the Palma Group (Saalmann et al., 2005a). The Passinho arc formed above an east-dipping subduction zone (Figs. 8A and 9) that led to its accretion to the passive margin of the Rio de la Plata craton (Figs. 8B and 9B). Between 850 and 700 Ma, subduction occurred to the west beneath the continental margin consisting of the Rio de la Plata craton and the attached Passinho island arc (Figs. 8B and 9B), giving rise to development of the calc-alkaline plutonic rocks of the Cambaí Complex as well as the volcanic arc of the upper Palma Group. Sedimentary input into the back-arc and forearc basins associated with this active continental margin was respectively derived mainly from the previously accreted Neoproterozoic juvenile Passinho arc or from rocks of the magmatic arc, with only small amounts of sedimentary input from the old Rio de la Plata craton in the hinterland. This input source explains both the positive εNd(t) values and low TDM model ages (1.1– 0.8 Ga) of the metasedimentary rocks as well as the juvenile signatures of the Cambaí Complex caused by the absence of significant contribution from old crust to the melts. Westward
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Saalmann et al.
W
Rio de la Plata craton
Passinho intraoceanic arc
E São Gabriel / Goiás Ocean
A 0.9–0.85 Ga
Encantadas block
São Gabriel block (D1–D3) Rio de la Plata craton
Passinho Vila Nova arc arc Palma group
Porongos sequence
Dom Feliciano belt tonalites
B 0.8–0.7 Ga
Cambaí complex (εNd(t) > 0)
Encantadas block Porongos belt Dom Feliciano (D1–D3) belt
Juvenile São Gabriel block (late D3)
Rio de la Plata craton
... ? ...
C
0.7–0.67 Ga
late tectonic granites (εNd(t) < 0)
future DCSZ Encantadas block
Rio de la Plata craton
Juvenile São Gabriel block (D4)
Porongos belt (D4) DCSZ
Kalahari craton
Dom Feliciano belt (D1)
Adamastor Ocean
Lower Gariep basin
?
D 0.67–0.62 Ga
“tangential” granites
?
possible terrane(s) on present-day shelf
Encantadas block Rio de la Plata craton
Juvenile São Gabriel block (D4)
Porongos belt (D5)
Camaquã basin
Camaquã DCSZ basin
Kalahari craton
Dom Feliciano belt (D2–D3)
Upper Gariep + Rocha basin arc?
E
? ?
0.62–0.54 Ga
post-tectonic granites
Caçapava granite Camaquã sediments
transcurrent granites post-tectonic granites
Figure 8. Cartoons depicting the plate tectonic evolution of southern Brazil during the Brasiliano orogenic cycle. Note that in this model the term Dom Feliciano belt is restricted to the Pelotas Batholith. DCSZ—Dorsal de Canguçu Shear Zone.
The assembly of West Gondwana—The view from the Rio de la Plata craton
A
B
0.9–0.8 Ga
15
0.8–0.7 Ga
A C C
GO/SGO P
GO/SGO L
P
E
RP SGB Pa
AO
RP K
Pa
C
A ANS ANT AUS C E IND L K RP P SC WA GO/SGO AO MO Pa SGB
0.7–0.6 Ga
A
C
P
L
SGB RP Pa
AO K
DF E
0.59–0.53 Ga SC
K
Amazon craton Arabian-Nubian Shield Antarctica Australia Congo-São Francisco craton Encantadas block India Luís Alves block Kalahari craton Rio de la Plata craton Paraná block Sahara craton West Africa craton Goiás/São Gabriel Ocean Adamastor Ocean Mozambique Ocean Passínho arc São Gabriel block
Subduction zone with magmatic arc
ANS
S ut
ur e
WA
East Afr ican Orogen
D
L E AO E
a mb ique
C A
RP SGB S GB
Moz
Pa
IND
L K
Pa
E
MO
ANT
AUS
Figure 9. Schematic reconstruction of inferred plate movements and collisional events in South America and southwest Africa between 0.9 and 0.53 Ga (modified and combined from Brito-Neves et al., 1999, and Pimentel et al., 1999, and supplemented with own data; East Gondwana after Boger and Miller, 2004). The dotted lines in the last sketch show the present-day outlines of the continents. SC—Sahara craton; WA—West Africa craton.
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Saalmann et al.
subduction caused southeastward thrusting and stacking of the metasediments and juxtaposed (ultra-)mafic volcanic rocks in an accretionary wedge. Dextral shear indicates oblique subduction. At the same time, between 800 and 700 Ma, the Porongos basin developed on stretched and thinned continental crust of the Encantadas microcontinent forming a passive margin. The collision of the Rio de la Plata craton with the Encantadas block occurred during the final stages of São Gabriel event at or shortly after 700 Ma (Figs. 8C and 9C), causing southeastvergent thrusting in the São Gabriel block and isoclinal folding, shearing, and amphibolite-facies metamorphism (kyanite, garnet, staurolite) in the Porongos belt located on the lower plate (D1 to D3). Shortening was in parts accommodated by lateral shear. The late tectonic Santa Zélia granite in southernmost Brazil records contributions from old continental crust in the juvenile block for the first time, as indicated by its slightly negative εNd(t) values (Saalmann et al., 2005a). This input may be attributed to melting of underthrust continental crust of the subducting lower plate. Late- to post-collisional ongoing convergence was accommodated by orogen-parallel transcurrent shearing: sinistral shear along the Dorsal de Canguçu Shear Zone and northwest-directed transpressive D4 nappe stacking in the Porongos belt (Fig. 8D), which started at ca. 670 Ma and ceased at ca. 620 Ma (Koester et al., 1997). In the Porongos belt, sinistral ductile shear was followed by brittle sinistral strike-slip faulting while 630- to 600-Ma ductile sinistral strike-slip shear zones and synkinematic granite intrusions were still occurring in the Dom Feliciano belt farther to the east (Frantz and Nardi, 1992; Philipp et al., 1993; Babinski et al., 1997; Leite et al., 2000; Fig. 8E). The eastern Camaquã subbasins formed as pull-apart basins due to transtension. Transition from transcurrent to extensional tectonics is recorded in the Dom Feliciano belt between 630 and 617 Ma (Frantz et al., 2000). Assuming a magmatic arc environment for the Dom Feliciano belt, the Encantadas block could represent a single microcontinent sandwiched between the Rio de la Plata and the Kalahari cratons in response to the closure of two oceans (Adamastor and São Gabriel/Goiás oceans). In this case, the Dom Feliciano belt formed due to the subduction of the Adamastor Ocean. This scenario is also presented in Figures 8D and 9C. The Encantadas block, forming the basement of the Dom Feliciano belt, was either separated from Africa (e.g., the Congo craton) and subsequently attached to the Rio de la Plata craton or it was previously part of the Rio de la Plata craton and became separated from it by opening of the São Gabriel/Goiás Ocean. Derivation from the Congo craton could be indicated by Paleoproterozoic gneisses in the Dom Feliciano belt that may be correlative with the Epupa Complex north and south of the Kaoko belt of southwestern Africa (Leite et al., 2000). Alternatively, the Encantadas block may already have been attached to the Kalahari craton without formation of Neoproterozoic oceanic crust between these two continental blocks. This would imply that the widespread remobilization
of Paleoproterozoic basement in the Dom Feliciano belt was not caused by subduction of oceanic lithosphere but induced by other processes (e.g., plume ascent or crustal thinning due to extension) and that deformation occurred along intracontinental strike-slip shear zones. This explanation is supported by the lack of ophiolites in the Southern Brazilian Shield along the Atlantic Ocean coast. In Figure 8 the subduction below the Dom Feliciano belt is therefore labeled with a question mark. However, the absence of Grenville rocks in the Encantadas Complex argues against a connection with the Kalahari craton because the 1.2- to 1.0-Ga Namaquã belt represents the basement of the Gariep belt. The absence of juvenile rocks and the predominance of reworked ancient crust observed in many belts in West Gondwana have been explained by relative proximity of colliding cratons, which were separated only by a narrow ocean (Cordani et al., 2003). Hence, the Adamastor Ocean separating the Encantadas block (as well as other cratons in South America) from the Kalahari and Congo cratons was only a narrow seaway in contrast to the broad Goiás/São Gabriel Ocean farther west. In another possible model, the Encantadas block was part of the Rio de la Plata craton and was separated from it during opening of the São Gabriel/Goiás Ocean. It might have been split off the craton and displaced parallel to it along major strike-slip shear zones comparable to the terranes in the North American cordillera (e.g., Nokleberg et al., 2000). The model of Basei et al. (2005) is partly adopted by interpreting the Rocha Group in Uruguay being an equivalent of the Oranjemund Formation of the upper Gariep basin (Fig. 8E). This equivalence is in accordance to the Grenvillian Kalahari basement found in the Punta del Este terrane in eastern Uruguay and detrital zircon ages (Basei et al., 2005) near 1.0 Ga, which indicate a provenance relationship with the Gariep orogenic belt in southwest Africa. However, the back-arc position of this basin behind a magmatic arc—a result of eastward subduction of the Adamastor Ocean below the Kalahari craton—remains unresolved and poorly constrained. According to Basei et al. (2005) the whole Dom Feliciano belt sensu lato (i.e., the Aiguá, Pelotas, and Florianópolis batholiths) formed a continuous belt at the margin of the Kalahari craton. However, as pointed out above, the Paleoproterozoic basement units of the Pelotas and Florianópolis batholiths are incompatible with 1.2- to 1.0-Ga basement of the Kalahari craton and the Punta del Este terrane, and hence, the model cannot be applied to these units. This incompatibility further indicates that the Dom Feliciano belt sensu lato does not represent a single coherent belt. Post-tectonic granites display ages of 600–580 Ma in both the Dom Feliciano belt (Frantz et al., 1999) and in the western São Gabriel block and the Taquarembó block (e.g., Remus et al., 1999, 2000b). Strike-slip shearing, however, continued in localized fault zones and lasted at least until 540 Ma, as recorded by fault-related granites like the Caçapava (562 Ma) and São Sepe granites (540 Ma) (Chemale, 2000; Remus et al., 2000a), as well as by the 540- to 530-Ma 40Ar/39Ar data of biotite from late tectonic shear zones in the Dom Feliciano belt.
The assembly of West Gondwana—The view from the Rio de la Plata craton LINKS TO GONDWANA ASSEMBLY Correlation with Other Brasiliano Belts In Brazil, Neoproterozoic juvenile rocks related to subduction of oceanic crust are not restricted to the São Gabriel block but also occur in the Ribeira belt (790 Ma and 635- to 620-Ma arc complexes in the Costeiro domain, Heilbron and Machado, 2003; ca. 630 Ma Pirapora do Bom Jesus ophiolitic complex, Tassinari et al., 2001) and in the western Brasília belt (Pimentel and Fuck, 1992; Junges et al., 2002; Laux et al., 2005; Fig. 10). The Brasília belt shows an accretionary history comprising the formation of
17
890- to 800-Ma intraoceanic arcs followed by development of a ca. 750-Ma active continental margin (Pimentel and Fuck, 1992; Junges et al., 2002) that resembles the early evolution in the São Gabriel block. In both areas subduction of oceanic crust started at ca. 0.9 Ga, and final ocean closure at ca. 0.63–0.60 Ga in the Southern Brasília belt (Brito Neves et al., 1999; Pimentel et al., 1999; Laux et al., 2005) can be correlated with the Dom Feliciano event in southern Brazil. This correlation implies the existence of a large ocean (Goiás Ocean) between the Amazonian and Congo–São Francisco cratons (Kröner and Cordani, 2003), which may have been connected to the oceanic basin in the São Gabriel farther south. In this case, a large oceanic realm that
N
500 km
A SL 4°S
BP
Ar
TBL
A
RPB 630– 550
CD
900– 700 CG
TBL
630– 610 Br
RA
SFC
T 630– 610
28°S
900– 700
LP SGB 54°W
Por
A 590–560
RNE 590–560
PR
PA
S
Br
P 630– 550
16°S
630–580
R
630–610 + 560–530
F 630–610
DF 630–610 46°W
Brasiliano belts Shield areas Neoproterozoic juvenile rocks
Figure 10. Distribution of cratons and Brasiliano belts in Brazil (modified from Alkmim et al., 2001, Cordani et al., 2003, with data from Caby et al., 1995, and data cited in da Silva et al., 2005). Heavy lines—major lineaments; broken heavy lines—assumed continuation of lineaments; dotted lines of cratons—uncertain craton boundaries. Cratons: A—Amazon; CG—Central Goiás; LP—Rio de la Plata; PA—Pampia; PR—Paraná; RA—Rio Apa; SFC—São Francisco–Congo; SL—São Luís. Belts and other features: A—Araçuaí belt; Ar—Araguaia belt; BP—Boroborema province; Br—Brasília belt; CD—Chapada Diamantina; DF—Dom Feliciano belt; F—Florianópolis Batholith; P—Paraguai belt; Por—Porongos belt; R—Ribeira belt; RNE—northeast branch Ribeira belt; RPB— Riaco do Pontal belt; S—Sergipano belt; SGB—São Gabriel block; T—Tebicuary river area; TBL—Trans-Brasiliano lineament.
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formed prior to 900 Ma would have separated Amazonia and Rio de la Plata from the Congo–São Francisco and Kalahari cratons. The beginning of the opening of the ocean is unclear. A correlation of Mesoproterozoic calk-alkaline dikes with igneous rocks within the Namaqua fold belt in southwestern Africa implies that the Rio de la Plata craton and southwestern Africa were still contiguous in Early-Middle Proterozoic times (Iacumin et al., 2001). In contrast to the South American belts, the Namaquã fold belt was affected by the 1.2- to 1.0-Ga Kibaran orogeny and, therefore, the Rio de la Plata craton must have separated from the south African sector prior to 1.2 Ga. The ca. 1.7-Ga unmetamorphosed dike swarms in the Rio de la Plata craton in Uruguay (Teixeira et al., 1999) and the 1.6-Ga tholeiitic dike swarms in central-eastern Argentina, which formed during extensional tectonics, could reflect the initial rifting and break-up stages, so that formation of a large oceanic basin to the east of the Rio de la Plata craton could have begun at ca. 1.6–1.3 Ga. Tectonic events at 630–600 Ma are widespread in Brazil (Fig. 10) and probably mark important collisional events in Brasiliano belts. This age has been recorded in the Araguaia belt between the Amazon and São Francisco cratons, in the Brasília belt, in the Ribeira and Dom Feliciano belts, and in the Borborema province. These belts mark the collision zones between large cratons and/or cratonic fragments (e.g., Encantadas block, Luís Alves block). Brasiliano belts aged 590–550 Ma occur in the coastal regions of Brazil, that is, in the southeastern Ribeira belt and Araçuaí belt to the east of the São Francisco craton (Fig. 10). They might be linked to belts of the same age in western Africa. In contrast to juvenile island arc accretion, which is characteristic of the 900- to 700-Ma belts, most of the young (630- to 550-Ma) Brasiliano belts seem to have evolved on old crust (Kröner and Cordani, 2003). The distinction of orogenic episodes comprising (1) 800- to 700-Ma juvenile volcanic arcs, (2) 640- to 630-Ma collision and crustal reworking, and (3) collisional events between 590 and 500 Ma is compatible with the concept of three stages, Brasiliano I, II, and III, of the Brasiliano orogenic cycle delineated for southern and southeastern Brazil (in the Mantiquera province) (e.g., Basei et al., 2000; Trouw et al., 2000; da Silva et al., 2005). The discrimination of episodes is based on compilation and integration of various age data sets. The data, however, also show that coastal belts in Brazil, such as the 630- to 600-Ma Ribeira and Dom Feliciano belts, cannot be directly linked to Pan-African belts in southwestern Africa (e.g., the Kaoko, Damara, and Gariep belts) because the main orogenic events in the latter occurred ~50–70 m.y. later. Instead, linkages of these belts are likely located on the present-day shelf regions off the coasts of eastern Brazil and western Africa. (West) Gondwana Assembly The Rio de la Plata craton is poorly exposed and extensively covered by the Paraná-Chaco basins so that its limits are mainly
inferred from geophysical data. In contrast to most previous studies, Kröner and Cordani (2003) distinguish the Paraná cratonic block to the north of the Rio de la Plata craton. The blocks are separated by a 630- to 600-Ma belt in the Tebiquary River area. This observation is based on a correlation and connection with the Ribeira belt implied by similar geochronological patterns. Many reconstructions (e.g., Weil et al., 1998; Dalziel et al., 2000; Meert, 2003) attach the Rio de la Plata craton to Amazonia, which is placed along the eastern margin of Laurentia as result of the Grenvillian orogeny. This configuration is debated, however, because a comparison of ore deposits in eastern Laurentia and West Gondwana argues against a genetic relationship (De Witt et al., 1999), and terranes in the Andes cannot be correlated with the Grenville belt in Laurentia (Dalla Salda et al., 1992; Finney et al., 2003). Moreover, Cordani et al. (2003) state that Grenvilleage rocks in Amazonia formed as a result of extension rather than collisional tectonics related to the formation of Rodinia. Geochronological data for the Rio de la Plata craton lack any evidence of a 1.1- to 0.9-Ga Grenvillian event (Hartmann et al., 2000a), at least in its eastern parts. Thus, Grenvillian belts cannot simply be traced into the Rio de la Plata craton. The absence of collisional tectonic events between 2.0 and 0.9 Ga in the Rio de la Plata craton in southern Brazil (Hartmann et al., 2000a; da Silva et al., 2005) and in the Encantadas Complex (da Silva et al., 1999) indicates that the Southern Brazilian Shield was not affected by Grenvillian orogenesis related to the amalgamation of Rodinia. In contrast, the Brasilian orogenic cycle related to Gondwana assembly had already started during the final stages of the Grenvillian orogeny. The only exception is the Punta del Este terrane in eastern Uruguay (Preciozzi et al., 1999; Basei et al., 2000), which represents a fragment of the Kalahari craton, based on provenance relationships with the Gariep orogenic belt (Basei et al., 2005). A position for the Rio de la Plata craton separate from the Amazon craton is supported by the Neoproterozoic Paraguai belt (Fig. 10), which is located between Amazonia and Rio de la Plata and is believed to have resulted from collision of the two cratons during the final stage of West Gondwana assembly (Alkmim et al., 2001). This position suggests that Amazonia and Rio de la Plata were not connected until ca. 600 Ma. As a consequence, recent Rodinia reconstructions (Alkmim et al., 2001; Cordani et al., 2003; Kröner and Cordani, 2003) show the Rio de la Plata craton as an isolated continental block separated by oceans from the Kalahari and Congo cratons as well as from Rodinia. Alternatively, the Rio de la Plata craton was located at a peripheral position in a passive margin setting of this supercontinent during the Mesoproterozoic Rodinia assembly. A peripheric or isolated position of the Rio de la Plata craton, at least in relation to the Rodinia supercontinent, with a long-lived passive margin at its eastern side can explain (1) the absence of any 2.0- to 1.0-Ga orogenic events and (2) a change at ca. 0.9– 0.88 Ga from a passive margin environment to a 200-Ma history of island arc accretion and active continental margin magmatism in the São Gabriel block, marking the beginning of the Brasiliano orogenic cycle. Figures 1, 9, and 11 give an overview of the
The assembly of West Gondwana—The view from the Rio de la Plata craton
A
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575–545 Ma AUS
ANT
IND
K DF C
LAUR
E L
RP
seafloor spreading
545–530 Ma subduction
AUS
ANT
IND
K
C
RP
LAUR
B
Figure 11. Plate tectonic events between 575 and 530 Ma (based on Rozendaal et al., 1999, modified and supplemented). (A) Subduction and closure of the Mozambique and Adamastor oceans, initiation of subduction along the Gondwanan Paleo-Pacific margin, with deformation along the Transantarctic Mountains (Ross orogen). Opening of the Iapetus Ocean. (B) Final stages of Gondwana assembly, with closure of the Mozambique Ocean and collision of East Gondwana with Kalahari and Indo-Antarctica as well as sinistral strike-slip deformation in southwest Africa (Kaoko, Gariep, and Saldania belts), terrane accretion in the Antarctic Ross orogen, and continued Iapetus spreading. ANT—Antarctica; AUS—Australia; C—Congo craton; IND—India; K—Kalahari craton; LAUR—Laurentia; RP—Rio de la Plata craton with Encantadas microcontinent (E) and Dom Feliciano belt (DF).
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major tectonic events in West Gondwana and illustrate the major suturing events in Gondwana assembly. 900–800 Ma Early Neoproterozoic events have been recorded in Brazil and in the Arabian-Nubian Shield. In South America, subduction of oceanic crust located to the east of the Amazon and Rio de la Plata cratons and accretion of the resulting arcs started at ca. 0.9 Ga, as recorded in the Goiás arc in the Brasília belt (Pimentel et al., 1999) and in the Passinho arc in southernmost Brazil. At the same time, juvenile arcs also formed in the Arabian-Nubian Shield (Abdelsalam and Stern, 1996; Stern, 2002). Most of the juvenile crust formed in intraoceanic convergent margin settings (Stern and Abdelsalam, 1998; Cosca et al., 1999). The existence of oceanic basins is also documented in passive margin sequences in the Araçuaí orogen, the southern Brasília belt, and in the Brusque Group (Santa Catarina) and Lavalleja Group (Uruguay) (Heilbron et al., 2004). 800–700 Ma A metamorphic peak at ca. 790 Ma occurs in the Brasília belt (Ferreira Filho et al., 1994), and an accretionary wedge developed in the southern part of the belt (Pimentel et al., 1999) related to a 750-Ma active continental margin. The Rio Negro I magmatic arc in the Ribeira orogen has the same age (ca. 790 Ma) (Heilbron et al., 2004). Subduction to the west beneath the eastern margin of the Rio de la Plata craton, comprising the attached Passinho island arc, led to development of an active continental margin in the São Gabriel block between 780 and 700 Ma. At the same time, the Porongos basin formed on stretched and thinned continental crust on the passive margin of the Encantadas block. In the Dom Feliciano belt farther east, crustal thinning led to extensive partial melting of the basement, resulting in granitoid emplacement and resetting of zircon xenocryst ages. Early subduction and arc accretion can also be observed in areas outside West Gondwana. Arc-arc collision occurred between 800 and 700 Ma in the Arabian-Nubian Shield (Abdelsalam and Stern, 1996). Arc-related magmatism in the Armorican Massif in Peri-Gondwana is represented by the ca. 746-Ma orthogneiss of the Pentevrian Complex (Egal et al., 1996; Samson et al., 2003). Initial subduction starting at ca. 730 Ma has also been recorded from Atlantic Canada (Doig et al., 1993; O’Brien et al., 1996) and Great Britain (Patchett et al., 1980; Tucker and Pharoah, 1991). Initial rifting occurred in the Damara belt at ca. 780 Ma (Hoffman et al., 1996) as well as in the Gariep belt at 780–740 Ma (Frimmel and Zartmann, 2000; Frimmel and Fölling, 2004). Rifting was linked to the opening of the Adamastor Ocean. 700–600 Ma Closure of the São Gabriel/Goiás Ocean led to collision of the Rio de la Plata craton with the Encantadas block. Oblique convergence is documented in sinistral ductile shearing and gran-
ite intrusion along the Dorsal de Canguçu Shear Zone that started at ca. 670 Ma and ceased at ca. 620 Ma (Koester et al., 1997). Deformation was followed by brittle strike-slip faulting and pullapart basin development, whereas left-lateral ductile deformation prevailed in the Dom Feliciano belt. In the latter, subvertical transcurrent shear zones developed at ca. 630–610 Ma (Frantz and Nardi, 1992; Philipp et al., 1993; Babinski et al., 1997; Leite et al., 2000). Relaxation of the strike-slip tectonics in the Dom Feliciano belt was followed by extensional tectonics and the emplacement of ca. 600-Ma peraluminous granites (Frantz et al., 2000). Voluminous 630- to 610-Ma granite magmatism is also characteristic for the Florianópolis batholith in Santa Catarina (da Silva et al., 2005). The time span between 630 and 600 Ma is a period of major tectono-metamorphic and magmatic activity in the Brasiliano belts of Brazil. It involves the final closure of the Goiás Ocean and collision of the Amazon and São Francisco cratons (Pimentel et al., 1999), possibly interacting with the Rio de la Plata/Paraná craton (Alkmim et al., 2001) and leading to crustal thickening, east-vergent nappe formation, and metamorphism in the Brasília belt (Pimentel et al., 1991, 1997; Alkmim et al., 2001), which was sandwiched between the two continental masses. At that same time, the Paraná block was welded to the São Francisco craton (Pimentel et al., 1999). Collision-related granitoid magmatism and metamorphism can be observed in the Ribeira belt (Campos Neto and Figueiredo, 1995; Töpfner, 1997; Heilbron and Machado, 2003). Granites up to 650 Ma in age derived from partial melting of older crust were emplaced in the Kaoko belt (Seth et al., 1998). However, the main tectono-metamorphic evolution in this belt is younger and postdates the main orogenies in South America (see below). Oceanic crust preserved in the Gariep belt formed between at least 630 and 600 Ma (Frimmel and Frank, 1998). Avalonia collided with the Amazonia margin at ca. 650 Ma, followed by formation of an Andean-type active continental margin (635–570 Ma) (Nance et al., 2002). The main phase of arc magmatism (640–570 Ma) is recorded in numerous plutonic rocks throughout much of Avalonia (Nance et al., 1991; O’Brien et al., 1996; Murphy et al., 1999), with most being emplaced at ca. 610–580 Ma. A 633- to 607-Ma juvenile arc assemblage has also been reported from Carolina (Samson et al., 1995; Wortman et al., 2000). In the Armorican Massif, granodiorite cobbles found in conglomerates show protolith ages of 670–650 Ma (Guerrot and Peucat, 1990) and can be linked with the early arc magmatism in this area. The east African orogen formed as a result of long-lived subduction and terrane accretion (Stern, 1994). A 640- to 620Ma oblique continent-continent collision from southern Tanzania northward has been attributed to collision of West Gondwana and Indo-Antarctica (Stern, 1994; Meert, 2003). However, other authors suggest that the 640- to 620-Ma event represents the development of a continental arc and that collision occurred between 590 and 500 Ma (Appel et al., 1998; Möller et al., 2000; Boger and Miller, 2004).
The assembly of West Gondwana—The view from the Rio de la Plata craton 600–550 Ma Voluminous post-tectonic 600- to 580-Ma granites intrude the Southern Brazilian Shield and are exposed in the Dom Feliciano belt (Frantz et al., 1999) as well as in the western São Gabriel and Taquarembó blocks (e.g., Remus et al., 1999, 2000b). However, strike-slip shearing continued in localized fault zones until at least 540 Ma. In the Brasília belt, crustal thickening was followed by intrusion of large volumes of post-orogenic K-rich calc-alkaline granites (Pimentel et al., 1999). In contrast to waning deformation in most Brasiliano belts in central and southern Brazil, the main orogenic episode in the Atlantic coast belts—namely, the Araçuaí belt and certain branches of the Ribeira belt—starts in 590–550 Ma (Söllner et al., 1989, 1991; Machado et al., 1996; Brueckner et al., 2000). Only a narrow width of oceanic lithosphere was generated in the Araçuaí belt (Pedrosa-Soares et al., 2001; Martins et al., 2004), and most belts are characterized by crustal reworking. In the Malmesbury belt, in the western part of the Saldania belt in South Africa, the geochemistry of 600- to 540-Ma syntectonic granites suggests a magmatic arc environment involving partial melting of ancient continental crust (Scheepers, 1995). This observation indicates subduction of the Adamastor Ocean. The upper successions of the Gariep basin, including the Oranjemund Formation and the Rocha Group (Uruguay) formed in a back-arc environment (Basei et al., 2005). Boger and Miller (2004) suggest that East Gondwana did not form a single block when it was amalgamated to West Gondwana but can be split into two plates: (1) East Gondwana or the Austro-Antarctic plate, comprising Australia and most cratons of Antarctica, except the northern Prince Charles Mountains; and (2) the Indo-Antarctic plate, made up of India, Sri Lanka, Madagascar, and northern Prince Charles Mountains (Fig. 1). These continental masses also grew by a series of accretionary and collisional events (Boger et al., 2002; Stern, 2002; Meert, 2003). Boger and Miller (2004) suggest that the Indo-Antarctic plate collided first with West Gondwana along its western margin between 590 and 550 Ma (east African orogeny) (cf. Stern, 1994; Meert and Van der Voo, 1997; Meert, 2003; see Figs. 9D and 11). It is striking that collisional events in western and eastern Africa occurred at the same time. The 600- to 550-Ma belts crop out in the coastal areas of Dronning Maud Land (east Antarctica), Sri Lanka, southern India, Madagascar, and parts of southeastern Africa (e.g., Hölzl et al., 1994; Miller et al., 1996; Jacobs et al., 1998; Kröner et al., 2001). The main phase of arc magmatism in the Armorican Massif was active until 570 Ma (Strachan et al., 1996). Most ages for basin development and arc magmatism in subduction zone settings in the Peri-Gondwana belts lie between 570 and 540 Ma, as in the Bohemian Massif (Zulauf et al., 1999) and slightly earlier, at ca. 600 Ma, in its southeastern parts (Finger et al., 2000), in the Saxothuringian belt of Armorica (Linnemann et al., 2000), in the Central Iberian Zone (Vidal et al., 1999), and in the Ossa Morena Zone of Iberia (Giese and Buehn, 1994; Bandrés et al., 2002). An Andean-type active continental margin setting
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persisted until ca. 570 Ma in Avalonia (Nance et al., 2002). In Carolina, juvenile arc accretion was followed by 580- to 540-Ma mature arc suites (Samson et al., 1995; Wortman et al., 2000). 550–500 Ma The main tectonomagmatic activity in the northeastern branch of the Ribeira belt occurred at 560–530 Ma (Machado et al., 1996; Valladares et al., 1996). The end of deformation in the Araçuaí belt is marked by the intrusion of 535- to 500-Ma posttectonic granitoids (Pedrosa-Soares et al., 2001). The youngest tectonometamorphic event recorded in Brazil is the Cambrian Búzios orogeny in the Ribeira belt, with a metamorphic peak at ca. 525–510 Ma (Schmitt et al., 2004). The main tectonism in southwestern Africa took place at ca. 550–530 Ma and thus partly overlaps with deformation in the Araçuaí belt and parts of the Ribeira belt. In the Kaoko belt, deformation is characterized by transpression and granite intrusions (Dürr and Dingeldey, 1996; Goscombe et al., 2003). Syncollisional granites have an age of ca. 550 Ma (Seth et al., 1998). Frimmel et al. (1996) and Frimmel and Fölling (2004) report the formation of an accretionary wedge followed by collision at ca. 560–530 Ma in the Gariep belt associated with a metamorphic peak at ca. 545 Ma (Frimmel and Frank, 1998). Sinistral transpression between 550 and 510 Ma can also be observed in the Saldania belt (Rozendaal et al., 1999), possibly still linked with subduction, because granitoids dated at 560–520 Ma (da Silva et al., 2000a) are interpreted as being emplaced in an active continental margin setting (Scheepers, 1995). Crustal thickening and metamorphism in the Damara belt occurred at ca. 530 Ma (Jung and Mezger, 2003). Meert (2003) notes that tectonism in the Damara and Gariep belts overlaps the 570- to 530-Ma Kuunga orogeny (Meert et al., 1995) in East Gondwana. The author suggests that deformation in the Damara belt was caused by a clockwise rotation of the Kalahari craton, which was hinged in the region of the Zambesi belt. The rotation could explain sinistral transpression in the coastal belts of southwestern Africa and may even have caused far-field effects, such as localized strike-slip faulting, in southern Brazil. Final closure of the southern Mozambique Ocean and collision of East Gondwana (cf. Boger and Miller, 2004) with the Kalahari craton (Kuunga orogeny) and Indo-Antarctica occurred between 540 and 500 Ma (Boger and Miller, 2004). Subduction along Gondwana’s Pacific margin started at ca. 560 Ma (Goodge, 1997), and magmatism was widespread throughout the Transantarctic Mountains from ca. 530 Ma onward (Boger and Miller, 2004). Deformation occurred at ca. 525–515 Ma (Goodge and Dallmeyer, 1992; Myrow et al., 2002) and was associated with accretion of island-arc rocks in Northern Victoria Land (Antarctica) (Gibson and Wright, 1985; Kleinschmidt and Tessensohn, 1987). Deformation in the Ross-Delamerian orogen of Australia and Tasmania began slightly later, at ca. 515 Ma (Crawford and Berry, 1992; Foden et al., 1999; Münker and Crawford, 2000). It is interesting to note that, although collision and amalgamation of cratons predominated in most regions of Gondwana, the 540- to 515-Ma interval represents the timing of rifting of
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Avalonia from Gondwana (Nance et al., 2002). This period is also the time for initial rifting and opening of the Iapetus Ocean (Grunow et al., 1996; Meert, 2003). CONCLUSIONS A number of terranes can be distinguished in southern Brazil. Volcanic arcs (São Gabriel block) have been accreted to the eastern margin of the Rio de la Plata craton. The formation of oceanic crust in the west between the Rio de la Plata craton and different continental terranes of the Brasiliano orogen is well established, and the São Gabriel block contains the relics of this ocean basin (São Gabriel/Goiás Ocean). The Encantadas microcontinent collided with the Rio de la Plata craton and São Gabriel block as a result of the closure of the São Gabriel/Goiás Ocean. The orogenic episodes can be correlated with the Brasiliano stages (I–III) elsewhere in Brazil. The serial closure of oceanic basins and the amalgamation of cratons in both East and West Gondwana occurred over a long time span of nearly 400 m.y. The “Pan-African” orogenic cycle encompasses the initial accretion of island arcs followed by collision of a number of cratons and cratonic fragments. The summary of tectonic events illustrates the episodicity of Gondwana formation. Deformation in individual belts cannot be viewed separately from regional and even global plate tectonic processes. For example, the striking synchroneity of 590- and 550-Ma collisional events in Brazil/western Africa and eastern Africa suggests that these events are genetically linked by the contemporaneous closure of both the Adamastor Ocean in the west and the Mozambique Ocean in the east. Likewise, rotation of the Kalahari craton as a result of collision with East Gondwana caused transpression in the Kaoko, Gariep, and Saldania belts. It is also interesting to note that the accretion of juvenile arcs at 900–850 Ma and 750–700 Ma in southern Brazil is temporally equivalent to comparable events in the Arabian-Nubian Shield, although these areas were far apart and cannot be directly linked. This synchroneity is also the case for 640- to 610-Ma ocean closures occurring in both regions. Regional plate tectonic processes are coupled in the global tectonic framework such that subduction in one region may cause possible far-field effects in an area thousands of kilometers away. Neoproterozoic global plate tectonic scenarios, however, await delimitation. Nevertheless, several orogenic episodes, albeit partly diachronous, can be distinguished in many Gondwanan belts on several continents: 1. 900- to 700-Ma island arc accretion and active continental margin development in South America and within the Arabian-Nubian Shield, and early magmatism in the Armorican Massif; 2. 650- to 600-Ma collision of cratons in South America, the Arabian-Nubian Shield, and continental arc development in the east African orogen, and accretion of Avalonia to Amazonia;
3. 590- to 550-Ma collisional belts in southwest Africa, eastern South America, and collision of east Africa with Indo-Antarctica (east African orogeny), and main phase of arc magmatism in Peri-Gondwanan terranes; and 4. 550- to 500-Ma subduction and magmatic arc development in Peri-Gondwana; final terrane docking in the Ribeira belt; and deformation in the southwest African Kaoko, Gariep, and Damara belts, possibly linked with the final ocean closure and collision of East Gondwana with South Africa and Indo-Antarctica (Kuunga orogeny). The data also show that West Gondwana was not assembled prior to 540–520 Ma, and that terrane docking and deformation within West Gondwana continued during the final collision with East Gondwana. The assembly of Gondwana consequently comprises the continuous amalgamation of cratons and microcontinents, until its final stages, rather than the collision of three or four large cratonic masses. ACKNOWLEDGMENTS We acknowledge the Herrmann-Willkomm-Stiftung, Frankfurt am Main, for travel grants to Brazil. Field support by the Centro de Estudos em Petrologia e Geoquímica, Instituto de Geociências, Universidade Federal do Rio Grande do Sul, is highly appreciated. Visit and work in the Laboratório de Geologia Isotópica at Universidade Federal do Rio Grand do Sul, Porto Alegre, was made possible by financial support from the Deutscher Akademischer Austauschdienst and Coordenação de Aperfeiçoamento de Pessoal de Nível Superior (Brazilian Government). We thank Fátima Bitencourt for supplying geological maps of the area. Victor A. Ramos and R. Damian Nance are thanked for constructive reviews of the manuscript. REFERENCES CITED Abdelsalam, M.G., and Stern, R.J., 1996, Sutures and shear zones in the Arabian-Nubian Shield: Journal of African Earth Sciences, v. 23, no. 3, p. 289–310, doi: 10.1016/S0899-5362(97)00003-1. Alkmim, F.F., Marshak, S., and Fonseca, M.A., 2001, Assembling West Gondwana in the Neoproterozoic: Clues from the São Francisco craton region, Brazil: Geology, v. 29, p. 319–322, doi: 10.1130/00917613(2001)0292.0.CO;2. Appel, P., Möller, A., and Schenk, V., 1998, High-pressure granulite facies metamorphism in the Pan-African belt of eastern Tanzania: P-T-t evidence against granulite formation by continent collision: Journal of Metamorphic Geology, v. 16, p. 491–509, doi: 10.1111/j.1525-1314.1998.00150.x. Babinski, M., Chemale, F., Jr., Hartmann, L.A., Van Schmus, W.R., and da Silva, L.C., 1996, Juvenile accretion at 750–700 Ma in southern Brazil: Geology, v. 24, no. 5, p. 439–442, doi: 10.1130/0091-7613(1996)0242.3.CO;2. Babinski, M., Chemale, F., Jr., Van Schmus, W.R., Hartmann, L.A., and da Silva, L.C., 1997, U-Pb and Sm-Nd geochronology of the Neoproterozoic Granitic-Gneissic Dom Feliciano belt, Southern Brazil: Journal of South American Earth Sciences, v. 10, no. 3–4, p. 263–274, doi: 10.1016/ S0895-9811(97)00021-7. Bandrés, A., Eguíluz, L., Gil Ibarguchi, J.I., and Palacios, T., 2002, Geodynamic evolution of a Cadomian arc region: The northern Ossa-Morena zone, Iberian massif: Tectonophysics, v. 352, p. 105–120, doi: 10.1016/ S0040-1951(02)00191-9. Basei, M.A.S., Siga, O., Jr., Masquelin, H., Harara, O.M., Reis Neto, J.M., and Precozzi, F., 2000, The Dom Feliciano belt of Brazil and Uruguay and its
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Stern, R.J., 2002, Crustal evolution in the east African orogen: A neodymium isotopic perspective: Journal of African Earth Sciences, v. 34, p. 109–117, doi: 10.1016/S0899-5362(02)00012-X. Stern, R.J., and Abdelsalam, M.G., 1998, Formation of continental crust in the Arabian-Nubian Shield: Evidence from granitic rocks of the Nakasib suture, NE Sudan: International Journal of Earth Sciences (Geologische Rundschau), v. 87, p. 150–160. Strachan, R.A., D’Lemos, R.S., and Dallmeyer, R.D., 1996, Late Precambrian evolution of an active plate margin: North Armorican Massif, France, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related periGondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 319–332. Strieder, A.J., Roldao, D.C., and Hartmann, L.A., 2000, The Palma VolcanoSedimentary Supersuite, Precambrian Sul-Riograndense Shield, Brazil: International Geology Review, v. 42, p. 984–999. Tassinari, C.C.G., Munhá, J.M.U., Ribeiro, A., Ciro, T., and Correia, C.T., 2001, Neoproterozoic oceans in the Ribeira belt (southeastern Brazil): The Pirapora do Bom Jesus ophiolitic complex: Episodes, v. 24, p. 245–251. Teixeira, W., Renne, P., Bossi, J., Campal, N., Argella, D., and Filho, M., 1999, 40 Ar-39Ar and Rb-Sr geochronology of the Uruguayan dike swarm, Rio de la Plata craton, and implications for Proterozoic intraplate activity in western Gondwana: Precambrian Research, v. 93, p. 153–180, doi: 10.1016/S0301-9268(98)00087-4. Tommasi, A., and Fernandes, L.A.D., 1990, O ciclo brasiliano na porçãoo sudeste da Plataforma Sul-americana: Um novo modelo: Congresso Uruguayo de geologica, Montevideo, 1st: Anais, v. 1, p. 107–114. Tommasi, A., Vauchez, A., Fernandes, L.A.D., and Porcher, C.C., 1994, Magmaassisted strain localization in an orogen-parallel transcurrent shear zone of southern Brazil: Tectonics, v. 13, p. 421–437, doi: 10.1029/93TC03319. Töpfner, C., 1997, Age and origin of Brasiliano-granitoids in the southern Ribeira mobile belt by means of U/Pb-zircon and Rb/Sr-whole-rock dating: South American Symposium on Isotope Geology, Campos do Jordão, extended abstracts, 1997, p. 314–316. Trouw, R., Heilbron, M., Ribeiro, A., Paciullo, F., Valeriano, C.M., Almeida, J.C.H., Tupinambá, M., and Andreis, R.R., 2000, The central segment of the Ribeira belt, in Cordani, U.G., Milani, E.J., Tomaz, A., and Campos, D.A., eds., Tectonic evolution of South America: 31st International Geological Congress Rio de Janeiro, Sociedade Brasileira de Geologia, Rio de Janeiro, 2000, p. 355–365. Tucker, R.D., and Pharoah, T.C., 1991, U-Pb zircon ages of late Precambrian rocks in southern Britain: Journal of the Geological Society of London, v. 148, p. 435–443. Valladares, C.S., Heilbron, M., Figueiredo, M.C.H., and Teixeira, W., 1996, Geochemistry and geochronology of Paleoproterozoic gneissis rocks of the Paraíba do Sul Complex (Quirino unit), Barra Mansa region, Rio de Janeiro, Brazil: Revista Brasileira Geosciências, v. 27, p. 111–120. Vidal, G., Palacios, T., Moczydlowska, M., and Gubanov, A.P., 1999, Age constraints from small shelly fossils on the early Cambrian terminal Cadomian Phase in Iberia: Geologiska Föreningens i Stockholm Förhandlingar, v. 121, p. 137–143. Weil, A.B., Van der Voo, R., Niocaill, C.M., and Meert, J.G., 1998, The Proterozoic supercontinent Rodinia: Paleomagmetically derived reconstructions for 1100 to 800 Ma: Earth and Planetary Science Letters, v. 154, p. 13–24, doi: 10.1016/S0012-821X(97)00127-1. Wildner, W., 1990, Caracterização geológica e geoquímica das sequências ultramáficas e vulcano-sedimentares da região da Bossoroca-RS [M.S. thesis]: Porto Alegre, Brazil, Instituto de Geosciências, Universidade Federal do Rio Grande do Sul, 170 p. Wortman, G.L., Samson, S.D., and Hibbard, J.P., 2000, Precise U-Pb zircon constraints on the earliest magmatic history of the Carolina terrane: Journal of Geology, v. 108, p. 321–338, doi: 10.1086/314401. Zulauf, G., Schitter, F., Riegler, G., Finger, F., Fiala, J., and Vejnar, Z., 1999, Age constraints on the Cadomian evolution of the Teplá-Barrandian unit (Bohemian Massif) through electron microprobe dating of metamorphic monazite: Zeitschrift der Deutschen Geologischen Gesellschaft, v. 150, p. 627–639.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Geodynamic evolution of the northwestern Paleo-Gondwanan margin in the Moroccan Atlas at the Precambrian-Cambrian boundary Andre Pouclet* Institut des Sciences de la Terre d’Orléans, UMR 6113, Université d’Orléans, B.P. 6759, 45067 Orléans cedex 2, France Abdellatif Aarab Ecole Normale Supérieure, Université, BP 5118, 10000 Rabat, Morocco Abdelilah Fekkak Faculté des Sciences, Université Chouaib Doukhali, B.P. 20, 24000 El Jadida, Morocco Mohammed Benharref Faculté des Sciences, Université Cadi Ayyad, B.P. 518, Marrakech, Morocco
ABSTRACT In the southern Moroccan Atlas, abundant volcanic and sedimentary formations, dated from the Ediacaran to Cambrian time, were set at the northwestern Paleo-Gondwanan margin, after the main Pan-African orogenic event. The Precambrian-Cambrian geodynamic transition is characterized by an Early Cambrian marine transgression. We examine the tectonic conditions of this transgression and the magmatic signatures of the volcanic rocks that were produced just before and around the PrecambrianCambrian boundary. Significant angular unconformities are evidenced, between the Late Neoproterozoic formations and the Cambrian deposits, in the central and eastern Anti-Atlas, which are due to a late Ediacaran NNE-SSW compressional event. The Late Neoproterozoic formations are related to an intracontinental volcanic chain of andesitic to rhyolitic lavas dated to the Ediacaran period. These calc-alkaline rocks were generated by melting of the mantle, previously metasomatized during the PanAfrican orogenic stage, and of continental crust. The Late Ediacaran to Early Cambrian formations are analyzed in the Agoundis-Ounein and Toubkal areas, southwest of the old block of High-Atlas. An important basaltic pile unconformably overlies the Ediacaran rhyolitic formation and is overlain by Tommotian sediments. These basalts are continental tholeiites generated by melting of a normal subcontinental mantle. They outpoured from an important N 30°-trending fissural system over a basin floor. Some lherzolite fragments have been sliced along southwest-northeast faults, in the Lower Cambrian sediments. They originated from a transitional mantle between continental and oceanic domains. Farther east of the central Anti-Atlas, the Tommotian Djbel Boho
*E-mail:
[email protected]. Pouclet, A., Aarab, A., Fekkak, A., and Benharref, M., 2007, Geodynamic evolution of the northwestern Paleo-Gondwanan margin in the Moroccan Atlas at the Precambrian-Cambrian boundary, in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, p. 27–60, doi: 10.1130/2007.2423(02). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Pouclet et al. volcano exhibits olivine basalts having an intraplate enriched asthenospheric signature type of ocean island basalt. The magmatic characteristics of the Late Ediacaran to Early Cambrian volcanic rocks, the structural features, and the presence of lherzolite fragments are consistent with a volcanic passive margin rift setting in a WNW-ESE extension regime. The meaning of this extensional event is discussed in relationships to the opening of a Cambrian basin and the drifting of the Avalonian terranes. Keywords: Early Cambrian, Paleo-Gondwana, passive margin, continental tholeiite, Atlas
INTRODUCTION Geologic Background Late Precambrian to Cambrian Time The Pan-African ocean closure caused the formation of Pannotia, in the Late Neoproterozoic. The west African, central African, south African, and South American cratons coalesced to form the Paleo-Gondwana supercontinent. In the Ediacaran time, the northwestern African part of this supercontinent was constituted by the suturing of the Anti-Atlasic Pan-African mobile belt
32°
undistinguished Early to Middle Paleozoic
toward the west African craton Proterozoic margin (Leblanc and Lancelot, 1980; Saquaque et al., 1989; Hefferan et al., 1992, 2000; Leblanc and Moussine-Pouchkine, 1994), followed by accretion or docking of the High-Atlas and Meseta terranes. The Pan-African suture zone locates along the Anti-Atlas major fault (AAMF, Fig. 1), where two ophiolitic complexes are preserved (Leblanc, 1981; El Boukhari et al., 1992; Admou and Juteau, 1998; Admou, 2000). An alternative location along the South Atlas fault has been proposed by Ennih and Liégeois (2001, 2003). However, it seems that the structural and petrological features of the ophiolitic complexes, the lithostratigraphic framework, and the interpretation of
8°
6° 32°
Early Paleozoic
Mk
Precambrian
CHA
Figs. 6, 7
WHA
Ag
SAF
Tb
Sg
Oz
A
EA
Sr
2
CAA 3 AAMF
TNT
Zn
Db
5
ure n- African sut Pa
Kr 1
6
4 Bz
Ig 30°
Og
7
A WA
West-African Craton
8?
30?
6°
Figure 1. Map of the Atlasic region. AAMF—Anti-Atlas major fault; CAA—central Anti-Atlas; CHA—central High-Atlas; EAA—eastern Anti-Atlas; SAF—south Atlas fault system; TNT—Tizi n’Test fault system; WAA—western Anti-Atlas; WHA—western High-Atlas. Precambrian inliers: Bz—Bou Azzer; Ig—Ighrem; Kr—Kerdous; Og—Ougnat; Sg—Saghro; Sr—Siroua; Zn—Zenaga. Ag—Agadir; Db—Djbel Boho volcano; Mk—Marrakech; Oz—Ouarzazate; Tb—Toubkal Massif; 1–7—sites of the Precambrian-Cambrian boundary. The large and the small frames are the location of the Figures 6 and 7, respectively. They are located in the eastern side of the WHA, which is the old block of High-Atlas characterized by the Precambrian rocks of the Siroua northern part. Note that the TNT and the SAF separate the High-Atlas from the Anti-Atlas. Heavy lines are major faults. Insert: location of the studied area, north of the West African Craton, straddling the Pan-African suture.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin the aeromagnetic data are in favor of the Anti-Atlas major fault location (Saquaque et al., 1992; Hefferan et al., 2002; Bouougri, 2003; Beraaouz et al., 2004; Bouougri and Saquaque, 2004; Inglis et al., 2004; Samson et al., 2004; Soulaimani et al., 2006). The Pan-African north oceanic domain subducted, between 750 and 660 Ma, below the Saghro terrane, which can be considered a mobile zone (Thomas et al., 2004). The Pan-African accretion was followed by the erection of a volcanic chain that covered the entire Anti-Atlas and parts of the High-Atlas in a postorogenic context. At the end of the Neoproterozoic, destruction of the volcanic mountains caused thick deposition of molassic conglomerates into a sinking trough of pull-apart types, followed by leveling of the relief. The Cambrian marine transgression took place above this paleotopography and extended from a new oceanic domain that opened northwest of Paleo-Gondwana. This opening resulted from a tectonic extension of the western margin and rift formation (Bernardin et al., 1988; Piqué et al., 1990, 1995; El Attari et al., 1997; Piqué, 2003; Soulaimani et al., 2003, 2004; El Archi et al., 2004). The transgressive sediments deposited above the Precambrian formations, either in an apparent conformity or with an angular unconformity. They constitute the local Adoudounian formations of the Taroudant and Tata groups. Radiometric data place these formations in the Early Cambrian, from 535 to 520 Ma (Ducrot and Lancelot, 1977; Compston et al., 1992; Landing et al., 1998; Levresse, 2001; Gasquet et al., 2005; Maloof et al., 2005; Table 1). However, in a detailed stratigraphic study of the northern margin of the western Anti-Atlas, Maloof et al. (2005) defined the Tabia Member and the Tifnout Member in the lower Adoudounian formations. The δ13C record of a set of stratigraphic columns shows a –6‰ nadir at the base of the Tifnout Member that could be the Ediacaran-Cambrian boundary, by comparison with the Siberian chemostratigraphy. This boundary is dated at 542.0 ± 0.6 Ma (Amthor et al., 2003). According to this interpretation, the base of the “Cambrian” transgression is dated at the late end of the Ediacaran Period. Post-Cambrian Folding Events The High-Atlas and the westernmost part of the Anti-Atlas registered the Variscan orogeny. Two major folding events took place in the late Visean age and in the late Carboniferous Period (Houari and Hoepffner, 2003; Hoepffner et al., 2005, 2006). The High-Atlas underwent a WNW-ESE major compression responsible for the NNE–SSW-trending reclined folds, or vertical and parallel folds in the less deformed southernmost area. Then a dextral wrench-dominated transpression along the north African craton margin developed ENE–WSW-trending folds in the northwestern and northern parts of the Anti-Atlas (Soulaimani et al., 1997; Houari and Hoepffner, 2003). In the rest of the Anti-Atlas, the Paleozoic terrains are subhorizontal or slightly tilted. The Alpine or Atlasic orogeny reworked the major faults of the Precambrian and Paleozoic substratum, which was uplifted in different blocks. The Mesozoic and early Cenozoic cover is thrust to the south, in a thin-skin tectonic style (Benammi et al., 2001).
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Cambrian Transgression The Precambrian substratum is exposed in the Anti-Atlas windows and in the old block of High-Atlas, east of the Paleozoic western High-Atlas (WHA, Fig. 1). It consists of sedimentary, volcanic, and plutonic rocks named P-I, P-II, and P-III (Choubert, 1963). P-I is the oldest basement cropping out in the western Anti-Atlas (WAA, Fig. 1) inliers. It is dated to Paleoproterozoic and belongs to the northwestern margin of the west African craton (Charlot, 1982; Aït Malek et al., 1998; Thomas et al., 2002; Walsh et al., 2002). Thanks to U-Pb dating (referenced in Thomas et al., 2004), the P-II and P-III terranes are placed in two chronostratigraphic supergroups (Table 1): (1) the Anti-Atlas Supergroup (the former lower P-II), dated to the Cryogenian Period and corresponding to sedimentary and magmatic rocks affected by the Pan-African orogenic phase in the mobile belt; and (2) the Ouarzazate Supergroup (the former upper P-II and P-III), dated to the Ediacaran Period and corresponding to the late- and postorogenic formations. The Taroudant and Tata groups correspond to the transgressive deposits above the former Precambrian formations. They are dated to the Early Cambrian and possibly to the late Ediacaran Period, according to the δ13C record, for the lowest deposits (see above). However, to simplify, the term Cambrian transgression will be used in this article. Questions arise concerning the geotectonic setting of the Cambrian transgression and its precise timing, because it occurred at the Precambrian-Cambrian boundary, a rather symbolic geological period. It has been assumed that the transgression occurred in a late Pan-African extensive continuum, as indicated by postorogenic molassic deposition in a sinking trough along NNE-SSW to NE-SW normal faults (Badra et al., 1992; Chbani et al., 1999; Piqué et al., 1999; Algouti et al., 2000, 2001; Benssaou and Hamoumi, 2003; Soulaimani et al., 2003, 2004). But the Cambrian deposits are rarely in perfect conformity above the late Precambrian terrains. In most cases, an erosional unconformity and a more or less important sedimentary gap is observed. In other cases, clear angular discordances are conspicuous. The unconformities could be explained by block tilting, which usually occurs in any extensive structural context, such as postorogenic collapse and rifting (Youbi, 1998; Jouhari et al., 2001, El Archi et al., 2004; Soulaimani et al., 2004). However, several WNW–ESE-trending anticlinal and synclinal fold axes have been observed in the Ouarzazate Supergroup formations, but not in the overlying “Cambrian” formations. Alternatively, one may support the hypothesis of a particular tectonic event, in the latest Neoproterozoic. Moreover, there is a drastic change in the geochemical compositions of the magmatic products, which are calc-alkaline and mainly acidic in the Late Neoproterozoic, and tholeiitic and alkaline basaltic in the Early Cambrian. In this article, (1) we describe the different types of Precambrian-Cambrian boundaries, in terms of structural and sedimentary features, in seven sites selected all over the Atlas; (2) we examine the nature of the magmatic products, in the Late Neoproterozoic and the Early Cambrian. As there is a crucial problem in
Toyonian
542
Tommotian NemakitDaldynian
530
534
C1
Divergent stage
Convergent stage
ca 660 (2)
Late orogenic stage
P-III basaltic volcanics
Basal Formation
Tamjout dolostone or Lower limestones
Lie de vin Formation
Upper limestones
“Schisto-calcaire”
“Série schisteuse”
Passive margin sediments—marine basin formations ca. 762 Ma (1)— intracontinental rifted basin formations
Oceanic arc- and continental margin arc-volcanics and associated basin sediments
Syntectonic plutons straddling the main tectonic phase
Main collisional-related Pan-African tectonic phase
400 m)
4
Tanalt Formation
Basal Formation (100 m)
Basal Formation (140 m)
3 Anzel
Rhyolite-andesite complex
N-Imiter
7
100 m
Middle Cambrian
Tagmout Tin Ouayour
Agoundis-Ounein
E-Kerdous Tamjout dolostone (70 m)
2
Tamjout dolostone (50 m)
Tamjout dolostone Basal Formation (30 m)
Precambrian
1
32 Pouclet et al.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin the polymictic conglomerate of Trifya, the thick upper member, includes blocks of andesites and ignimbrites of the Ouarzazate Supergroup, and of the Bleida granodiorite dated to 579 Ma (Inglis et al., 2004). Thus the Trifya Formation is dated to the Ediacaran Period, and the angular unconformity between this formation and the overlying terranes (particularly the Early Cambrian Tamjout dolostone) must be dated to the Late Ediacaran. Site 6 is located at the southern border of the Saghro window, at the edges of the Tagmout Tin Ouayour detrital basin. The Late Neoproterozoic formations consist of rhyolitic flows and pyroclastic beds dipping 20°–30° to the southeast and belonging to the Ouarzazate Supergroup. Above an erosional unconformity, a 200-m-thick subhorizontal detrital basin sequence is made of heterometric conglomerates and interbedded epiclastites and sandstones in the lower part, and of a fining-upward alternation of sandstones and siltstones in the upper part. This sequence is conformably overlain by a new 10-m-thick conglomeratic bed and by a succession of meter-sized layers of sandstones, dolostones, and shales. About 30 m above this upper conglomerate, a detrital level includes the Micmacca Breccia, a fossiliferous debris attributed to the Middle Cambrian (Neltner, 1938; Hupé, 1952). Details of the stratigraphic column are given by Benziane et al. (1983). Because the uppermost units are dated to the Middle Cambrian and there is no unconformity or sedimentation break, the lower detrital sequence may correspond either to the upper Ouarzazate Supergroup or to the eastern continental equivalent of the western marine Early Cambrian deposits. Thus the Precambrian-Cambrian boundary is either at the subrhyolitic unconformity or at the uppermost conglomerate layer. Site 7 is located at the northern border of the Saghro window, at the Imiter inlier edge. The Late Neoproterozoic formations are made of rhyolitic flows and of pelitic and pyroclastic layers, all dipping 30° to the northwest and belonging to the Ouarzazate Supergroup. The youngest rhyolitic activity is dated at 550 ± 3 Ma (Levresse, 2001; Cheilletz et al., 2002). The Cambrian sequence is subhorizontal or gently dipping to WNW and shows a clear angular unconformity with the Ouarzazate Supergroup volcanic rocks. It begins with a meter-sized conglomerate of rhyolitic and andesitic pebbles in a carbonated cement. Above, sedimentation is fining upward, with sandstones and shales, including decimeter-sized limestone beds. The shales are fossiliferous, with fragments of trilobites (Paradoxides) and brachiopods, which are dated to the early Middle Cambrian. Thus the Early Cambrian deposits are missing. Evidence for a Late Precambrian Tectonic Event This overview of the Precambrian-Cambrian boundary sites shows that different situations may have prevailed at these sites. The “Cambrian” units were deposited above the Precambrian formations made of acidic volcanic and volcanoclastic formations (Sites 3 and 7) and of heterometric conglomerates rich in rhyolitic and granitic elements (Sites 1, 4, 5, and 6). In the northwestern area, the Cambrian units overlay a thick occurrence of
33
basaltic lava flows (Site 2). The Cambrian deposition is conformable above the Late Neoproterozoic conglomerates (Sites 1 and 6) and above the basalt pile (Site 2); or it is more or less unconformable above the same conglomerate (Sites 4 and 5) and above the acidic volcanic rocks (Sites 3 and 7). The conglomeratic formation itself is either conformable (Site 1) or unconformable (Site 6) with the underlying acidic volcanic rocks. The transgressive formations are dated to latest Ediacaran Period or to Early Cambrian time in the western, northwestern, and central areas, but to the Middle Cambrian in the eastern area. The unconformity of the “Cambrian” transgressive deposits above the Late Neoproterozoic formations implies a late Precambrian tectonic event. Many authors, after Choubert (1963), attribute these discordances to local tilting of blocks in an extensional regime (Azizi-Samir et al., 1990; Piqué et al., 1999; Soulaimani et al., 2003, 2004). Normal faults trending north-south to northeast-southwest may be related to a northwest-southeast main extension that continues in the Cambrian time. However, in central and eastern Anti-Atlas, the late Neoproterozoic formations are not only tilted, but folded. South and east of Ouarzazate and southeast of Bou Azzer, one can observe in the Ouarzazate Supergroup formations a set of WNW–ESE-trending fold axes, belonging to hectometer-scaled parallel folds with 50°–60° dipping flanks. Such structures resulted from a SSW-NNE compressive event. This folding is absent in the overlying Cambrian formations, which are subhorizontal in the central and eastern Anti-Atlas. The meaning of this tectonic phase can be documented by investigating the composition of the volcanic rocks below and above the unconformity. EDIACARAN MAGMATIC ACTIVITY OF THE OUARZAZATE SUPERGROUP The Late Neoproterozoic post-Pan-African collision formations consist of detrital deposits and abundant volcanic rocks known as the Ouarzazate Supergroup. The magmatic activity is typically calc-alkaline, with moderately potassic andesites and high-potassic dacites, rhyolites, and cognate subvolcanic granites. These formations constitute the Ediacaran Atlasic Volcanic Chain, an important volcanic chain that built across the whole Anti-Atlas and in the old block of High-Atlas. This chain extends, in a N 65° direction, from the Atlantic coast to the border of Algeria, at the Boukaïs inlier (Seddiki et al., 2004). The total length reaches 850 km and the width, 80–150 km. About twenty accurate U-Pb zircon datings are available (Charlot, 1982; Mifdal et al., 1982; Mifdal and Peucat, 1985; Aït Malek et al., 1998; Levresse, 2001; Cheilletz et al., 2002; Thomas et al., 2002; Walsh et al., 2002; Gasquet et al., 2005). The ages range from 578 ± 5 Ma to 543 ± 9 Ma. A new extensive sampling of this volcanic chain (Table 2) exemplifies the geochemical characteristics, taking into account the previous compilation of Bajja (2001) and the data of Youbi (1998), Ezzouhairi (2001), and Zahour (2001). There are no geochemical variations in space, from west to east or along the chain.
34
Pouclet et al.
TABLE 2. CHEMICAL ANALYSES OF THE VOLCANIC AND GRANITIC ROCKS OF THE EDIACARAN ATLASIC VOLCANIC CHAIN Location Sample number
Kerdous Ker-7 Flow And
Imiter Pbf-2 Dike And
Kelaat Mgouna Ay-1 Plug And
Im-5 Flow And
Imiter Im-4 Flow And
Im-6 Flow And
Bou Azzer Da-26 Flow And
Imiter Im-7 Flow And
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
50.81 1.36 17.64 12.64 0.14 6.69 1.85 3.98 1.00 0.31 3.45 99.87
51.77 1.06 15.82 8.08 0.11 6.36 3.39 1.34 3.87 0.35 7.98 100.13
51.91 1.12 18.70 9.70 0.12 4.75 2.14 3.28 3.78 0.41 4.04 99.95
54.39 1.00 17.88 7.91 0.13 3.01 2.31 4.10 3.43 0.25 5.50 99.91
54.53 1.02 17.75 7.40 0.10 2.56 3.86 3.78 3.14 0.27 5.45 99.86
54.62 0.99 17.86 8.23 0.13 2.60 3.18 5.39 1.94 0.28 4.69 99.91
54.68 0.68 13.18 6.95 0.15 4.49 9.11 2.56 0.74 0.13 6.42 99.09
54.82 1.01 17.89 8.43 0.11 2.68 3.16 5.76 1.22 0.27 4.52 99.87
55.01 0.85 16.57 8.11 0.17 4.82 5.10 3.22 2.95 0.19 2.95 99.94
55.85 0.90 17.39 8.12 0.18 3.30 3.60 3.87 3.56 0.20 2.52 99.49
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
154 264.2 57.5 106.6 15.16 25.0 229 17.50 203 7.50 1.12 385 4.66 0.67 4.74 2.10
158 126.0 25.8 51.4 21.26 166.6 154 19.55 182 5.59 6.52 1916 4.25 0.40 4.50 2.85
180 86.2 24.6 24.7 22.12 162.6 229 20.84 125 6.72 3.26 1146 3.12 0.57 5.66 2.29
121 10.2 20.7 9.0 25.25 124.6 231 22.61 168 8.67 4.21 1024 4.47 0.73 4.57 1.76
122 7.1 18.2 7.7 25.10 144.7 210 23.99 166 8.91 7.07 762 4.69 0.72 4.58 1.78
120 12.5 18.1 8.7 23.90 73.6 391 20.70 150 8.03 3.91 787 3.92 0.64 3.94 1.66
107 371.9 24.9 41.1 18.13 38.3 550 19.11 119 5.95 2.37 479 3.36 0.49 5.18 1.45
115 13.3 18.6 9.9 22.50 55.4 395 22.00 159 8.26 2.94 463 4.18 0.64 4.03 1.52
162 117.0 24.6 43.1 19.40 108.0 502 17.10 135 5.57 2.19 841 3.52 0.44 4.25 2.26
156 46.5 17.0 33.3 19.02 140.3 348 25.14 144 5.98 1.55 961 4.15 0.63 7.35 5.36
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
19.37 42.71 5.89 24.06 5.09 1.49 4.16 0.64 3.54 0.63 1.70 0.24 0.23 1.52
29.97 68.53 8.97 37.62 7.00 2.20 5.52 0.70 3.69 0.65 1.87 0.25 0.27 1.54
16.00 39.00 5.22 22.40 5.07 1.39 4.47 0.62 3.66 0.67 1.79 0.29 0.26 1.85
28.00 57.20 6.99 28.40 6.26 1.77 4.82 0.71 4.12 0.82 2.30 0.39 0.34 2.15
28.40 57.90 7.26 29.60 6.36 1.58 5.31 0.77 4.65 0.84 2.24 0.35 0.40 2.19
22.40 48.90 6.13 25.30 4.87 1.12 4.65 0.69 3.78 0.72 1.94 0.31 0.30 2.23
22.90 45.37 5.41 21.27 4.32 1.10 3.69 0.57 3.30 0.65 1.84 0.27 0.27 1.79
23.10 50.80 6.36 26.50 5.09 1.51 4.95 0.73 3.94 0.77 2.08 0.29 0.33 2.08
17.80 37.30 4.74 19.40 3.92 1.04 3.46 0.53 2.87 0.60 1.65 0.26 0.26 1.61
16.01 34.77 4.45 18.60 4.41 1.06 4.27 0.69 4.25 0.85 2.48 0.38 0.40 2.48
Rock type
Kelaat Mgouna Ouarzazate Fw-w Sk-9 Dike Flow And And
(wt%)
Continued
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
35
TABLE 2. CHEMICAL ANALYSES OF THE VOLCANIC AND GRANITIC ROCKS OF THE EDIACARAN ATLASIC VOLCANIC CHAIN (continued) Location Kelaat Mgouna Sample number Wiz-W Plug Rock type And
Imiter Fbo-7 Flow And
Kerdous Bou Azzer Ouarzazate Kelaat Mgouna Ker-8 Da-24 Sk-4 Ay-6 Flow Flow Flow Flow And And And And
Imiter Pbf-1 Flow And
Km-3 Dike And
Kelaat Mgouna Ay-10 Flow And
Wiz-5 Plug And
(wt%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
56.17 0.88 17.23 8.71 0.15 3.25 6.78 3.36 1.42 0.20 1.58 99.73
57.53 0.71 16.94 7.91 0.15 3.27 4.18 4.72 2.54 0.23 1.82 100.00
57.92 0.98 17.40 8.21 0.14 2.56 2.98 4.20 3.17 0.25 2.05 99.86
58.29 1.15 14.93 7.35 0.09 1.33 4.43 7.24 0.36 0.33 3.61 99.11
58.50 1.27 15.23 8.11 0.07 1.64 3.47 4.21 3.42 0.47 3.06 99.45
59.08 0.89 16.81 8.16 0.14 2.54 3.18 4.51 2.67 0.29 1.70 99.97
59.32 0.57 13.63 6.94 0.10 5.10 3.47 2.54 2.37 0.15 5.72 99.91
60.65 1.03 15.19 8.00 0.08 1.78 1.95 4.39 4.25 0.36 2.22 99.90
61.09 0.90 14.36 8.27 0.09 3.93 1.15 3.91 1.89 0.24 4.01 99.84
61.46 0.89 16.33 6.30 0.10 1.95 3.49 4.26 2.88 0.25 1.99 99.90
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
170 57.9 19.8 35.9 22.50 65.4 520 17.30 120 5.13 4.52 739 3.20 0.43 3.02 1.37
103 51.3 17.4 36.2 21.30 92.1 665 16.64 191 5.15 1.31 1347 5.00 0.40 5.10 3.61
99 12.1 4.8 10.3 17.94 136.9 73 21.49 186 7.36 3.89 943 4.43 0.59 4.97 2.96
95 55.3 8.6 17.6 16.41 36.7 199 16.02 207 8.93 0.43 764 5.31 0.68 6.75 2.30
87 9.6 16.2 12.0 20.56 78.4 265 37.84 236 10.10 0.98 1824 6.21 0.87 7.50 5.20
128 38.1 14.0 21.9 21.09 101.4 501 22.78 210 7.39 0.87 1050 4.75 0.55 5.43 2.42
94 496.9 25.9 123.6 17.46 68.2 184 15.07 127 5.92 1.49 739 3.01 0.43 4.78 3.09
81 10.9 13.7 8.5 19.64 102.8 62 29.95 226 9.43 0.46 950 5.32 0.73 7.73 3.07
132 51.2 13.1 21.6 19.60 59.0 93 16.20 103 5.86 1.87 529 2.33 0.46 4.40 1.95
73 20.1 9.4 11.8 20.90 87.4 368 22.90 181 7.88 2.38 910 4.79 0.72 7.68 3.81
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
15.50 34.60 4.53 19.50 4.10 1.18 3.41 0.54 3.19 0.65 1.72 0.27 0.27 1.74
22.70 49.20 6.11 25.40 4.96 1.44 3.91 0.54 2.88 0.59 1.61 0.26 0.25 1.59
15.73 28.84 3.59 15.03 3.53 1.13 3.95 0.64 3.82 0.73 1.96 0.27 0.26 1.75
29.12 76.07 10.17 44.97 9.62 2.84 6.90 0.78 3.39 0.51 1.13 0.13 0.11 1.19
27.95 63.11 8.12 33.69 7.50 1.77 6.94 1.09 6.49 1.29 3.68 0.55 0.58 3.69
25.50 52.10 6.21 24.80 5.19 1.48 4.10 0.63 3.55 0.76 2.00 0.29 0.31 1.92
18.13 35.54 3.95 15.36 2.86 0.80 2.55 0.40 2.44 0.49 1.45 0.22 0.25 1.46
28.60 62.70 7.83 32.90 5.97 1.46 5.68 0.86 4.66 0.89 2.66 0.37 0.40 2.61
15.80 32.70 3.87 16.30 3.33 0.87 3.01 0.47 2.54 0.53 1.44 0.21 0.23 1.50
27.20 56.70 6.92 29.60 5.54 1.44 4.82 0.69 4.19 0.81 2.28 0.35 0.36 2.36 Continued
36
Pouclet et al.
TABLE 2. CHEMICAL ANALYSES OF THE VOLCANIC AND GRANITIC ROCKS OF THE EDIACARAN ATLASIC VOLCANIC CHAIN (continued) Location Sample number
Ouarzazate Sk-14 Sk-18 Flow Flow Dc-And Dc-And
Toubkal Aa-413 Flow Dc
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
63.37 0.71 16.23 8.33 0.06 0.79 1.20 5.17 3.02 0.21 1.36 100.45
63.47 0.67 16.13 5.93 0.13 1.59 4.46 3.76 3.53 0.19 0.82 100.68
64.04 0.67 17.07 5.19 0.08 1.24 1.24 5.73 3.33 0.20 1.36 100.15
66.23 0.69 15.00 5.40 0.16 1.07 1.34 4.56 3.90 0.16 1.40 99.91
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
33 24.0 6.3 13.6 21.93 98.4 220 43.15 323 16.09 5.24 1096 8.22 1.31 12.29 4.18
77 30.0 11.5 20.7 20.63 93.3 455 23.03 213 10.24 1.58 1157 5.59 0.96 10.34 5.48
42 14.0 8.0 12.1 22.05 90.5 238 18.22 160 7.78 1.85 1113 4.53 0.65 6.20 4.62
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
38.97 87.78 10.82 42.16 8.78 1.56 7.36 1.14 6.93 1.42 4.24 0.67 0.77 4.77
33.57 66.88 7.79 29.22 5.67 1.32 4.67 0.69 4.10 0.78 2.22 0.33 0.37 2.28
21.64 45.89 5.77 22.46 4.55 1.30 3.87 0.58 3.30 0.62 1.71 0.25 0.28 1.72
Rock type
Kelaat Mgouna Bou Azzer Wiz-1 Da-19-1 Plug Flow Dc Dc
Ifni Ifn-2 Flow Dc
Toubkal Aa-404 Flow Dc
Bou Azzer Da-20-1 Flow Dc-Rh
Ouarzazate Sk-16 Flow Dc-Rh
Kerdous Ker-1 Pebble Rh
66.70 0.62 14.52 5.01 0.09 2.38 0.94 4.33 3.37 0.15 1.85 99.96
67.42 0.51 15.70 4.47 0.00 0.79 0.10 2.72 7.74 0.15 0.79 100.39
67.64 0.44 14.05 4.63 0.05 0.78 1.54 5.16 4.05 0.09 1.50 99.93
69.14 0.23 12.36 4.18 0.05 0.79 2.75 3.10 2.77 0.04 3.52 98.93
69.19 0.41 14.88 5.86 0.05 0.36 0.34 4.83 3.53 0.06 0.78 100.29
71.72 0.12 13.61 5.73 0.00 0.00 0.00 6.77 1.82 0.05 0.00 99.82
22 15.6 3.6 11.7 20.70 127.0 236 29.60 239 10.80 2.10 1008 6.61 0.93 10.20 5.14
65 49.4 12.6 18.3 17.25 78.1 154 30.83 198 8.30 1.28 965 5.41 0.67 7.32 2.05
35 23.2 6.2 16.7 22.36 225.5 84 35.79 267 13.27 6.56 996 7.61 1.21 12.60 4.31
48 18.2 6.5 13.7 14.43 80.9 50 26.67 190 7.44 0.43 630 5.48 0.78 10.90 4.19
38 27.5 2.4 10.9 16.15 94.4 82 21.16 220 6.86 2.81 483 6.28 0.63 9.70 2.16
5 18.3 3.0 16.0 23.56 116.7 104 27.10 364 16.07 3.23 1826 8.90 1.34 11.20 5.22
25 18.2 1.1 12.4 13.46 33.7 59 10.70 179 5.21 0.35 447 5.16 0.44 11.31 3.13
35.30 75.50 9.11 36.10 6.90 1.56 6.20 0.90 5.09 1.04 2.93 0.45 0.48 3.02
47.61 95.70 11.93 48.67 9.75 2.07 7.24 1.00 5.48 1.06 3.01 0.45 0.50 3.09
46.31 98.50 11.83 45.30 8.90 1.17 7.28 1.09 6.42 1.25 3.63 0.55 0.58 3.79
42.95 89.54 10.45 39.10 7.15 1.58 5.75 0.84 4.82 0.90 2.49 0.35 0.35 2.33
11.23 22.98 2.46 10.65 2.55 0.59 2.65 0.44 3.00 0.71 2.29 0.38 0.43 2.70
16.65 47.73 5.04 20.05 4.39 0.79 3.86 0.68 4.51 0.96 3.01 0.49 0.57 3.50
8.18 17.13 2.20 8.92 1.87 0.20 1.60 0.25 1.56 0.35 1.18 0.21 0.33 1.75
(wt%)
Continued
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
37
TABLE 2. CHEMICAL ANALYSES OF THE VOLCANIC AND GRANITIC ROCKS OF THE EDIACARAN ATLASIC VOLCANIC CHAIN (continued) Location Sample number
Ouarzazate Sk-8 Flow Rh
Kelaat Mgouna Gt-1 Dome Rh
Ifni Ifn-1 Flow Rh
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
71.80 0.17 14.19 3.63 0.00 0.51 0.20 3.39 5.35 0.04 1.12 100.40
72.25 0.29 13.24 3.80 0.05 0.23 0.93 3.64 4.68 0.11 0.67 99.89
74.15 0.07 11.90 3.25 0.00 0.00 0.00 0.00 10.39 0.04 0.00 99.80
75.46 0.08 11.44 2.97 0.00 0.00 0.09 0.53 8.86 0.06 0.46 99.95
60.31 0.82 17.21 6.59 0.15 2.46 2.34 4.53 3.76 0.25 1.85 100.27
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
7 19.7 1.8 14.9 22.37 176.6 81 31.89 290 12.67 2.72 1745 7.68 1.15 13.24 5.02
16 11.8 2.9 12.1 22.19 179.6 89 50.60 350 14.29 2.38 755 8.74 1.28 14.79 4.87
17 21.8 1.0 15.4 15.28 180.9 23 24.47 383 17.99 3.08 1433 9.22 1.08 12.84 3.50
6 15.7 2.3 27.4 14.25 208.9 56 37.42 200 16.49 0.74 1745 7.99 1.51 18.27 3.28
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
17.71 40.06 4.52 17.17 3.84 0.61 3.97 0.74 5.11 1.08 3.28 0.51 0.55 3.49
48.10 102.10 12.03 46.30 8.83 1.11 7.97 1.38 7.98 1.54 4.42 0.68 0.68 4.93
25.97 55.48 6.31 23.56 4.35 0.21 3.55 0.55 3.49 0.77 2.49 0.43 0.51 3.16
43.90 87.90 11.15 43.80 8.24 0.88 6.84 1.15 6.67 1.40 4.31 0.72 0.75 5.11
Rock type
Kelaat Mgouna Toubkal Ftc Aa-414 Aa-411 Dike Apophysis Pluton Rh Mdior Dior
Ifn-4 Pluton Gr musc
Ifni Ifn-3 Pluton Gr Bi
Kelaat Mgouna Ay-5 Pluton Syen-Gr
64.06 0.68 15.29 5.66 0.21 2.23 2.43 3.76 3.93 0.15 1.49 99.89
70.79 0.36 15.32 2.63 0.00 0.47 0.18 3.84 5.12 0.12 1.09 99.92
73.67 0.24 13.83 2.34 0.00 0.65 0.13 2.62 5.19 0.10 1.05 99.82
75.94 0.08 12.90 1.49 0.00 0.14 0.10 4.13 4.52 0.07 0.46 99.83
76 47.5 15.2 23.0 20.43 105.3 331 23.31 159 8.46 0.74 1584 4.34 0.69 5.49 3.54
89 28.1 11.7 16.5 17.38 111.4 203 23.99 227 7.80 0.83 829 6.83 0.70 11.33 2.66
17 9.8 1.8 9.8 21.99 155.9 85 19.89 229 12.80 3.66 872 6.54 1.07 10.55 1.32
17 22.2 3.4 6.5 17.68 292.5 56 19.69 142 9.40 4.49 627 4.41 1.43 13.84 3.29
1 5.4 0.8 6.2 17.77 128.5 37 20.56 184 7.93 0.77 781 4.90 0.53 10.40 2.71
23.22 46.28 5.98 24.40 5.09 1.54 4.57 0.69 4.01 0.78 2.23 0.32 0.34 2.17
20.94 46.67 6.19 24.99 5.27 1.01 4.56 0.71 4.27 0.82 2.39 0.35 0.37 2.44
22.02 49.80 6.71 24.99 4.73 0.55 3.58 0.58 3.65 0.74 2.26 0.37 0.38 2.61
31.12 63.56 8.35 31.18 6.40 0.56 4.65 0.67 3.64 0.67 1.93 0.29 0.29 1.95
7.58 17.50 2.12 7.98 2.24 0.38 2.28 0.46 3.01 0.70 1.93 0.34 0.36 2.23
(wt%)
Note: Analytical laboratory of the Research Center for Petrography and Geochemistry of Nancy, France. Major elements by inductively coupled plasma emission spectrometry, and minor elements by inductively coupled plasma mass spectrometry. Analytical uncertainties are calculated to 2% for major elements, 5% for the 1000- to 100-ppm abundant elements, and from 5 to 10% for minor and trace elements. And—andesites; Dc-And—dacitic andesite; Dc—dacite; Dc-Rh—rhyodacite; Rh—rhyolite; Mdior—microdiorite; Dior—diorite; Gr musc—muscovite granite; Gr Bi—biotite granite; Syen-Gr—syenogranite. LOI—loss on ignition.
38
Pouclet et al.
The same volcanic activity took place above the different areas of the previous Pan-African domain: the former passive margin of the west African craton to the southwest (western Anti-Atlas), the suture zone of the middle part (central Anti-Atlas), and the mobile zone with the former active margin of the northeast edge (eastern Anti-Atlas) (Fig. 1). One can note, however, very reduced amounts of andesite on the western side. Petrographical and Geochemical Features In the central and eastern Anti-Atlas, the typical lithostratigraphic succession of the volcanic products is the following: More or less abundant meter-sized andesitic flows overlie detrital sediments deposited into different basinal areas. They are interbedded with and overlain by pyroclastites, epiclastites, fine sandstones, and silty shales. The lavas are mainly aphyric in texture. Compositions vary from basaltic andesite to dacitic andesite. After this first mafic activity, the volcanism rapidly evolves to acidic products. Numerous ignimbritic flows of dacitic to rhyolitic composition overcame the whole Anti-Atlas, in association with rhyolitic protrusions and massive flows. A few dacitic to andesitic flows are intercalated. In the eastern Anti-Atlas, the acidic pile and interbedded clastic sediments are crosscut by porphyritic andesite dikes. In the western Anti-Atlas, only acidic lavas have been identified. Finally, throughout the Anti-Atlas, large plutons of subvolcanic granites invaded the former volcanic formations, causing a general thermal metamorphism. The last activity consists of rhyolitic dike swarms.
Ti
Petrographically, andesites show a magmatic paragenesis more or less replaced by a greenschist metamorphic assemblage. Phenocrysts and microphenocrysts of plagioclases are partly transformed to albite and epidote. Compositions of their preserved cores range from bytownite to labradorite in basaltic andesites and andesites (An 78–50), and from labradorite to andesine in dacitic andesites (An 50–28). Pyroxenes are poorly preserved and pseudomorphosed to green amphibole. They have composition of calcic augite (Mg% 47–38, Fe2+% 11–20, Ca% 43–36) showing a Fe-enrichment trend. Their low TiO2, Al2O3, and Na2O contents are characteristic of pyroxenes from tholeiitic and calc-alkaline rocks. In the discrimination diagrams of Leterrier et al. (1982), pyroxenes cannot belong to alkaline intraplate basalts; they overlap the fields of orogenic and non-orogenic basalts (Fig. 3). No orthopyroxene has been clearly identified. Microcrysts of Timagnetite and ilmenite are common. Magmatic Mg-hornblende and biotite may be present. Phenocrysts of amphibole and biotite characterize dacitic andesites and dacites. Groundmass is entirely transformed to actinote, chlorite, albite, and epidote. Chemical compositions are frequently affected by alteration and postmagmatic processes, spilitization, weak thermal metamorphism, and hydrothermalism. Consequently, the major element contents have to be considered cautiously, mainly for alkalies, CaO, and, sometimes, MgO. The norm calculation is unavailing. Some minor elements are also questionable, such as the mobile lithophile elements. However, most of the incompatible elements can be used for magmatological purposes. The chemical features are: 51 < SiO2% < 62, 0.6 < TiO2% < 1.4, 1.3 < MgO% < 6.7, 0.2
Ediacaran andesites
0.10
V4
Cambrian basalts
Ti + Cr
V5
0.05
V4
0.08
Cambrian basalts
V5
alkali basalts (86%)
0.06
Ediacaran andesites
0.04
other basalts (92%)
nonorogenic basalts (81%)
0.03
0.04 0.02
0.02
orogenic basalts (80%)
0.01
A 0.00 0.60
0.70
0.80
0.90
Ca + Na
1.00
1.10
B
0.00 0.5
0.6
0.7
0.8
0.9
Ca
Figure 3. (A) Ca + Na vs. Ti and (B) Ca vs. Ti + Cr pyroxene composition diagram after Leterrier et al. (1982), for the andesites of the Ediacaran Atlasic Volcanic Chain and for the Cambrian basalts (V4, V5) of the Agoundis-Ounein area. (A) discriminates the alkali basalts to the other basalts; (B) discriminates the orogenic basalts to the non-orogenic basalts (straight lines after Leterrier et al., 1982).
Geodynamic evolution of the northwestern Paleo-Gondwanan margin < MgO/(MgO + FeOt) < 1.5, and, with the exception of rocks with high loss on ignition, 4.8 < Na2O + K2O < 8.6. The alkaline ratio, 0.9 < 2K2O/Na2O < 1.9, indicates a potassic tendency. The incompatible element profiles (Fig. 4A) show fractionation of the most incompatible elements, with LaN/YbN ranging from 4.6 to 17.4, enrichment of the lithophile elements, and moderate Nb and Ta anomalies (0.2 < NbN/LaN < 0.5). The associated Sr and Eu negative anomalies can be due to plagioclase fractionation. Compared to continental arc-related andesites represented by averaged andesites of Chile (Fig. 4B), the Anti-Atlas andesites are more enriched in the most incompatible elements and are less anomalous in Nb-Ta. They resemble intracontinental andesites, such as those from Erciyes and Ararat, two volcanic centers in Turkey (Kürkçüoglu et al., 1998; Yilmaz et al., 1998; Fig. 4B). The acidic volcanic rocks consist of dacites and rhyolites characterized by phenocrysts of alkaline feldspaths and quartz. Microphenocrysts of clinopyroxene and phenocrysts of biotite and hornblende are common in dacites. Ignimbritic rocks show an usual eutaxitic structure. The plutonic rocks are composed of two main petrographical types: (1) diorite, biotite monzogranite, and K-feldspar–rich granite; and (2) biotite and primary muscovite granite, sometimes rich in garnet, tourmaline, and rare cordierite. Magmatic minerals are highly altered to actinote, epidote, chlorite, and oxides (Fe-Mg minerals), and to sericite and albite (feldspars). The silica contents range from 63 to 75%. Alkalies are high, 7.3–10.5%, with a ratio of 2K2O/Na2O up to 1.2, indicating a high potassic feature. The incompatible element contents are in the range of those of the andesites, with a low to moderate fractionation, 2.3 < LaN/YbN < 13.1, and moderate negative Nb-Ta anomalies, 0.3 < NbN/LaN < 0.7 (Fig. 4C). Thus the acidic lavas cannot have evolved from andesitic magmas by simple differentiation, although high Sr, Eu, and Ti negative anomalies are due to mineral fractionation. Few selected cogenetic plutonic rocks are analyzed for comparison. Diorites share the same chemical patterns with andesites and dacites, and biotite- or muscovite-bearing granites, with rhyolites. The magmatologic features are shown in the Ta/ Yb versus Th/Yb diagram adapted from Pearce (1982; Fig. 5). Andesites and dacites overlap the calc-alkaline and shoshonitic area. Some dacites and rhyolites are significantly enriched in Th and mildly in Ta. The chemical trend is intermediate between the crustal contamination and the fractional crystallization trends and can be explained by assimilation and fractional crystallization (AFC) combined processes. A similar interpretation is proposed for the mafic and acid volcanic rocks of Turkey, for example, the eastern Anatolian volcanic suites (Yilmaz et al., 1998; Fig. 5). Probable Geotectonic Setting The volcanic and sedimentary formations of the Ediacaran Atlasic Volcanic Chain unconformably overlie the formations of the mobile belt, which were folded and metamorphosed during the Pan-African collision. Thus the volcanic chain formations postdate the collision. They also overlie the late orogenic dioritic to granitic plutons of the mobile belt. They were folded and
39
tilted before the Cambrian transgression. They have undergone a moderate thermal metamorphism caused by the magmatic activity and granite emplacement. The calc-alkaline signature of the magmatic products may suggest an active continental margin setting. But one must note the lack of any suture zone for any previous oceanic area, which would be parallel to the chain. Moreover, there are no subduction-related marine sediments and no tectonic and metamorphic witnesses of a major convergence that would have led to the vanishing of an oceanic domain. Structurally, a continental intraplate setting is the most probable context. The petrographic (e.g., pyroxene composition) and magmatologic features of the volcanics agree with such a setting. The calcalkaline affinity can be explained by melting of a previously contaminated mantle by earlier subduction and collision events, and also by crustal assimilation and hybridization, as shown by the high-potassic acidic rock abundance. Indeed, the Sr-Nd isotopic data obtained for the Imiter inlier Ediacaran rocks indicate a mixing of mantle and lower crustal sources (Gasquet et al., 2005). A modern equivalent is the Anatolian volcanic system of Turkey. LATE EDIACARAN TO EARLY CAMBRIAN MAGMATISM In the Atlasic northwestern area, southwest of the old block of High-Atlas, the Adoudounian formations are conformably deposited above a basaltic pile. We carried out complete crosssections of this pile and observed a clearly unconformable contact of the base of the basaltic stack above the rhyolitic complex of the Late Neoproterozoic volcanic chain (Ediacaran Atlasic Volcanic Chain). Subsequently, an extended sampling has been performed for the volcanic products of the lower basaltic pile and also of the Early Cambrian strata, in the Agoundis-Ounein region. Detailed investigations of the structural relationships between the basaltic pile and the underlying rhyolitic complex have been done in the Toubkal massif. Agoundis-Ounein Region The Agoundis-Ounein region is located southeast of the western High-Atlas, at the edge of the old block of High-Atlas (Site 2, Fig. 1; Fig. 6). The earlier Cambrian sediments overlie the volcanic and the granitic formations attributed to the Ouarzazate Supergroup and known as the Ouzellarh “promontoire” (Choubert, 1963). In the upper valley of the Agoundis River and east of the Ounein plain, the transgressive Cambrian sediments appear to be conformable above a pile of basaltic flows (Fig. 2). A geological map of this region has been done (Fig. 7). The lithological succession is shown in a log (Fig. 8). Along the Agoundis valley, strata are dipping 30°–40° to WNW in a normal monoclinal succession, a structural feature gained during the Variscan tectonic event. The same old Cambrian formations crop out in the eastern part of the area. On the southeast side, they thrust themselves with transport to the northwest. In the south, they thrust, with important reverse folding, above the Cretaceous sediments. These thrust
1000
Pouclet et al.
Rock / Primitive Mantle
40
A
100
10
1 Rb Ba Th Nb Ta La Ce Pr Sr Nd Zr Hf Sm Eu Gd Ti Dy Y Yb 1000
B Erciyes
Rock / Primitive Mantle
Ararat
Figure 4. Incompatible element diagrams normalized to Primitive Mantle for Late Neoproterozoic rocks of the Ediacaran Atlasic Volcanic Chain. (A) Mafic rocks; (B) comparison with Chile andesites and Turkey andesites. Average composition of Chile andesites after Thorpe et al. (1984) and Hickey et al. (1986). Erciyes and Ararat compositions after Kürkçüoglu et al. (1998) and Yilmaz et al. (1998). (C) Acidic rocks with addition of three eastern Saghro rhyolite analyses after Levresse (2001). Normalization values after Sun and McDonough (1989).
100
Anti-Atlas andesites
10
Chile andesites 1 Rb Ba Th Nb Ta La Ce Pr Sr Nd Zr Hf Sm Eu Gd Ti Dy Y Yb 1000.0
Rock / Primitive Mantle
dacites rhyolites granites
C
100.0
10.0
1.0
0.1 Rb Ba Th Nb Ta La Ce Pr
Sr Nd Zr Hf Sm Eu Gd Ti Dy Y Yb
10.00
Rh-Gr
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
Calc-alkaline basalt
41
Upper crust
DacAnd
Figure 5. Ta/Yb vs. Th/Yb diagram after Pearce (1982) for the andesites and dacites (Dac-And), rhyolites, and granitoids (Rh-Gr) of the Ediacaran Atlasic Volcanic Chain magmatic suites. EA—Eastern Anatolian volcanic suites after Yilmaz et al. (1998). Control trends of magmatic evolution: AFC—assimilation and fractional crystallization; CC—crustal contamination; FC—fractional crystallization; SZ—subduction zone. Upper crust composition after Taylor and McLennan (1985).
Ocean Island Basalt EA
A rra y
1.00
an t
le
SZ
CC AFC FC
M
Th/Yb
Shoshonite
Primitive Mantle
0.10 0.01
0.10
1.00
10.00
Ta/Yb
8°
N
Wirgane
Imlil
Toubkal
Ifni Lake Amsozerte
31°
31°
AGOUNDIS Ijoukak
OUNEIN 10 km
8° Post-Cambrian formations Ediacaran rhyolitic rocks
Cambrian sediments Ediacaran granitic rocks
Late Ediacaran (?) basalts Neoproterozoic volcano-sediments
Toubkal feeding dikes stratification
Figure 6. Sketch geological map of the Toubkal Massif and Agoundis-Ounein areas. The Toubkal dike system corresponds to the feeding dikes of the overlying Late Ediacaran to Early Cambrian basaltic flows. Heavy lines are major faults.
42
Pouclet et al. Assif n'Fis
8°10
8°5
Ijoukak AÔt Moussa
31°
Rh
31°
Wankrim Assi
f
N
Tourkout
S1
V1
S2
Taghbar Ago und
2 -V
is
SL V3
C1
C2-S3-C3-S4 Wijdane Talat n'Ou Lawn
V4
Lz
30°55
30°55
V5
t
la 'Ta
Sr
Agadirane
n rar Ad
Sp
C1
S5 Assi
f
Tamsaloumt
a ufr
n'O
ra
uf
a
Assi
f
OUNEIN
n'Ou
ar dr
Cg
l fal u O
n'O
n'
A
Agd
im
0 8°10
5 km
8°5
Sp
S5 V5 Cg
C2-S3-C3-S4 SL V3
V4
S2-V2
C1 S1 V1 Rh
Figure 7. Geological map of the Ounein-Agoundis area. V1, S1, C1, S2, V2, V3, C2, S3, C3, S4, V4, Cg, V5, and S5 are Late Ediacaran to Middle Cambrian Formations. Lz—lherzolite slices; Rh—rhyolitic-dacitic complex of the Ediacaran Atlasic Volcanic Chain; Sp—spilitic chimneys; Sr—serpentinite and ophicalcite beds. Arrows indicate observation sites of rhyolitic complex and discordant contacts with the V1 basaltic pile. Assif means “river.”
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
Stratigraphic Formations
tectonics resulted from the Alpine-Atlasic shortening. The rhyolitic substratum appears at the base of the southern thrust. The upper Early Cambrian and Middle Cambrian formations crop out to the west and to the southwest of the Ounein area.
Magmatic events
S5 laccoliths and spilitic flows V5 Cg
basalt flow V4b ophicalcite and lherzolitic breccia
S4
spilitic chimneys and flows V4a
lherzolite slices
C3 S3
C2
laccoliths V3
S2 pyroclastic flows V2 C1
100 m
S1
V1
43
basaltic pile V1 rhyolitic complex
Figure 8. Stratigraphic column of the Agoundis-Ounein area.
Stratigraphic Succession Above the rhyolitic complex, we distinguish five volcanic occurrences or units, V1–5, five detrital formations, S1–5, three dolostone and limestone formations, C1–3, and one main conglomerate, Cg. The volcanic Unit V1, 400–500 m thick, is a stack of metersized underwater basaltic flows. It overlies the rhyolitic complex, with a clear erosional unconformity that we observed at two sites, in the upper Agoundis valley and east of Agadirane (arrows in Fig. 7). The structural setting of the massive rhyolite bodies is not determinable, but in the upper Agoundis valley, intercalated deposits of epiclastites, volcanic breccia, and polymictic conglomerates (subrounded and angular elements of rhyolites, ignimbrites, dacites, andesites, and granites) dip 45° to the southwest. As the overlying basalt flows dip 30° to the northwest, an angular unconformity is evidenced. The base of the basaltic pile consists of a heterogeneous pyroclastic bed of variable thickness (50–150 cm) deposited above a highly altered and eroded rhyolitic material. This bed includes fine fragments and centimeter-sized angular blocks of scoriaceous basalts, rhyolites, ignimbrites, quartz, meta-argilosiltites, and meta-silty sandstones. It is overlain by a plurimeter-sized pyroclastic flow of basalt and by the stack of massive flows. The lower rhyolitic and basaltic formations of the discordance zone are affected by three sets of faults in the following order: (1) N110° normal faults straightened and reworked in right-lateral wrench faults, (2) N20° reverse faults and thrusts dipping 35° to WNW, and (3) vertical N80° right-lateral strike-slip faults. Some hectometer-scale slices of highly altered rhyolitic material are intercalated in the lower basaltic pile because of the N20° reverse faulting. Some basaltic dikes, having the same composition as the flows, crosscut the rhyolite complex and the basal pyroclastic layers but not the overlying sediments. They are 50-cm to 1-m wide and trend N80°, N90°, N110°, N120°, and N170°. The basaltic stack is made of 1.50- to 3-m-thick massive flows and of few intercalated pyroclastic flows. The top of the unit consists of plurimetersized beds of pyroclastites. No pillow lavas have been observed. However, the base and the top of each flow are vitrified and fragmented into hyaloclastic layers. Typical underwater quench textures are observed. Lavas are poorly vesicular except in the upper part of the pile. Spilitization products (lodes and veins of epidote, chlorite, and calcite) and development of celadonite are common. All these features indicate an underwater flow above a subhorizontal or gently sloping floor. Basalts are either microlitic porphyritic or highly phyric, with aggregates of large plagioclase. Magmatic minerals are phenocrysts of olivine replaced by saponite and mantled by secondary magnetite, more or less albitized plagioclase, clinopyroxene pseudomorphosed to calcic amphibole, microcrysts of Ti-magnetite (15.0 < TiO2% < 18.9),
44
Pouclet et al.
and rare ilmenite. The groundmass is wholly transformed to secondary minerals (chlorite, smectite, epidote, albite, and calcite). The porphyritic plagioclase-rich facies prevailed in the middle and upper piles. Preserved plagioclase compositions range from An 60.9 to An 50.6 for the phenocrysts, and to An 37.9 for the microcrysts. The order of crystallization—olivine, plagioclase, and clinopyroxene—corresponds to olivine tholeiites. Chemical compositions are presented below.
100
Lherzolite spinel V4 xenolith spinel
90
80
100 Cr / (Cr + Al)
70
Stratiform complexes
60
Ophiolites
50
40
Abyssal peridotites
30
20
Galicia Margin
10
0 0
20
40
60
80
100
100 Fe2+ / (Fe2+ + Mg) Figure 9. Diagram of 100Fe2+/(Fe2+ + Mg) vs. 100Cr/(Cr + Al) for spinels of the lherzolite and of the V4 xenolith of the Ouniein area. Spinel composition ranges of ophiolites, abyssal peridotites, and stratiform complexes after Dick and Bullen (1984), Irvine (1967), and Duke (1983). Spinel composition of the Galicia Margin peridotites after Evans and Girardeau (1988). The lherzolite spinels are close to those of the Galicia Margin.
The formation S1, 140 m thick, starts with a conglomerate of rounded pebbles of basalt in a volcanosedimentary matrix of epiclastites, fine sand, and clay. Above there is a succession of green to purple siltstone and claystone beds. Interbedded centimeter-sized layers of a fine siliceous carbonate are added in the upper part. Carbonate deposition prevailed in the C1 formation, which is 70 m thick, with alternating layers of siliceous dolostones and of clay-rich dolomitic limestones. Planar-shaped stromatolitic structures are common. Detrital deposition renews in the S2 formation, which is 250 m thick. The base of S2 is rich in volcanic products, the V2 occurrence, with thick conglomerates of basaltic and chert pebbles in a volcaniclastic matrix, hyaloclastite layers, and spilitic flows. The overlying and main member of the formation is made of alternating thin layers of green to purple siltstones and claystones. In the upper part, interbedded dolomitic limestone layers are added, as are many centimeter-sized layers of black phtanite. The C2 formation, 200 m thick, is mainly a thick-bedded dolomitic limestone. The base is characterized by a plurimetersized deposition of a matrix-supported limestone conglomerate. In the massive limestone, micritic layers are rich in Archeocyatha. Magmatic occurrences of doleritic to micrograined dioritic rocks appear as typical concordant 2- to 10-m-thick lenticular laccoliths. These reservoirs belong to a slightly younger volcanic activity we named the V3 occurrence. Five of these laccoliths outcrop in both sides of the Agoundis valley (SL V3, Fig. 7). Their lower and upper margins are quenched and show a hyalo-phyric texture, all minerals being highly altered. Texture of the inner rock is porphyritic micrograined with abundant phenocrysts of albitized plagioclase, some phenocrysts of clinopyroxene, and amphibole replaced by chlorite, in a matrix of the same minerals plus Ti-magnetite, ilmenite, titanite, and interstitial quartz. Chemical compositions are discussed below. The S3 formation is characterized by an increasing number of silty clay layers alternating with limestone beds. Then renewal of calcareous deposition constitutes the C3 thick-bedded limestone formation. The S3 and C3 total thickness is poorly estimated, because of complex normal and reverse faulting, but it exceeds 250 m. The prominent feature of the C3 formation is the abundance of meter-sized domed stromatolites. Also very important is the discovery of slices of moderately serpentinized peridotites along southwest–northeast-trending faults. The peridotite is a spinel lherzolite made of olivine Fo 91.4–89.5, orthopyroxene (Mg% 89.6–88.5, Fe% 10.5–9.2, Ca% 1.8–0.9), clinopyroxene (Mg% 47.7–45.7, Fe% 4.5–3.9, Ca% 49.7–47.7), and magnesian hercynite spinel (Sp% 70.7–55.0, Hercy% 25.7–18.4, Mg-Cr% 15.2–6.9, Cr% 6.3–2.0) that is poor in TiO2. Only a few fine spinel grains are Cr-rich (Sp% 37.0, Hercy% 30.6, Mg-Cr% 17.3, Cr% 14.3; Fig. 9). Rare pentlandite crystals are associated with secondary magnetite in the serpentine veins. A weak thermal effect was responsible for development of aggregated fine tremolite prisms (Mg/[Mg + Fe] 0.92–0.98, Al2O3% 2.8–4.5). The compositions of pyroxenes with high octahedral Al contents in orthopyroxene
Geodynamic evolution of the northwestern Paleo-Gondwanan margin (OPX) and clinopyroxene (CPX), and high Cr contents in CPX (Ko% 3.1–3.6 and Jd% 5.4–9.2 in the solid solution percentages) correspond to mantle peridotite minerals. High Na2O and Al2O3 contents of the CPX (1.2 < Na2O % < 1.8, 5.9 < Al2O3% < 6.9) are characteristics of subcontinental lherzolites (Kornprobst et al., 1981). However, the positive correlation displayed between Na and Cr is distinct from that of the continental peridotite CPX. The same correlation is shown by the CPX of the Galicia Margin peridotites, representative of transitional upper mantle between continental and oceanic domains (Evans and Girardeau, 1988). Low TiO2 contents and low Cr ratios of the spinel agree with a continental mantle peridotite rock composition. The Fe-Mg and Cr-Al number ranges (100Fe2+/[Fe2+ + Mg] 22.0–31.3, 100Cr/[Cr + Al] 8.9–21.7; Fig. 9) are similar to those of spinels of the Galicia Margin peridotites (Evans and Girardeau, 1988). Limestone conglomerate and breccia ended the carbonate deposition. Detrital beds are overlain by alternation of decimetersized beds of clay, silty clay, and argillaceous limestones constituting the 200-m-thick S4 formation. The middle part of this formation is characterized by an intense effusive spilitic activity, with chimneys and flows, which constitutes the V4a volcanic phase. Five main chimneys, 2–8 m in diameter, have been observed. They perpendicularly intrude the sediment beds, which are straight up and invaded by calcite, chlorite, and epidote veins and blobs. The interbedded flows, 1–3 m thick, are based and topped by hyaloclastites, all being highly spilitized. The center part of the thickest flows show a hyalopilitic porphyritic texture, with some microphenocrysts of olivine replaced by saponite and Mg-chlorite, more or less abundant phenocrysts of clinopyroxene (Mg% 41.1–46.9, Fe% 8.9–15.7, Ca% 43.2–44.8) and plagioclase labradorite, and microcrysts of Ti-magnetite (5.9 < TiO2% < 16.9) in a groundmass made of secondary minerals (albite, chlorite, epidote, celadonite, and calcite). No mantle xenoliths have been found. However, in a centimeter-sized enclave of quartz, we analyzed a crystal of chromhercynite spinel (Fig. 9), the origin of which is questionable. In the upper S4 unit, clastic limestone or calcarenite decimeter-sized beds were deposited in the silty clay, but also a few layers, half a meter thick, of brecciated ophiclacites rich in serpentinite blocks. These blocks recall the C3 lherzolite slices, but the rock is highly crushed, serpentinized, and carbonated. Petrographic and chemical determinations are not possible. The top of the S4 formation consists of two or three thick and massive basaltic flows, the V4b volcanic phase. These flows are devoid of any underwater outpouring witness (lack of spilite products, hyaloclastites, and quenching textures), indicating a shallow water or subaerial flow. Coarse columnar jointing developed in the thickest flow. Texture is commonly microlitic phyric but is doleritic subophitic in the center of the flow. Magmatic minerals consist of microphenocrysts of olivine and Timagnetite (10.2 < TiO2% < 11.4), phenocrysts of clinopyroxene (Mg% 46.0–38.2, Fe% 7.9–20.3, Ca% 39.7–47.3), plagioclase (An 59.3–48.5), and rare titanite in a groundmass of microcrysts of magnetite, ilmenite, clinopyroxene, and plagioclase. Apart from the spilitization process for the V4a lavas, the two volcanic
45
group rocks V4a and V4b exhibit the same petrologic features, with the following order of crystallisation: olivine, clinopyroxene ± plagioclase ± magnetite, and plagioclase, corresponding to an olivine-tholeiite or moderately alkaline basalt magma. In the discrimination diagrams of Leterrier et al. (1982), pyroxenes indicate a tholeiitic to alkaline magmatic affinity, typical of transitional suites, and a non-orogenic basalt signature (Fig. 3). A 5- to 30-m-thick conglomerate, Cg, overlies the eroded top of the V4b lava flows. It includes subrounded to well-rounded pebbles and sand-sized fragments of basalts, cherts, silexites, and limestones in a poor silty clay matrix. Then deposition rapidly evolves from coarse to fine greenish siliceous sand constituting the lower part, around 100 m thick, of the S5 formation. The rest of this formation consists of an alternation of fine sand and silty clay, with upward-increasing of the clay component. The lower S5 formation contains witnesses of a moderate volcanic activity: lava flows and laccoliths of the V5 event. The lava flows are spilitized and exhibit the same lithologic and petrographic features as the V4a flows. A few small laccoliths or sills are emplaced just above the conglomerate Cg. The rocks are fine grained and show a basaltic doleritic to dioritic petrographic composition. A large laccolith, of decameter thickness, widely outcrops to the west. Its base and top are severely quenched, and a thermal effect welded the sandstones. The inner rock is microdioritic, with a composition close to that of the V3 laccoliths. Texture is porphyritic micrograined with abundant large phenocrysts of plagioclase (An 49.5–34.0), some phenocrysts of clinopyroxene (Mg% 40.7–43.9, Fe% 20.7–14.3, Ca% 41.8–38.2), and amphibole replaced by chlorite, in a matrix of the same minerals plus Timagnetite (9.2 < TiO2% < 11.4), Mn-rich ilmenite (1.6 < MnO% < 3.4), titanite, and interstitial quartz. Pyroxenes are slightly more evolved than those of the V4 lavas. The Ti, Cr, and Ca + Na contents are consistent with the evolved trend (Fig. 3). Some rounded small xenocrysts of garnet have been analyzed. They contain tiny inclusions of K-feldspar and pyrrhotite. They are mantled by a reaction rim of mixed quartz, actinote, and ilmenite. Their composition (almandine % 57.4–75.8, pyrope % 11.6–33.1, grossular % 1.9–8.9, spessartine % 0.3–3.7) may correspond to garnets of high-grade gneisses. To sum up, the sedimentary and volcanotectonic evolution can be interpreted as follows. The basal major volcanic pile V1 is related to a major volcanotectonic event. Underwater basalts outpoured from fissural system above a subhorizontal eroded ground made of continental-set rhyolitic material. Erosional and angular unconformities are evidenced. The initial tectonic setting was partly reworked by the Variscan folding and the Atlasic faulting. However, the N20° faults seem to be inherited from primitive structures. Also, the basalt dike trends are not well understood, because of the lack of extended observations. Dikes may set either along faults or along oblique tension gashes. At the end of the lava supply, the basin floor subsided, possibly in response to the reservoir emptying. Detrital accumulations (S1) progressively fill up the basin and a shallow-water carbonate deposition
46
Pouclet et al.
developed (C1). A new moderate volcanotectonic event (V2) is associated with a new basin deepening and detrital deposition (S2). Again, the sedimentary filling-up permitted the development of a carbonate platform (C2–C3) in a marine transgression context (basal conglomerate). A subsidence process or a climatic variation may explain the intermediate detrital alternation (S3). A third important volcanotectonic phase took place. It was announced by the mantle peridotite slice occurrences, which strongly suggest important crustal extension and mantle denudation. The laccolith intrusions (V3) are attributed to this phase. Then a deep-water condition prevailed for detrital deposition (S4) and volcanic activity (V4a). Reworking of mantle material attests to continuation of the extension and the associated magmatic production. Above, increasing amounts of limestone beds indicate increasingly shallow water until a possible emergence with subaerial lava flowing (V4b). A drastic change is shown by the thick coarse conglomerate of rounded pebbles (Cg) linked to new marine transgression. However, the volcanic activity continued (V5). This activity declined progressively, while a very thick, fining-upward and probably deep-water detrital sedimentation (S5) filled the basin. These sediments widely outcrop south of the Ounein plain. No more witnesses of volcanic products have been observed in this southern area. But farther to the north, in the late Early and Middle Cambrian formations of the WesternMeseta, rare intercalated volcanic rocks are known. They display the composition of continental tholeiites and alkaline basalts (Ouali et al., 2000, 2003; El Hadi et al., 2006). Regional Correlations and Datings V1 was formerly included in the Ediacaran Ouarzazate Supergroup, as an andesitic or basaltic activity continuing the rhyolitic activity, in the western High-Atlas and western AntiAtlas (Proust, 1973; Azizi-Samir et al., 1990; Piqué et al., 1990, 1995, 1999; Piqué, 2003; Table 1). The observed disconformity of V1 above the rhyolitic complex and the conformity with the overlying Cambrian units allow us to place the V1 basaltic activity at the beginning of the “Cambrian” evolution story. This activity occurred after the tectonic event that affected the Ouarzazate Supergroup in the old block of High-Atlas and in the central and eastern Anti-Atlas. The youngest products of this supergroup are dated to 550 ± 3 and 552 ± 5 Ma in the eastern Anti-Atlas (Levresse, 2001; Gasquet et al., 2005) and to 559 ± 6, 562 ± 5, and 565 ± 7 Ma in the western Anti-Atlas (Thomas et al., 2002; Walsh et al., 2002). Consequently, the V1 activity is younger than 550 Ma. The sedimentary formations above the V1 basalts are well known in the central and western Anti-Atlas as the Taroudant and Tata Groups (Thomas et al., 2004). Stratigraphic correlations are given in Table 1. S1 corresponds to the Adoudounian detrital Basal Formation. Occurrences of conformably set basaltic flows at the base of this formation are known in the Siroua area of the central Anti-Atlas and the northeast and west of the western Anti-Atlas (Leblanc, 1977; Demange, 1980; Algouti et al., 2001; Bajja, 2001; Benziane et al., 2002; Soulaimani et al.,
2003, 2004). Their thickness rapidly decreases to the south and the southwest. In other places, the Adoudounian Basal Formation overlies the Neoproterozoic terranes, either unconformably in the central Anti-Atlas (Sites 3, 4, and 5; Fig. 2), or conformably in the southwest of the western Anti-Atlas (Site 1; Fig. 2), as discussed above. We may conclude that the V1 basaltic pile, as well as its S1 overlying detrital deposits, is correlated with the Adoudounian Basal Formation. C1 corresponds to the Tamjout dolostone, massive beds of which constitute a famous lithostratigraphic marker. It is at the base of this formation and to the northern edge of the western Anti-Atlas, that Maloof et al. (2005) determined a –6‰ nadir in the δ13C record, which they attributed to the Ediacaran-Cambrian boundary dated at 542.0 ± 0.6 Ma (Amthor et al., 2003). If it is true, the Adoudounian Basal Formation, including the V1 basaltic formation, is dated to the late Ediacaran. Northwest of the Bou Azzer inlier, important volcanic activity of the Djbel Boho (or Alougoum) occurred during the dolostone deposition. Andesitic to trachytic lavas have been described (Choubert, 1963). Analyses of interbedded flows, sampled in the middle part of the dolostone formation, reveal an alkaline composition (see below). The U-Pb datings of two rocks yield 534 ± 10 and 533 ± 2 Ma (Ducrot and Lancelot, 1977; Levresse, 2001). These ages are recalculated to 529 ± 3 and 531 ± 5 Ma by Gasquet et al. (2005). Consequently, the C1, or Tamjout formation, is dated to the Early Tommotian Stage, but also to the Nemakit-Daldynian Stage, if its base is at 542 Ma. It spent more than 10 Ma in this stage, a rather long time. Another possibility is that the δ13C low-value excursion observed at the base of the Tamjout Formation is not the one that is considered to mark the Ediacaran-Cambrian boundary. S2 corresponds to the “Lie de Vin” Formation. The U-Pb zircon ages of volcaniclastic material related to the V2 activity give 525.4 ± 0.5, 522 ± 2, and 521 ± 7 Ma (Compston et al., 1992; Landing et al., 1998; Maloof et al., 2005). However, Archaeocyatha of the overlying limestone beds C2 (the upper limestone formation) are dated to the Atdabanian (Debrenne and Debrenne, 1978). If it is true, because the Atdabanien is defined between 530 and 525 Ma (Tucker and McKerrow, 1995), the oldest range of the age-dating uncertainty must be taken. S3 and C3 correspond to the “Schisto-calcaire” Formation, and S4 to the “Série schisteuse.” A U-Pb zircon age of 517 ± 1.5 Ma was obtained for an ash bed of the latter formation (Landing et al., 1998). Consequently Cg and S5 can be dated to the Middle Cambrian, which is demonstrated by the trilobite fauna (Boudda and Choubert, 1972; Sdzuy, 1978; Geyer, 1998). Toubkal Massif The Toubkal Massif is a thick stack of basaltic rocks (up to 500 m) that constitutes the higher mountains of the High-Atlas, culminating at the Toubkal summit (4167 m; Fig. 6). The upper part of the pile and the overlying deposits are missing in this area. Fortunately, the base of the basalts above the rhyolitic complex can be extensively seen in all the deep valleys of the massif.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin Previous investigations of these volcanic rocks are due to Zahour et al. (1999) and Zahour (2001). The rhyolitic lavas are attributed to a postcollisional calc-alkaline volcanism, whereas the basalts are related to an intraplate continental activity. Lithostratigraphy The lithostratigraphic succession is exposed along a southeast-northwest cross-section of the Toubkal Massif, from Amsozerte to the upper valley of Imlil (Fig. 10). The Neoproterozoic formations consist of a plutonovolcanic acidic complex made up of granitoid plutons and dacitic to rhyolitic lava domes and flows. To the southeast, the Amsozerte pluton is a quartz-diorite comprising amphibole, biotite, andesine, K-feldspar, and quartz. It shares calc-alkaline composition (Table 2, Fig. 4) with granodioritic and Bi-granitic batholitic plutons that constitute most of the Siroua inlier. Numerous microdioritic bodies crop out in the Ifni Lake area, between the diorite and the overlying dacite unit. They are considered as apophyses of the diorite pluton. The >1,000-m-thick dacitic and rhyolitic unit is an association of lava domes and massive flows. All the massive flows are capped by pyroclastic flows of typical eutaxitic textured ignimbrites. No significant structural oriented pattern can be seen in the massive flows. In compensation, the platy cleavage parallel to the final cooling of the ignimbrites shows a 30°–40° post-setting dip of the pile, to the northeast. Very abundant basaltic dikes crosscut the acid volcanics. They belong to a N30° trending swarm. They fed the overlying basaltic pile, as shown by relationships between dike flows and
by their petrographic and chemical likenesses. Dike thickness ranges from 2 to 15 m. In the upper valley of the Ifni Lake, south of the Toubkal Mountain, dikes coalesce at their bases to give small chambers, but they are thinning upward. In that area, they may constitute 50% of the country rocks. Direction varies from N10° to N45°, being most frequently N30°. All the dikes are dipping 60° to the southeast. The chilled margin is 5–20 cm wide, depending on the dike thickness. About 65% of the dikes exhibit a composition of porphyritic basalts rich in plagioclase phenocrysts. The others are aphyric basalts. Contact between ignimbrites and the overlying basaltic stack is sharp and almost flat, dipping 30° to the northwest. The overbedded basaltic units have the same dip, except close to NNW-SSE faults. The basal unit is a 1- to 30-m-thick brecciated agregate of rhyodacitic fragments and scoria-sized basaltic pyroclasts. It is overlain by a 20- to 50-m-thick basaltic pyroclastic flow accumulation, and then by the upper stack of thick lava flows. South of Imlil, just below the basaltic pile, we observed a few meter-wide sedimentary channels, filled by quartz-rich fine sandstones and by detrital deposits of mixed lapilli-sized fragments of basalt pyroclasts and quartz grains in a silty matrix. This deposit is a witness to the reworked mixed formation of the substratum at the time of the basaltic outpouring. The upper slopes of the valleys and all the summits of the Toubkal Massif are made of stacked basaltic flows similar to those of the Agoundis-Ounein V1 unit. Development of celadonite and the occurrence of intercalated beds of hyaloclastites indicate an underwater flowing. However, no pillow lava structures were observed.
NW
SE Toubkal
Ifni Lake edges
3000
Amsozerte
Assif n'Isougouane
4000
47
2000
Late Ediacaran (?) basalt flows Ediacaran ignimbrites and dacites-rhyolites
basalt and microgabbro dykes
Ediacaran microdiorite and diorite Figure 10. Cross-section in the Toubkal Massif from Amsozerte (southeast) to the northwestern valley of the massif (see Fig. 6). Note the WNW 30° tilting of all the formations. Vertical axis in meters. Heavy lines are major faults.
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Pouclet et al.
Structural Insights The main tectonic feature of the Toubkal Massif is the N60°trending major strike-slip faulting. These faults are part of the Tizi n’Test and south Atlas fault system, which separates HighAtlas and Anti-Atlas (TNT, SAF, Fig. 1). Displacements along this system were right-lateral during the Variscan orogeny and left-lateral in the Mesozoic pre-Atlasic stage (Proust et al., 1977). Before the Permian time, the crustal substratum experienced a 30° tilting to WNW (Fig. 10). The last Atlasic compression caused reverse faulting of the WSW-ENE shear zones and thrust of the Cretaceous sedimentary cover. The Variscan and Atlasic events reworked a previous system of northeast-southwest and northsouth normal faults, northwest and west dipping. Reverse and dextral motions are observed. However, the initial normal fault position can be seen along the important N45° trending faults, south and north of the Toubkal summit. The normal movement was subcontemporaneous with the basaltic activity, as shown by the thick accumulation of pyroclastic material at the northwest lower side of the faults. These synvolcanic tectonic features of normal faults associated with the N30° dike swarm indicate a WNW-ESE extensional regime. Age Data A geological mapping of the Siroua area including several geochronological data points has been done recently (Thomas et al., 2002). Late- or post-orogenic diorites and granodiorites of the Siroua northeastern batholith yield U/Pb ages of 586 ± 8, 579 ± 7, and 575 ± 8 Ma. Neighboring rhyolites and associated granites belonging to the Ediacaran Atlasic Volcanic Chain are slightly younger: 571 ± 8, 562 ± 5, and 559 ± 6 Ma. In the HighAtlas (old block), a rhyolite, below the easternmost occurrence of the flood basalts, is dated at 578 ± 5 Ma (Mifdal et al., 1982). These data give the age of the substratum of the basaltic pile. To sum up, after the late Neoproterozoic intracontinental magmatic activity, deformation of the substratum is followed a new extensive fracturing. A N30°-trending fissural basaltic activity took place in a WNW-ESE extensional regime. Lava flows outpoured in a shallow water environment and in a nearly flat topography, except for the northeast-southwest cliffs facing northwest, as shown by the paucity of basal detrital sediments. Geochemical Composition of the Late Ediacaran to Early Cambrian Magmatic Products Analyses of the V1–V5 volcanic rocks of the AgoundisOunein and Toubkal areas are given in Table 3. The magmatic products consist of basaltic rocks in lava flows, dikes, chimneys, and plugs (V1, V2, V4, V5) and of doleritic to evolved dioritic rocks in laccoliths (V3, V5). For the basaltic rocks, SiO2 contents range from 46.3 to 61.3%, TiO2 from 0.9 to 2.8%, and MgO from 2.0 to 8.9%. Mg number values (100MgO/[MgO + FeO]), which range from 24.5 to 48.8, correspond to evolved basalts. Seawater alteration and spilitization effects have modified alumina, lime, and alkali abundances (as shown by celadonite and
albite developments). Alkali ratio values are unavailing, as is the norm calculation. For the laccolithic evolved rocks, SiO2 contents range from 61.8 to 65.1%. The alkaline ratio, 1.3 < 2K2O/Na2O < 1.4, indicates a potassic trend. The hygromagmaphile patterns are pointed out in Figure 11. Basalts of the lower pile V1 are moderately fractionated, with LaN/YbN ratios ranging from 3.5 to 10.1. The most evolved basalts display higher contents of incompatible elements. The lithophile elements are also enriched. Several anomalies are noticeable. They are negative for Nb and Ta compared to La (0.3 < NbN/LaN < 0.7); more or less negative for Sr; and slightly negative, in a few rocks, for Eu and Ti. Associated Sr and Eu negative anomalies can be due to plagioclase fractionation; they are significant in the most evolved rocks, but also in some mafic rocks. The V2 lavas display the same pattern. Compositional range of the V4 and V5 basalts is similar to that of the V1 (Fig. 11B), with 3.0 < LaN/YbN < 5.1 and 0.4 < NbN/LaN < 0.7. Some Ba and Sr contents have been disturbed by spilitization process. The V3 and V5 microdiorites are more fractionated and enriched in lithophile elements (10.2 < LaN/YbN < 11.3), but not in the high field strength elements; this characteristic precludes a simple differentiation of the evolved rocks by fractional crystallization. However, strong Sr and Ti negative anomalies attest to crystal fractionation. A complex AFC process can be suggested, when considering the Ta/Yb versus Th/Yb diagram (Fig. 12). The mafic lavas may have originated from a mildly enriched mantle source close to that of the continental tholeiites. The evolved microdiorites experienced a crustal contamination. Witnesses of such a contamination could be the garnet-bearing xenoliths of granulitic gneisses of the V5 lavas. Consequently, one may assume the presence of a continental crust below the Ounein area. Low incompatible element enrichments and moderate Nb-Ta negative anomalies of the more mafic V1 basalts are in good agreement with the average continental tholeiite profile (Fig. 11C). Such a magmatic signature suggests a continental lithospheric source. Compared to the Ediacaran Atlasic Volcanic Chain andesites, the V1 basalts display weak but significant differences. As a whole, andesites are more enriched in the most lithophile elements, more anomalous in Nb and Ta, and depleted in the heavy rare earth elements (Fig. 11C). Some data are available in the literature for the Early Cambrian lavas of the Anti-Atlas. A trachyte sill of the Djbel Boho volcano, sampled in the Bou Azzer northern area, is analyzed by Levresse (2001). Recent investigation of this volcano by Alvaro et al. (2006) yields chemical and petrographic data. Products of the Djbel Boho volcano are intercalated, in the northwest of Bou Azzer, in the Tamjout dolostone (the C1 formation), where we sampled the basaltic lava flows. Because of its stratigraphic position, this activity took place between the V1 and V2 phases of the Agoundis-Ounein area, but 130 km farther to the ESE (Fig. 1). Two analyses of olivine-basalts have been performed (Table 3). Chemical composition is very different from that of the Agoundis-Ounein volcanics. The Djbel Boho basalt is alkali sodic, TiO2-rich, and has the typical incompatible element
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
49
TABLE 3. CHEMICAL ANALYSES OF V1–V5 LATE EDIACARAN TO CAMBRIAN VOLCANIC ROCKS OF THE AGOUNDIS-OUNEIN AND TOUBKAL AREAS AND OF THE CAMBRIAN DJBEL BOHO VOLCANO Volcanic stage Location Sample number
V1 Agoundis Aa-300 Flow B
Aa-216 Flow B-Pl
Toubkal Aa-408 Dike B-Pl
Aa-409 Dike B-Pl
Aa-405 Dike B-Pl
Aa-412 Dike Dol
Agoundis Ag-4 Flow B-Pl
Aa-3 Flow B-sp
Aa-217 Dike B-sp
Ag-2 Flow B-Pl
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
47.21 2.76 16.34 13.46 0.12 4.89 2.99 4.50 3.39 0.67 3.64 99.97
47.80 1.66 19.63 9.49 0.08 4.62 2.11 3.75 4.96 0.39 5.51 100.00
48.34 2.16 16.05 12.69 0.24 7.04 2.63 4.91 1.85 0.54 3.78 100.23
48.54 1.11 16.39 10.39 0.16 8.91 0.61 2.13 4.78 0.18 5.55 98.75
48.74 2.11 15.48 11.60 0.25 7.36 3.21 4.55 1.27 0.48 5.69 100.74
48.95 2.49 15.19 13.43 0.22 4.83 6.37 3.78 1.91 0.54 1.64 99.35
49.07 1.65 18.24 9.26 0.12 6.38 2.26 4.99 2.17 0.34 4.66 99.14
49.18 1.22 16.63 8.87 0.13 4.32 5.74 5.72 2.24 0.30 5.54 99.89
49.19 1.85 16.92 9.58 0.13 5.70 4.02 6.11 0.88 0.33 4.71 99.42
49.80 1.79 18.73 9.13 0.11 5.78 2.41 5.14 2.50 0.35 4.16 99.90
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
409 30.7 34.1 24.7 22.16 72.2 167 57.37 362 21.49 2.36 1390 8.50 1.64 4.97 3.35
295 80.8 35.5 45.0 22.04 101.5 475 32.07 208 12.06 5.82 1536 4.68 0.88 2.55 1.64
212 136.0 38.8 84.6 21.58 38.5 120 42.96 261 14.11 0.43 539 5.85 1.08 2.01 1.35
182 313.3 40.4 128.7 18.02 35.0 67 19.60 92 3.55 6.99 447 2.73 0.26 1.92 2.36
219 182.2 40.5 98.9 20.64 25.7 171 41.36 238 12.96 0.71 272 5.38 1.00 1.86 1.11
255 34.4 43.9 57.3 22.65 56.8 397 44.37 272 15.70 0.95 586 6.25 1.20 3.03 1.56
223 75.9 28.8 27.0 25.62 49.1 208 35.02 213 13.02 2.70 573 5.48 1.02 3.27 0.97
139 31.9 24.1 23.4 17.50 30.8 274 23.58 144 8.38 0.45 663 3.40 0.62 1.96 1.04
276 30.5 27.8 16.0 23.87 10.8 242 24.22 143 6.08 0.50 289 3.42 0.42 1.27 0.86
233 68.2 25.4 25.2 24.36 55.1 237 37.28 227 13.77 2.90 763 5.91 1.10 3.46 0.94
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
49.83 110.50 14.63 58.44 11.02 2.65 10.43 1.67 10.02 1.97 5.43 0.80 0.81 5.27
21.02 47.56 6.41 28.06 6.57 2.27 6.19 0.98 5.64 1.12 3.00 0.51 0.48 3.18
20.08 51.03 7.86 35.85 8.35 2.60 8.18 1.27 7.64 1.52 4.26 0.62 0.65 4.10
9.80 20.59 2.79 12.21 3.07 0.77 3.26 0.55 3.37 0.70 2.03 0.29 0.31 1.98
26.15 56.46 7.91 35.08 8.32 2.66 8.37 1.26 7.48 1.44 3.99 0.58 0.59 3.88
26.72 60.00 8.40 36.59 8.60 2.67 8.39 1.32 7.87 1.55 4.32 0.63 0.65 4.23
19.50 46.35 6.55 27.84 7.31 1.98 6.99 1.09 6.41 1.24 3.30 0.48 0.45 3.01
12.77 33.51 4.86 21.27 5.71 1.66 5.02 0.77 4.37 0.88 2.25 0.36 0.36 2.21
14.10 29.60 3.88 20.13 5.89 2.09 5.74 0.83 4.36 0.86 2.36 0.36 0.31 2.10
19.96 46.55 6.85 30.34 7.45 2.32 7.06 1.08 6.27 1.41 3.56 0.51 0.54 3.61
Rock type (wt%)
Continued
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Pouclet et al.
TABLE 3. CHEMICAL ANALYSES OF V1–V5 LATE EDIACARAN TO CAMBRIAN VOLCANIC ROCKS OF THE AGOUNDIS-OUNEIN AND TOUBKAL AREAS AND OF THE CAMBRIAN DJBEL BOHO VOLCANO (continued) Volcanic stage Location Sample number
V1 Agoundis Aa-215 Flow B-Pl
Aa-3b Flow B-Pl
Aa-310 Flow B-Pl
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
50.06 1.31 19.12 8.94 0.06 5.71 1.86 5.28 2.81 0.25 4.47 99.87
50.54 1.43 18.71 8.96 0.08 3.23 5.89 3.21 3.41 0.29 4.16 99.91
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
215 70.0 26.0 38.5 18.70 55.3 339 25.40 154 9.42 3.67 521 3.40 0.64 2.57 1.31
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
15.25 35.22 4.94 21.78 4.85 1.51 4.69 0.72 4.45 0.86 2.44 0.36 0.40 2.48
Rock type
V2
V3 Aa-201 Laccolith Mdior
V4 Ounein A-6 Flow B
Aa-213b Flow B-sp
Aa-20 Flow B-sp
Aa-219 Laccolith Mdior
Aa-221c Laccolith Mdior
Aa-45a Flow B
58.56 0.69 16.40 6.83 0.08 2.00 2.52 4.94 3.55 0.20 4.03 99.80
61.27 0.93 15.46 6.26 0.05 3.50 1.59 5.71 1.17 0.27 3.65 99.86
51.05 1.39 17.83 9.53 0.16 5.90 2.23 5.70 1.87 0.39 3.84 99.89
64.24 0.40 14.95 2.24 0.00 1.48 2.93 1.10 8.56 0.15 3.95 100.00
64.74 0.55 15.74 4.27 0.04 1.28 2.98 4.42 2.84 0.13 2.92 99.91
65.10 0.55 15.66 4.23 0.04 1.26 2.57 4.34 3.04 0.21 2.93 99.93
46.31 1.71 16.77 10.74 0.11 7.09 8.09 2.08 2.96 0.29 3.48 99.63
46.63 2.57 16.67 12.98 0.18 5.09 7.66 3.76 0.07 0.38 3.88 99.87
211 35.1 19.5 19.4 20.12 67.8 374 26.78 169 9.47 4.61 679 3.92 0.72 2.40 1.31
120 13.3 16.2 16.1 18.92 44.6 80 18.17 151 6.60 3.52 445 4.38 0.60 5.20 2.56
67 17.0 18.8 15.4 22.09 15.4 144 37.55 315 15.54 0.87 484 7.73 1.23 1.72 4.98
180 24.8 24.9 16.1 23.20 20.6 157 28.15 189 5.12 0.85 507 4.07 0.40 1.76 1.47
31 12.8 5.9 9.6 17.42 132.8 117 28.48 357 13.19 1.41 2226 7.98 1.19 9.11 4.28
48 43.1 8.1 24.5 22.38 64.8 314 26.05 318 12.44 1.33 808 7.40 1.15 9.13 3.99
42 14.6 7.3 10.7 19.89 61.5 370 24.58 309 12.19 1.04 887 7.11 1.15 8.64 3.67
266 141.0 0.3 91.0 23.10 42.8 538 22.50 140 8.98 1.99 685 3.83 0.76 1.08 0.73
273 14.4 38.3 26.2 24.35 11.7 453 46.60 271 8.04 0.27 124 5.92 0.67 1.95 1.02
19.44 42.85 5.54 25.68 5.54 1.70 5.29 0.79 4.52 0.92 2.46 0.40 0.40 2.64
17.58 36.70 4.62 18.15 3.73 0.93 3.24 0.51 3.15 0.65 1.89 0.31 0.34 2.12
54.54 123.87 14.55 54.94 10.81 2.00 7.92 1.11 6.17 1.23 3.66 0.57 0.61 3.82
14.31 39.15 5.89 26.72 6.48 1.92 5.52 0.85 4.79 0.97 2.53 0.40 0.46 2.64
38.18 75.64 8.51 30.77 5.76 1.36 5.12 0.79 4.19 0.80 2.39 0.37 0.41 2.45
40.62 79.37 9.01 32.57 6.22 1.53 4.68 0.75 4.23 0.76 2.32 0.37 0.39 2.55
36.75 74.04 8.00 28.67 5.45 1.31 4.42 0.70 4.02 0.76 2.28 0.39 0.38 2.55
12.32 32.74 4.19 18.75 4.22 1.51 4.62 0.73 4.25 0.92 2.08 0.29 0.31 1.72
17.84 44.41 6.06 28.57 7.50 2.36 7.16 1.22 7.65 1.54 4.10 0.63 0.65 4.25
(wt%)
Continued
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
51
TABLE 3. CHEMICAL ANALYSES OF V1–V5 LATE EDIACARAN TO CAMBRIAN VOLCANIC ROCKS OF THE AGOUNDIS-OUNEIN AND TOUBKAL AREAS AND OF THE CAMBRIAN DJBEL BOHO VOLCANO (continued) Volcanic stage Location Sample number
V4 Ounein Aa-225 Flow B
Aa-44 Flow B
Aa-202 Plug B-sp
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K 2O P2O5 LOI Total
47.05 1.73 16.44 10.56 0.08 6.92 6.28 3.58 1.98 0.31 5.01 99.94
47.85 2.26 16.39 13.07 0.18 5.13 6.92 4.01 0.00 0.37 3.64 99.82
(ppm) V Cr Co Ni Ga Rb Sr Y Zr Nb Cs Ba Hf Ta Th U
265 174.7 33.1 91.2 18.84 24.6 707 23.27 150 9.33 3.34 625 3.57 0.76 1.28 0.71
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Lu Yb
13.68 34.24 4.60 20.14 5.15 1.72 4.99 0.71 4.36 0.83 2.19 0.32 0.33 2.04
Rock type
Aa-7 Plug B-sp
V5 Ounein Aa-209 Laccolith Mgab
Bou Azzer Djbel Boho Dba-1 Dba-2 Flow Flow Ol-B Ol-B
A-1 Laccolith Mgab
Aa-205 Laccolith Mdior
50.49 2.32 16.53 8.94 0.06 5.84 2.84 6.87 0.11 0.48 4.04 98.52
53.48 1.92 15.83 9.07 0.14 3.81 3.13 6.07 0.92 0.42 5.09 99.88
52.95 1.93 15.49 8.70 0.21 3.99 3.36 5.73 0.59 0.40 6.50 99.85
53.73 2.08 15.80 9.14 0.27 3.29 3.66 6.66 1.10 0.39 3.43 99.55
61.83 0.84 17.10 5.49 0.05 1.61 3.28 4.29 2.87 0.24 2.33 99.93
48.13 3.09 15.26 12.78 0.06 3.75 4.03 4.45 2.48 1.07 4.74 99.84
48.74 3.10 15.44 12.73 0.06 4.04 3.87 5.01 2.00 1.07 4.35 100.41
248 18.3 37.0 32.8 23.43 11.5 392 42.72 246 7.50 0.63 130 5.60 0.63 1.95 1.01
321 73.4 27.4 29.9 21.06 11.1 125 30.86 160 7.78 0.00 125 3.73 0.66 1.67 1.39
128 7.0 15.9 6.2 21.98 10.3 259 35.31 240 9.75 0.43 661 5.66 0.87 2.40 1.71
133 5.0 12.6 5.0 22.56 11.4 157 33.65 229 9.77 0.41 310 5.39 0.85 2.58 1.77
191 7.0 0.3 7.1 22.30 14.0 474 34.00 215 9.23 1.95 739 5.73 0.82 2.33 1.80
66 51.9 11.0 27.9 20.82 59.3 596 23.56 260 12.71 0.56 868 6.11 1.12 8.31 2.97
158 7.5 23.9 10.3 25.00 53.4 96 34.85 317 57.39 1.09 320 6.85 3.62 4.92 1.55
144 10.6 24.2 13.4 23.02 43.5 107 39.59 321 57.88 0.99 350 6.91 3.53 4.90 1.86
17.29 41.78 5.83 26.91 6.78 2.28 7.03 1.21 6.95 1.52 3.86 0.61 0.59 4.11
12.06 30.97 4.85 23.52 6.24 2.13 6.64 1.09 6.44 1.05 2.66 0.40 0.31 2.31
20.14 49.03 7.03 32.36 8.40 2.77 7.77 1.16 6.86 1.35 3.50 0.54 0.55 3.44
18.69 46.72 6.00 27.25 6.62 2.04 6.47 1.03 6.02 1.16 3.40 0.52 0.46 3.35
18.86 47.96 6.51 27.41 6.69 1.82 6.67 1.10 6.38 1.44 3.52 0.49 0.53 3.06
34.40 71.26 8.15 30.20 6.06 1.50 4.81 0.74 4.11 0.81 2.35 0.37 0.33 2.27
37.73 81.17 10.08 41.61 8.74 2.85 7.90 1.17 6.64 1.26 3.37 0.48 0.47 3.06
39.57 86.41 10.87 44.73 9.92 3.42 9.04 1.39 7.80 1.43 3.79 0.52 0.49 3.29
(wt%)
Note: Analytical laboratory of the Research Center for Petrography and Geochemistry of Nancy, France. Same method as for Table 2. B—basalt; B-Pl—plagioclase phyric basalt; Dol—dolerite; B-sp—spilitic basalt; Mdior—microdiorite; Mgab—microgabbro; Ol-B—olivine basalt. LOI—loss on ignition.
Rb Ba Th Nb Ta
Sr
Nd Zr
Hf Sm Eu Gd Ti
La Ce Pr
Sr
Nd Zr
Hf Sm Eu Gd Ti
continental tholeiites
V1-V2 basalts
La Ce Pr
Ediacaran andesites
Rb Ba Th Nb Ta
V1 V2
Dy Y
Dy Y
Yb
C
Yb
A
1
10
100
1000
1
10
100
1000
Rb Ba Th Nb Ta
OIB
Rb Ba Th Nb Ta
Sr
Nd Zr
La Ce Pr
Sr
Nd Zr
V1-V2 basalts
La Ce Pr
V1-V2 basalts
Dy Y
D
Yb
Hf Sm Eu Gd Ti
Dy Y
Yb
Djbel Boho volcano Ol-Basaltes
Trachyte
Hf Sm Eu Gd Ti
V3 V4 V5
B
Figure 11. Incompatible element diagrams normalized to Primitive Mantle for the Late Ediacaran to Cambrian lavas. (A) Patterns of the V1 and V2 lavas. (B) Patterns of the V3, V4, and V5 lavas compared with the V1–V2 compositional range. (C) Comparison of the V1–V2 basalts with the Ediacaran Atlasic Volcanic Chain andesites and with average composition of continental tholeiites after Holm (1985). (D) Patterns of the olivine-basalts and trachyte of the Tommotian Djbel Boho volcano; comparison with the V1–V2 compositional range; OIB—average pattern of intraplate oceanic island basalt. OIB composition and normalization values after Sun and McDonough (1989).
1
10
100
1000
1
10
100
1000
Rock / Primitive Mantle Rock / Primitive Mantle
Rock / Primitive Mantle
Rock / Primitive Mantle
52 Pouclet et al.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin
53
10.00 Shoshonite Upper crust
Ediacaran andesites
Ocean Island Basalt DB-basalt
Calc-alkaline basalt
1.00
SZ
0.10 0.01
A rra
CC
tle
AFC FC
M an
V5 V4 V3 V2 V1
Figure 12. Ta/Yb vs. Th/Yb diagram after Pearce (1982) for the Late Ediacaran and Cambrian lavas. Comparison with the Ediacaran Atlasic Volcanic Chain andesites. CT—Continental Tholeiite after Holm (1985); DB—Djbel Boho. SZ to FC are defined in Figure 5.
CT
y
Th/Yb
DB-trachyte
Primitive Mantle
0.10
1.00
10.00
Ta/Yb
pattern of oceanic island basalt (OIB; Figs 11D and 12). The trachyte profile confirms the alkaline magmatic signature. Its incompatible element enrichment together with the Ba, Sr, Eu, and Ti negative anomalies indicate a derivation of the trachyte from the basalt by fractional crystallization. The OIB compositions are related to intraplate asthenospheric magmas. A continental rifting environment can be suggested, as is proposed by Alvaro et al. (2006). Significance of the Mantle Lherzolite Slices of the Ounein Area The chemical composition of the Ounein lherzolite is provided in Table 4. One may note a low TiO2 content and moderate amounts of alumina and lime. In the Atlas region, two other occurrences of ultrabasic rocks are known: in the Sidi Flah Cryogenian inlier to the west of the Saghro window, and in the Bou Azzer and Siroua inliers of the central Anti-Atlas. The Sidi Flah rocks outcrop as eleven small bodies intercalated into the Cryogenian sediments along submeridian- to ENE-trending faults (Fekkak et al., 2003). The rocks consist of totally serpentinized peridotites with relictual orthocumulate and adcumulate textures of lherzolites, dunites, and wherlites. Owing to the preserved chrome-spinel composition and to the sedimentary and tectonic context, the ultrabasic rocks are attributed to an intracontinental basin (Fekkak et al., 2003). The Bou Azzer ultramafic rocks belong to ophiolitic complexes related to suboceanic mantle or to island arc substratum accreted along the Pan-African suture, during the main Pan-African orogenic stage (Leblanc and Lancelot, 1980; Saquaque et al., 1989; Admou, 2000). The rocks consist
of serpentinized harzburgites, dunites, and wherlites associated with layered pyroxenites and gabbros. Selected revised and new analyses of these ultramafic rocks are given in Table 4. In comparison, the rare earth element contents of the Ounein lherzolite show an almost flat pattern with weak La and Ce depletion (Fig. 13), which precludes a normal subcontinental mantle origin. In marked contrast, the Sidi Flah lherzolite is enriched and displays a fractionated pattern typical of subcontinental mantle peridotites. The associated dunites are strongly depleted in the incompatible elements, as is also the case for the Bou Azzer ophiolite dunite and harzburgite. Only the more enriched wherlite and pyroxenite can be plotted. They display oceanic mantle-type depleted profiles. Compared to the Galicia Margin peridotites, which represent the transitional upper mantle between continental and oceanic ones, the Ounein lherzolite is less depleted. It can be interpreted as originating from a transitional mantle that has moderately participated in the production of small amounts of basaltic magma. DISCUSSION Summary of the New Results We examine the geotectonic context at the Late Neoproterozoic–Cambrian boundary for the Early Cambrian marine transgression. In the Anti-Atlas and southwest of the old block of High-Atlas, major to minor unconformities are apparent (Fig. 2). The angular unconformities cannot be explained by simple block tilting, because distinct fold axes trending WNW-ESE are
54
Pouclet et al. TABLE 4. CHEMICAL ANALYSES OF THE OUNEIN LHERZOLITE AND OF SELECTED ULTRAMAFIC ROCKS FROM SIDI FLAH AND BOU AZZER Location Sample number Rock type
Ounein Aa-2 Lherzolite
Sidi Flah Sfl-7 Lherzolite
Sfl-10 Dunite
Bsk-2 Dunite
Bou Azzer Baz-3 Dunite
Baz-5 Wherlite
Baz-8 Plagioclase pyroxenite
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
40.52 0.06 2.15 8.35 0.11 39.54 1.84 0.14 0.00 0.10 7.50 100.31
39.02 0.00 0.62 7.62 0.06 38.18 0.16 0.00 0.00 0.10 13.67 99.43
38.78 0.00 0.59 8.62 0.04 37.30 0.22 0.00 0.00 0.10 12.89 98.54
42.23 0.00 0.50 7.40 0.10 34.62 2.74 0.00 0.00 0.12 10.90 98.61
37.24 0.00 1.60 10.29 0.27 37.33 0.14 0.00 0.00 0.08 10.30 97.25
48.52 0.07 2.47 9.62 0.22 24.84 8.86 0.05 0.00 0.08 5.15 99.88
49.56 0.13 10.58 7.16 0.17 14.28 13.71 1.64 0.11 0.06 2.45 99.85
(ppm) V Cr Co Ni Ga Sr Y Zr Ba
47 2689.2 109.0 2099.1 2.01 7 1.56 3 +6) basalts associated in the field, several occurrences show acidic magmas with intermediate Nd isotope signatures. This is the case of the Krˇ ivoklátRokycany rhyolites (Table 4), which have εNd values close to
223
zero for three of four analyzed samples. As another example, two samples from the Vesser bimodal suite, containing 58 and 71 wt% silica, have εNd500 values of −0.5 and +0.2, respectively (Bankwitz et al., 1994). These Nd isotope features indicate that the average source of these rhyolites was neither enriched nor depleted in terms of time-integrated Sm/Nd ratio. This almost chondritic Nd isotope signature is open to several geological interpretations. First, it might reflect closed-system differentiation from a mafic magma extracted from a source with broadly chondritic Sm-Nd characteristics, as documented for many continental flood basalt suites (irrespective of the specific explanation given to this observation). Although such basalts have not been documented in the Vesser area, this hypothesis cannot be totally dismissed, bearing in mind the poor preservation potential of these subaerial rock types. Note in this respect that two samples of ca. 490-Ma orthogneisses from the high-grade Góry Sowie analyzed by Kröner and Hegner (1998) have initial εNd values of +0.2, similar to some of the mafic plutonic rocks (metagabbronorites) of the same domain, which are geochemically similar to continental tholeiites (Kryza and Pin, 2002). Second, the εNd values close to zero might reflect igneous mixing, in adequate proportions, of typical continental material (with negative εNd values) and a mafic magma extracted from a time-integrated depleted mantle source (i.e., with positive εNd), such as those occurring in the Vesser suite (εNd500 up to +7.6; Bankwitz et al., 1994) and in the Krˇ ivoklát-Rokycany complex (εNd500 approximately +5). This mixing might have occurred through crustal contamination during ascent through, and/or storage in, the crust via bulk assimilation or AFC processes. Such a scenario may be supported, in the Krˇ ivoklát-Rokycany complex, by the occurrence of an andesite with εNd500 +2.6 and a rhyolitic sample with εNd500 +1.8, because these two samples appear to bridge the gap between mafic and felsic rocks and possibly document an AFC process. Attempting to model quantitatively these inferred mixing processes is beyond the scope of this review and would remain a rather academic exercise, insofar as the potential crustal end-members are very poorly constrained in terms of Nd concentrations and isotope signatures. Alternatively, the Sm-Nd data might be satisfied readily if a source with suitable isotope characteristics occurred at depth in the local crust and was able to partially melt during the ca. 500-Ma episode. This scenario might have involved anatexis of interlayered amphibolite and peraluminous metasediments, as experimentally studied under lower crustal conditions by Skjerlie and Patiño Douce (1995). In another interpretation, the mixed source might have been generated during the erosion-sedimentation cycle. This hypothesis cannot be a priori dismissed for the Krˇ ivoklát-Rokycany complex rhyolites, based on scarce Nd isotope data on Late Proterozoic graywackes of the Barrandian area, which point to εNd500 values near −1 (Pin and Waldhausrová, unpublished data). In such a scenario, the mafic end-member (basalts and andesites) of the Krˇ ivoklát-Rokycany complex might represent magmas extracted from a hydrous upper mantle domain inherited from Late Proterozoic subduction processes,
224
Pin et al.
whereas the rhyolites would mainly reflect crustal melting of a hybrid source (graywackes) generated, several tens of Ma earlier, by sedimentary mixing in a Late Proterozoic basin that trapped both volcanogenic detritus from juvenile sources and old recycled clastic components. In conclusion, extrusive silicic magmas occurring either alone or, more commonly, as part of broadly bimodal mafic-felsic associations, include: 1. Rhyolites of pure or predominantly crustal derivation, representing, at least in part, the extrusive counterpart of the much more voluminous orthogneisses; 2. Rhyolites or trachytes, and even plagiogranites, of exclusively mantle origin, corresponding to felsic derivatives of abundant, associated enriched or depleted basaltic magmas; and 3. Rhyolites of inferred hybrid origin, generated either as a result of a high degree of crustal contamination of mantle-derived magmas ascending through the crust, or by partial melting of mixed sources (e.g., interlayered sediments and mafic rocks, or graywackes consisting of a sedimentary mixture of epiclastic components of ancient crustal origin and juvenile components fed by the erosion of mantle-derived, broadly coeval igneous rocks). Causes of Crustal Melting The generation of copious volumes of magma from crustal protoliths is well documented by the geochemical—and, particularly, Nd isotope—features of the ca. 500-Ma orthogneisses and some of the broadly coeval felsic metavolcanics. This crustal derivation prompts the question of the heat source for partial melting. Although the existence of certain low-temperature granitoids can be inferred (e.g., Miller et al., 2003), it is generally accepted that S-type felsic magmas erupted or emplaced at shallow crustal level were H2O undersaturated and generated in the lower crust through HT melting. Melting temperatures >800 °C are indicated by geothermometry studies (see references in Clemens, 2003), in agreement with the 850–900 °C range of dehydration melting experiments. Such elevated temperatures and the abnormally high heat flow they imply cannot be reached by crustal thickening, but require advective heat input from the mantle (e.g., Thompson, 2000). Therefore, it is concluded that mantle was the source of the excess heat responsible for widespread crustal melting during the 500-Ma event. More specifically, it is inferred that hot asthenosphere uprising during progressive stretching of the overlying lithosphere provided both an increased basal heat flow and basaltic partial melts, which underplated and intruded the continental crust, thereby causing copious partial melting of fertile lithologies. This scenario is supported by the occurrence of coeval basaltic magmatism in many of the occurrences discussed in this work and the recognition that periodic, multiple intrusions of basaltic magmas over a time span of a few hundred thousand years provides a very efficient way to promote partial melting in the lower crust (e.g., Petford and Gallagher, 2001).
Inferred Tectonic Setting Continental lithosphere extension provides a suitable mechanism to trigger granulite facies and partial melting in the lower crust at a regional scale, particularly when asymmetric extension and crustal-scale detachments are involved (Sandiford and Powell, 1986). Indeed, two favorable conditions promoting partial melting of the lower crust are combined in extensional settings, specifically, decompression through dehydration melting reactions and increased heat supply through the crust-mantle boundary caused by lithospheric mantle thinning. The large-scale spatial and temporal association of crustal melts with mantlederived magmas typical of rifting contexts further demonstrates that intrusion of mafic magmas occurred in many places and allowed for convective heat transfer into the surrounding crust. The combination of a “hot” tectonic setting with the presence of lithologies characterized by high melt productivity (Neoproterozoic graywackes) favored the generation of mobile magmas that emplaced as high-level granites (now orthogneisses) or erupted as lavas. A large-scale rifting context leading to continental breakup is independently indicated by the sedimentary record of the Late Cambrian–Early Ordovician wherever the Variscan tectonometamorphic overprinting was not too strong (e.g., Falk et al., 1995; Kemnitz et al., 2002). The occurrence of such examples as the Krˇ ivoklát-Rokycany complex, with undeformed mafic rocks similar to continental tholeiites and felsic rocks showing affinities with A-type granitoids, demonstrates that the rifting event failed in some cases, merely producing a shallow marine basin. This failure allowed the ca. 500-Ma volcanics to escape tectonometamorphic overprinting during Variscan events. Besides this example of aborted break-up, the temporal evolution of igneous rocks toward N-MORB (Vesser) suggests progressive rifting, evolving toward oceanic spreading. Independently, the relics of HP metamorphism documented in several other occurrences indicate that some rifted segments were brought down to mantle depths during their subsequent evolution. This HP imprint is interpreted as a record of subduction of former passive margins, following consumption of negatively buoyant, attached oceanic lithosphere. Along with scarce ophiolitic relics (e.g., Pin, 1990), this subduction process demonstrates that the 500-Ma rifting episode reached an ocean-spreading stage. The term anorogenic might convey the false connotation that rifting occurred in the middle of a large stable craton. Instead, the rifted domain was situated within a broad band of relatively juvenile continental lithosphere rimming the northern edge of the north African (Gondwana) cratonic domain. This lithosphere was largely newly formed before and during the Cadomian (panAfrican) orogeny, probably as the result of long-lived igneous and sedimentary accretion in a Pacific-type active margin setting. We do not concur, however, with authors supporting a still-active subduction as the driving mechanism for the 500-Ma rifting and break-up event. It seems more likely that active subduction ceased significantly earlier, either before Cadomian tectonics through switching to a transform regime (cf. Nance et al., 1991), possibly
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif following a ridge-trench encounter, or at the time of Cadomian collisional processes (as the final result of oblique convergence), some 50 Ma before the major break-up episode. Igneous rocks emplaced during the intervening time span, which might be interpreted conventionally as late Cadomian magmas, consist of ca. 540-Ma I-type granodiorite plutons of crustal derivation (as shown by Nd isotopes) in Lusatia and eastern Erzgebirge (e.g., Korytowski et al., 1993; Kröner et al., 1994; Linnemann et al., 2000; Tichomirowa et al., 2001; Dörr et al., 2002) on one hand, and ca. 510–520-Ma I-type granitoids and gabbros in the Teplá-Barrandian domain (e.g., Dörr et al., 1998, 2002) on the other. Whatever the geodynamic cause(s) for these older igneous events, the change from I-type to S-type sources for granitoid magmas during the Cambrian period might reflect sequential partial melting events of a vertically zoned late- to post-Cadomian crust, mostly composed of pre-Cadomian meta-igneous materials at deeper levels, overlain by a thick sedimentary pile including both Late Proterozoic graywackes and Cambrian clastics. In a very tentative interpretation, the ca. 540-Ma magmas and coeval LP-HT metamorphism (Zulauf et al., 1999) might have been generated as the result of slab break-off (cf. Atherton and Ghani, 2002) following the Cadomian collision. Based on the occurrence of very thick deposits accumulated in shallow, rapidly subsiding depressions under continental, then marine, conditions (Chlupácˇ et al., 1998) and structural data pointing to dextral transtension (Zulauf et al., 1997), it is likely that a rift-related regime prevailed throughout the Cambrian system in the Barrandian domain. For this reason, and in the lack of detailed geochemical and isotope data, it is considered that the ca. 510–520-Ma plutons of the Teplá-Barrandian record an early igneous pulse within the broader context of a protracted period of oblique extension, as already proposed by Zulauf et al. (1997). Active rifting triggered by the impingement of a mantle plume on the lithosphere is a popular model, which was suggested to explain the Cambro-Ordovician rifting in the Bohemian Massif (e.g., Floyd et al., 2000). However, it is not possible to clearly document a deep-mantle plume origin for rift-related basalts on geochemical grounds, and we are reluctant to invoke a “plume” as the best explanation for mantle melting and continental break-up. Indeed, purely plate-tectonic processes leading to so-called “passive” rifting and continental break-up are probably sufficient causes to generate significant melts from the asthenosphere (e.g., Smith and Lewis, 1999; Anderson, 2000; Favela and Anderson, 2000). No large excess temperature (i.e., no plume) is required if relatively fertile mantle underlay the stretched, rifted lithosphere. We suggest that this was indeed the case beneath the newly accreted Cadomian crustal domain. BROADER IMPLICATIONS Beyond their geodynamic bearing on the evolution of the continental lithosphere in the Bohemian Massif, our results have broader implications on two general issues, namely, the debated status of “A-type granitoids” and the characterization of the
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inferred tectonic setting of granitoids through the use of geochemical discrimination diagrams. This review demonstrates that highly contrasting silicic magmas may be generated, broadly concomitantly, in a single province of medium size (~500 km across) under a similar extensional tectonic regime. These magmas, ranging from crustally derived, metaluminous to peraluminous granitoids, to mantlederived, sometimes peralkaline silicic magmas, through rocktypes with intermediate characteristics, can be safely considered as anorogenic granitoids, although only a minority displays “Atype” geochemical characteristics as usually defined (Collins et al., 1982; Whalen et al., 1987). It has been known for decades that subalkaline or even peraluminous silicic igneous rocks occur in certain rifting environments and form a distinct group of anorogenic granitoids (e.g., Hanson and Al-Shaieb, 1980; Anderson and Thomas, 1985; Finger et al., 2003), besides the typical “Atype” granitoids or “within-plate granites” (Pearce et al., 1984). This observation highlights the fact that A-type granitoids, defined by both geochemical characteristics and inferred tectonic setting, are only an end-member among a much larger class of rocks, ranging from peraluminous to peralkaline, as shown in this study. Typical A-type granitoids are commonly interpreted as partial melts of lower crustal rocks (felsic granulites) that suffered earlier melt extraction (e.g., Collins et al., 1982) or merely H2O loss during a metamorphic event (Skjerlie and Johnston, 1992). Alternatively, they might represent partial melts of crustal igneous rocks of tonalitic to granodioritic composition (Creaser et al., 1991). In other cases (mafic-felsic bimodal associations), they correspond to differentiates of basaltic magmas (e.g., Eby, 1992) or to partial melts from underplated mafic bodies (e.g., Poitrasson et al., 1995). In any case, unusually high temperatures are required to trigger partial melting of relatively refractory source rocks, implying the involvement of thermal ± mass input from underlying mantle, as commonly occurs during lithospheric extension. This involvement accounts for the systematic association of A-type magmas with rift-related settings. However, if the rifted continental crust is relatively immature and contains fertile rock-types at depth, as inferred for the northern Bohemian Massif at 500 Ma, partial melting of such lithologies is inescapable, thereby producing relatively large volumes of peraluminous to metaluminous granitoids. This first style of anorogenic granitoids would be generated at an early stage from the most fertile lithologies present in the melting domain. At a later stage, more typical A-type magmas could be produced from more refractory lower crustal lithologies, provided that sufficient input of heat allows these rocks to melt. A-type granitoids are anticipated to occur alone in cratonic segments characterized by an ancient, relatively refractory lower crust. Conversely, their association with metaor peraluminous granites would characterize less-differentiated crustal segments that contain fertile lithologies. As already emphasized (e.g., Creaser et al., 1991), the A-type granite terminology is questionable because of its inherent ambiguity (geochemistry versus tectonics), and classifications based on factual mineralogical-geochemical criteria, and/or on inferred
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sources (e.g., the I-S scheme) seem preferable. In this framework, the ca. 500-Ma felsic magmas consist mainly of S-type, with minor I-type and/or even M-type. Diverse petrogenetic processes were involved, including differentiation from mantle magmas and partial melting of various crustal source materials. This multiplicity of processes gave rise to the observed geochemical and isotopic diversity, but whatever their chemical fingerprint, all these rocks were generated in an extensional tectonic regime in the broader geodynamic context of continental break-up. For this reason, they can all be termed anorogenic granitoids, despite their great diversity of source materials and petrogenetic mechanisms. Following a common practice in the study of ancient basaltic rocks, geochemical discrimination diagrams are often used to infer ancient tectonic environments of granitic rocks. This approach is potentially very misleading, as can be shown by the plotting of the vast majority of the ca. 500-Ma felsic rocks within the “volcanic arc granite” and “collision granite” fields, in contradiction to the typical within-plate affinity of associated basaltic rocks and geological constraints, wherever these have not been erased by Variscan overprint. This failure of chemical discrimination diagrams is interpreted to reflect the fact that granite magmas mainly mirror their (mostly crustal) sources and do not have any simple geochemical relationship with the geodynamic setting prevailing at the time of their genesis (see discussion in Oberc-Dziedzic et al., 2005a). Indeed, chemically similar granitic magmas could be produced by broadly similar degrees of partial melting of similar source materials, irrespective of the local geodynamic setting, provided that melting can occur. Rather, the “volcanic arc” signatures commonly found in the ca. 500-Ma orthogneisses are interpreted to reflect the ultimate provenance of their inferred sedimentary source, specifically, Late Proterozoic graywackes, which contained a significant contribution from subduction-related igneous sources (e.g., Jakeš et al., 1979; Drost et al., 2004). If this interpretation is accepted, the geochemical discrimination diagrams for the 500-Ma granitoids are biased by inheritance, and they do not convey any useful information on the tectonic setting at the time of magma generation. Clearly, ancient geodynamic settings should be inferred from a combination of various types of geological evidence, among which granite geochemistry should be used with extreme care.
that evolved from mantle-derived basalts (of both enriched and depleted types), abundant peraluminous orthogneisses emplaced, at least in part, as shallow intrusions, demonstrate that copious amounts of S-type granitic magmas were generated during the same event. These hot, mobile magmas, showing some geochemical resemblance with “volcanic arc” and/or “syncollision” granitoids, were produced by partial melting in the lower crust. Based on geochemical features and U-Pb age patterns of inherited zircons, it is inferred that the major source materials involved were metasediments, broadly similar to outcropping Neoproterozoic graywackes. These protoliths contained variable proportions of ancient (2 Ga and older), mature, recycled components and geochemically less mature components with a recent (ca. 540 Ma), more juvenile input. The high-temperature dehydration melting process was triggered by the advection of mantle heat, as allowed by the context of continental lithosphere extension and documented by broadly coeval basaltic magmatism on the scale of the igneous province. The large volumes of felsic magmas produced are interpreted to mirror the abundance of very fertile lithologies, such as metagraywackes, in the melting domain. In this scenario, following a proposal of Anderson and Bender (1989) for an older example of anorogenic granite magmatism, the large melt productivity would basically reflect the relatively juvenile and still largely undifferentiated nature of the local crustal segment accreted during the Cadomian orogeny. ACKNOWLEDGMENTS We gratefully acknowledge the perceptive and constructive comments of the reviewers, Dr. Fritz Finger and Dr. Peter Floyd. The article is based, in large part, on results of research carried out under the long-term bilateral cooperation between Département de Géologie, Centre National de la Recherche Scientifique, Université Blaise Pascal, France, and the Institute of Geological Sciences, University of Wrocław, Poland. The Barrande Project between the Czech Republic and France is also acknowledged. Maciek Kryza helped to computerize the diagrams. This article is a contribution to the International Geological Correlation Program Project 497. APPENDIX: ANALYTICAL METHODS
CONCLUSION This review highlights the diversity of rock-types and inferred source materials involved in felsic magmatism during the ca. 500-Ma event. Based on converging lines of evidence, including the geochemistry of concomitant basalts, the tectonostratigraphic context, and the igneous rock association, an extensional regime is clearly documented, as already emphasized by several previous studies. The 500-Ma igneous event is therefore interpreted to be unrelated to any active subduction or to any prior collisional orogeny, as was suggested by some earlier studies, but instead to be basically anorogenic and reflecting continental break-up. Besides volumetrically subordinate volcanics
New Major and Trace Element Data for Krˇivoklát-Rokycany Volcanic Rocks The major and trace element data were obtained at the Centre de Recherche Pétrographique et Géochimique, Nancy, France, by inductively coupled plasma (ICP) atomic emission spectrometry and ICP mass spectrometry, respectively, following methods described by Carignan et al. (2001). New Sm-Nd data for the Orlica-S´niez˙nik Massif Orthogneisses Sm-Nd isotope analyses were made in Clermont-Ferrand following isotope dilution, separation chemistry, and thermal ionization mass spectrometry methods described by Pin and Santos Zalduegui (1997).
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif The precision of 143Nd/144Nd ratios is based on within-run statistics and quoted as the standard error on the mean at the 95% confidence level (2 standard errors). During the analyses, the average results and corresponding standard deviations (SD) obtained on Nd isotopic reference materials were m = 0.511966, SD = 0.000015 on eight measurements for the AMES R French standard, and m = 0.512114, SD = 0.000005 for six determinations of the JNdi-1 Japanese standard, equivalent to 143 Nd/144Nd = 0.511857 for the previously widely used La Jolla isotopic standard (Tanaka et al., 2000).
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Geological Society of America Special Paper 423 2007
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts (“spilites”) from the Central Barrandian domain (Bohemian Massif, Czech Republic) Christian Pin* Département de Géologie, Centre National de la Recherche Scientifique and Université Blaise Pascal, 5 rue Kessler, 63038 Clermont-Ferrand Cedex, France Jarmila Waldhausrová Nad Hercovkow 422, 18200 Prague, Czech Republic ABSTRACT On the basis of immobile trace elements and Nd isotope signatures, the Barrandian metabasalts may be ascribed to two major groups, extracted from contrasting mantle sources: 1. A depleted group, with strong light rare earth element depletion, elevated Zr/Nb ratios (>30), and highly radiogenic Nd isotopes (εNd600 from +7.8 to + 9.3). Multi-element patterns normalized to normal mid-ocean ridge basalt all show negative anomalies of Nb, and to a lesser degree, Zr and Ti. Eight samples may define a 605 ± 39-Ma whole-rock isochron with εNdi of +8.8 ± 0.2. 2. An enriched group, comprising both mildly enriched (Zr/Nb 12–18) and strongly enriched (Zr/Nb 4–7) samples, with εNd600 ranging from +8.2 to +3.8. The depleted group is interpreted to reflect generation from depleted mantle sources fluxed by subduction-related components, probably in an intraoceanic back-arc basin. In contrast, the younger enriched group is typical of the within-plate style of mantle enrichment and documents the extinction of the subduction-related component. The switch from suprasubduction zone to within-plate magmatism suggests that new mantle material flowed into the former arc and back-arc system sources. This flow might have occurred simply as a result of ocean-ward migration of the subduction zone. Alternatively, the subduction fluxing might have stopped as a result of impingement of a spreading ridge with the intraoceanic trench, leading to mutual annihilation, a switch to a transform plate boundary, and opening of a slab window that allowed the inflow of new mantle and the generation of late-stage, within-plate enriched basalts. In terms of modern analogues, the Neoproterozoic of the Barrandian and other Cadomian regions of western Europe resemble arc and back-arc systems from the western Pacific region, where large intraoceanic subduction systems fringe major continental masses with a complex mosaic of microplates and magmatic arcs, including intervening basins floored either by oceanic crust or attenuated continental crust. Keywords: Sm-Nd, geochemistry, Cadomian, metabasalts, paleogeography *E-mail:
[email protected]. Pin, C., and Waldhausrová, J., 2007, Sm-Nd isotope and trace element study of Late Proterozoic metabasalts (“spilites”) from the Central Barrandian domain (Bohemian Massif, Czech Republic), in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, p. 231–247, doi: 10.1130/2007.2423(10). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Pin and Waldhausrová
INTRODUCTION The geological record of early stages of accretion and evolution of continental lithosphere in western and central Europe was largely obscured, or even erased, by extremely strong tectonometamorphic overprinting, which occurred during the Variscan (Hercynian) cycle, from Cambro-Ordovician continental breakup to Carboniferous continent-continent collisional orogeny. For this reason, the few large, coherent domains that escaped pervasive reworking during Paleozoic times deserve special interest. Besides the Ossa-Morena zone of southwest Iberia and the Cadomian block of north Brittany and Normandy (France) in western Europe, the Teplá-Barrandian unit and the Bruno-Vistulicum block in central Europe provide the best examples to investigate well-preserved Late Proterozoic igneous and sedimentary formations and to gain insight into how and when the pre-Variscan lithosphere was built. In the central part of the Bohemian Massif, the TepláBarrandian unit (Fig. 1) is composed of a low- to very-low-grade unit separated by major Variscan shear zones from the highly metamorphosed rocks of the surrounding Saxo-Thuringian segment of northwest vergence in the north, and Moldanubian segment of southeast vergence in the south (e.g., Glasmacher et al., 2002, and references therein). The Teplá-Barrandian block, and particularly its southeastern part (Barrandian area) displays a Precambrian basement, deformed and weakly metamorphosed during the Cadomian orogeny, and unconformably overlain by unmetamorphosed siliciclastic and carbonate sedimentary rocks of Cambrian to Middle Devonian age. This Paleozoic cover was only mildly deformed as a broad synclinorium prior to the deposition of Upper Carboniferous and Lower Permian continental sequences. Owing to these favorable circumstances, the Barrandian formations offer an excellent opportunity to study a relatively undisturbed segment of Late Precambrian crust, which might be representative of the much broader domain involved in, and strongly reworked by, the Cadomian and Variscan tectonometamorphic events. In particular, the nature of the materials forming the Late Precambrian crust (i.e., ancient crystalline continental basement versus juvenile crust and/or sediments) is pivotal to any model aiming to interpret the growth and evolution of European lithosphere. The Neoproterozoic sedimentary pile of the Barrandian unit contains relatively abundant mafic metavolcanic rocks, often referred to as “spilites” (e.g., Fiala, 1977). Insofar as their original features and mantle source can be deciphered through lowgrade alteration processes, these rocks may provide interesting clues to the geotectonic setting that prevailed during the formation of the Barrandian basin prior to the Cadomian orogeny and complement insights gained from the study of sedimentary country-rocks (e.g., Jakeš et al., 1979; Drost et al., 2004). In this work, the Sm-Nd isotope system, together with a set of trace elements selected on the basis of their relatively immobile behavior during postmagmatic processes, were used to reassess the geochemical characteristics of these igneous rocks and place constraints on
their inferred mantle sources. The possible geodynamic implications of these results are explored in the broader scope of the European Cadomian orogenic domain. GEOLOGICAL BACKGROUND Detailed information and interpretations on the evolution of the Teplá-Barrandian unit, including correlations with other Cadomian segments, can be found in a recent synthesis offered by Krˇ íbek et al. (2000). Briefly, the Precambrian of the Central Barrandian domain (e.g., Chaloupsky et al., 1995) consists mostly of a very thick (probably >10 km), monotonous sequence of siliciclastic hemipelagites and turbidites. The basement is unknown, but seismic reflection data (9HR profile; Tomek and Dvorˇ áková, 1994; Tomek et al., 1997) revealed different structures for the upper and lower parts of the Barrandian crust. The upper part is characterized by reflectors with variable positions and dips, in line with the mild Cadomian folding described by Holubec (1995a,b). In contrast, below ~10 km, the 9HR profile imaged SSE-dipping multiple reflectors imbricated down to a depth of 20–25 km. This imbricate structure might be interpreted as an early Cadomian complex modified to some extent during the Variscan orogeny (Tomek et al., 1997), the reflectors possibly corresponding to individual thrust planes within a subductionaccretion complex similar to those found in modern arcs. Following Kettner’s pioneering work, most lithostratigraphic schemes proposed for this domain are based on the presence or absence of metabasalts, variably transformed into so-called “spilites.” According to recent proposals (e.g., Mašek, 2000), three major units can be distinguished, from bottom to top: (1) the Blovice Formation (~7000 m thick?), containing frequent occurrences of mafic effusive igneous rocks, interbedded with shales, siltstones, graywackes, and minor rock-types, including pyroclastics, black shales, black cherts, and rare carbonates (diamictites are also reported to occur); (2) the Davle Formation (~2000 m), composed of graywackes and shales associated with intermediate to felsic volcanics and volcaniclastics; and (3) the Šteˇchovice Group, composed almost exclusively of a thick pile (up to ~5000 m) of flyschlike clastic rocks, including conglomerates, devoid of siliceous rocks, but containing alkaline volcanics. The Blovice and Davle formations are gathered into the Kralupy-Zbraslav Group, broadly corresponding to the “Pre-spilitic” and “Spilitic” of former schemes, whereas the Šteˇchovice Group corresponds to the “Post-spilitic series.” It should be stressed, however, that the stratigraphic contact between the Davle and Blovice formations was never observed, and that it would be preferable to establish different lithostratigraphic schemes in particular subdomains of the Barrandian Neoproterozoic (Röhlich, 2000). According to new mapping (Holubec 1995b), the Barrandian Neoproterozoic can be divided into three lithostratigraphic groups (Fig. 1): the Lower Group consists of the Rabštejn and Úslava groups; the Middle, Zvíkovec Group; and the Upper, Rakovník Group. Although the structure of the deep part of the Barrandian Neoproterozoic can only be inferred from seismic reflection
Figure 1. Lithostratigraphic map of the Barrandian Neoproterozoic (Holubec, 1995b), with major lithostratigraphic groups folded during Cadomian orogeny. Tectonic structures: 1—foliation, mostly parallel to bedding; 2—anticline, syncline; 3—thrust, tectonic boundary; 4—fault; 5—granitoid. Full circles—localities studied; see Appendix for details. Names of volcanic belts according to Fiala (1977).
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts 233
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Pin and Waldhausrová
profile 9HR, the near-surface geology displays relatively mild regional deformation, with three generations of folds and thrusts (Holubec, 1995a). An unconformity has been proposed but not fully demonstrated between the stratigraphically higher Zvíkovec and Rakovník groups (Holubec, 1995a). The corresponding tectonic phase would have caused shortening of the lower part of the Barrandian package. Our study is focused on mafic volcanics of the Blovice Formation (sensu Mašek, 2000) in the central and northwestern part of the Barrandian region and does not deal with the Davle Formation, exposed only in the southeastern part of the Barrandian domain. The alkaline volcanics of the Rakovník Group (Holubec, 1995b), a lithostratigraphic equivalent of the Šteˇchovice Group occurring in the northwest flank of the Barrandian, were also studied. Although a deep marine environment is inferred for the deposition of siliciclastic sediments of the Blovice Formation, very shallow conditions (lagoons, evaporitic flats) were locally achieved near the edges of volcanic islands, as shown by oolithic and stromatolitic carbonates (sometimes silicified) and pseudomorphs after gypsum (e.g., Skocˇek and Pouba, 2000; Vavrdová, 2000) closely associated with mafic igneous rocks. Based on microfossils of cyanobacterial and algal origin, along with acritarchs, a Late Neoproterozoic (Vendian) age is ascribed to these sediments (e.g., Konzalová, 1980, 2000; Pacltová, 2000; Vavrdová, 2000). In the framework of recent geological time scale, these formations might therefore be considered as belonging to the Ediacaran system, spanning the ca. 630–542-Ma period (Gradstein et al., 2004). The age of deposition of the Šteˇ chovice Group is constrained to be younger than ca. 570 Ma by isotopic dilution thermal ionization mass spectrometry (ID-TIMS) U-Pb ages measured for two rhyolitic boulders from conglomerates (585 ± 7 and 568 ± 3 Ma; Dörr et al., 2002) and for the youngest detrital zircons from a graywacke (564 ± 16 Ma by sensitive high-resolution ion microprobe [SHRIMP] U-Pb dating; Drost et al., 2004). Important intercalations of predominantly mafic volcanic rocks, including frequent pillow lavas (Fiala, 1977) and often associated with metal-rich black shales (Pasava, 2000), occur within the Neoproterozoic sedimentary pile. These volcanics are exposed along several major belts, broadly parallel to the northeast–southwest-trending Variscan structures: the Svojšín belt and the Strˇ íbro-Plasy volcanic belt in the northwest, the main central volcanic belt in the central Barrandian, and the southern volcanic zone (Fiala, 1977) in the southeast (Fig. 1). The samples analyzed in this study were collected from the central and northwestern volcanic belts and from the alkaline volcanics occurring in the flyschlike sediments of the Rakovník Group. Based on previous studies (Fiala, 1977, 1978; Pelc and Waldhausrová, 1994; Waldhausrová, 1997a), three magmatic suites have been defined among the Neoproterozoic volcanics from the central and western Barrandian (Blovice Formation): (1) a primitive, tholeiitic suite, occurring in the lower part of the stratigraphic pile; (2) an alkaline suite, found in the upper part of the pile and representing the youngest Neoproterozoic volcanics;
and (3) a chemically transitional suite, stratigraphically lying between the tholeiitic and alkaline suites. The major part of the volcanics were affected by prehnite-pumpellyite metamorphicfacies overprinting prior to the deposition of Cambrian cover sediments (Bernardová and Chab, 1974), but metamorphic grade increases to greenschist- to amphibolite-facies conditions toward the southwest, northwest, and north. A petrographic description, including electron microprobe analyses of the main relictual igneous phases and metamorphic minerals of the Barrandian metavolcanics, is given by Waldhausrová (1997a). GEOCHEMICAL RESULTS Twenty-two samples of meta-igneous rocks from the northwest, central, and southern volcanic zones have been selected for major and trace element chemistry and Sm-Nd isotopes (see Appendix for sample locations). All but one sample (Si-2, corresponding to a trachytic lava) are of broadly basaltic composition, reflecting the overwhelming proportion of mafic rocks in this part of the Barrandian area. In addition two samples of metagraywackes (Viš-1 and Nml-1) and one of black shale (Kruš-2) have also been analyzed to get a crude estimate of the chemical and Nd isotope features of the metasedimentary country-rocks. The major and trace element data were obtained at the Centre de Recherche Pétrographique et Géochimique, Nancy, France, by inductively coupled plasma (ICP) atomic emission spectrometry and ICP mass spectrometry, respectively, following methods described by Carignan et al. (2001). The data are listed in Table 1. Sm-Nd isotope analyses were made in Clermont-Ferrand following isotope dilution, separation chemistry, and thermal ionization mass spectrometry methods described by Pin and Santos Zalduegui (1997). The results are given in Table 2, together with Nd isotope compositions corrected for in situ radiogenic decay of 147Sm, assuming geological ages of 570 and 600 Ma, respectively, and reported using the ε-notation. These results illustrate the range of initial isotope signature arising from a 30-Ma uncertainty in the true geological age. It can be seen that, in general, a ±30-Ma variation causes a shift in ε-value barely outside analytical uncertainty, highlighting that εNd values reported in Table 2 are not very sensitive to the uncertainty on the igneous emplacement age of the protoliths. The previous geochemical studies of the Barrandian volcanic rocks (see Waldhausrová, 1997a,b, and references therein) have used a set of data for major and trace elements, some of which are potentially mobile during seawater alteration and lowgrade metamorphism. Specifically, although the assessment of paleotectonic setting of the mafic metavolcanics was based on resistant elements, such as Ti, Mn, P (Mullen, 1983) and the lanthanides, the total alkali versus silica (TAS) classification scheme for volcanic rocks (Le Maitre et al., 1989) was used. However, Si, K, and Na, along with alkaline-earth elements (Ca, Sr, Ba), might have been redistributed to variable extent under the conditions to which the Barrandian rocks were subjected. Only elements extremely insoluble in aqueous fluids (Cox, 1995) and widely
44 400 76 272
0.95 3.71 0.76 4.68 2.07 0.81 3.14 0.58 3.93 0.85 2.52 0.38 2.58 0.40 22.8
Co Cr Ni V
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Y
3.08 9.36 1.60 8.76 3.11 1.19 4.09 0.72 4.77 0.98 2.86 0.43 2.89 0.44 27.4
38 349 148 253
4 301 240
46.2 17.00 10.00 0.12 7.44 10.7 2.57 0.27 1.30 0.13 4.11 99.87
1.78 5.63 1.01 5.61 1.99 0.95 2.68 0.47 3.10 0.65 1.84 0.27 1.73 0.27 18.1
32 388 72 191
2 153 224
47.1 22.3 4.62 0.09 4.92 10.9 3.65 0.26 0.82 0.08 5.11 99.83
1.30 3.97 0.72 4.33 1.77 0.75 2.60 0.48 3.32 0.73 2.10 0.32 2.12 0.33 19.9
43 432 205 197
3 134 26
46.5 17.4 8.71 0.13 9.41 11.0 2.55 0.16 0.73 0.07 3.09 99.83
1.99 6.72 1.23 7.16 2.73 1.08 3.87 0.69 4.65 0.97 2.83 0.42 2.84 0.43 27.8
45 368 74 287
1 126 26
47.9 14.7 10.5 0.15 7.64 10.6 3.43 nd 1.19 0.12 3.52 99.73
3.94 12.0 2.13 11.8 4.22 1.65 5.81 1.04 6.77 1.45 4.31 0.65 4.33 0.66 42.4
44 110 129 408
2 217 45
49.5 13.6 13.6 0.17 6.8 8.24 1.6 nd 1.78 0.16 4.24 99.77
1.38 4.39 0.89 5.44 2.27 0.91 3.40 0.65 4.29 0.93 2.66 0.40 2.62 0.39 24.9
44 461 112 282
9 129 123
49.9 17.8 6.59 0.09 6.58 8.85 4.27 0.47 1.00 0.11 4.15 99.86
2.42 7.10 1.21 6.68 2.53 0.95 3.53 0.65 4.29 0.92 2.65 0.40 2.60 0.40 26.0
42 226 46 274
3 90.2 47
48.9 14.7 9.52 0.17 7.47 11.1 3.22 0.13 1.05 0.11 3.44 99.81
3.57 11.6 2.10 11.8 4.29 1.43 5.84 1.10 7.34 1.55 4.56 0.68 4.63 0.70 42.7
50 110 50 428
1 58.1 26
51.7 14.4 11.00 0.21 5.64 6.85 4.34 0.10 1.95 0.19 3.38 99.81
2.43 7.78 1.43 7.93 2.89 1.11 3.85 0.70 4.57 0.96 2.79 0.42 2.80 0.44 26.4
45 381 120 415
4 169 123
50.00 16.6 11.3 0.12 6.40 7.34 2.93 0.15 1.33 0.13 3.68 99.87
2.72 7.89 1.29 6.93 2.46 0.98 3.3 0.58 3.87 0.81 2.36 0.36 2.36 0.35 23.3
45 344 78 236
6 123 124
47.2 15.5 9.31 0.14 8.72 9.77 2.88 0.25 0.92 0.11 5.08 99.84
6.86 16.8 2.42 11.5 3.28 1.38 3.89 0.67 4.43 0.91 2.60 0.38 2.52 0.4 24.6
43 350 122 228
7 158 157
46.6 16.4 8.31 0.12 8.58 9.47 2.95 0.96 1.16 0.15 5.15 99.84
6.53 16.1 2.31 11.0 3.22 1.26 3.89 0.67 4.40 0.90 2.58 0.38 2.56 0.39 23.9
42 375 118 227
1 70.5 31
4.19 11.2 1.60 8.08 2.46 0.86 3.15 0.56 3.71 0.78 2.28 0.34 2.23 0.35 22.6
46 415 179 218
4 321 100
45.2 45.3 17.1 13.9 9.18 8.43 0.08 0.12 11.9 6.89 5.43 13.1 1.55 3.38 0.10 0.17 1.16 0.93 0.15 0.13 7.97 8.06 99.83 100.44
7.28 17.5 2.51 12.0 3.49 1.35 4.18 0.70 4.52 0.93 2.65 0.39 2.56 0.39 26.9
39 429 138 250
6 279 158
46.2 16.7 9.67 0.13 8.79 10.7 2.28 0.36 1.35 0.21 3.46 99.88
36.3 76.2 9.28 37.9 7.8 2.46 7.05 1.08 6.02 1.12 3.04 0.43 2.78 0.41 31.5
41 102 51 247
25 498 440
49.1 18.6 9.51 0.11 3.69 3.78 5.09 1.33 3.16 0.55 4.93 99.91
36 211 65 245
14 246 353
46.6 18.4 8.60 0.14 5.96 11.3 3.11 0.81 1.10 0.16 3.72 99.87
14.6 14.4 29.4 29.0 3.55 3.49 14.6 14.4 3.38 3.31 1.19 1.18 3.46 3.44 0.56 0.55 3.51 3.46 0.71 0.69 2.000 1.98 0.29 0.29 1.96 1.96 0.3 0.3 19.7 19.13
37 220 66 256
14 255 358
46.3 18.4 8.52 0.14 5.94 11.4 3.04 0.81 1.09 0.17 4.03 99.86
47.8 94.9 10.9 42.6 8.67 2.73 8.49 1.36 8.42 1.69 4.78 0.72 4.81 0.75 47.1
29 4 7 297
36 368 586
49.5 13.7 13.1 0.19 3.70 7.58 3.17 1.52 2.82 0.52 4.08 99.91
49.3 97.5 11.2 43.8 8.93 2.80 8.69 1.40 8.63 1.73 4.95 0.74 4.96 0.78 49.2
31 13 9 330
52 289 603
48.9 14.1 13.5 0.21 4.22 8.51 2.96 1.43 2.83 0.51 2.71 99.9
22.5 45.1 5.84 22.7 4.63 1.08 4.02 0.65 4.000 0.82 2.45 0.38 2.67 0.41 23.8
6 28 15 55.3
56 112 631
70.8 14.8 3.73 K2O) of these rocks (see also Jakeš et al., 1979; Drost et al., 2004). In comparison, the black shale sample Kruš-2 displays broadly similar chondrite-normalized patterns for Th, Nb, and LREE, but it has deeper anomalies of Eu and particularly Ti, and contains half the HREE. Its distinctly negative εNd600 value (−3.7) reflects a larger contribution to the finer-grained detrital supply of epiclastic component(s) with time-integrated LREE enrichment, probably derived from ancient sialic crust. Model ages relative to modeldepleted mantle (DePaolo, 1988), interpreted to reflect the average crustal formation age of the mixture of detrital components present in the sample, are 1.1–1.2 Ga (graywackes) and 1.5 Ga (black shale), largely in excess of deposition age. This difference highlights the role of old, recycled components in the sedimentary input of the Neoproterozoic basin. In summary, on the basis of chondrite- and N-MORB-normalized diagrams, selected ratios of immobile incompatible trace
100
10
1
Th Nb La Ce Nd Zr Sm Eu Ti Gd Dy Y Er Yb Lu
Figure 6. Chondrite-normalized patterns of selected immobile incompatible trace elements for three samples of metasediments. The black shale Krus-2 differs from the two metagraywackes (Nml-1 and Vis-1) in showing higher contents of HREE. Normalization values from Sun and McDonough (1989).
elements, and Nd isotope signatures, the Barrandian metabasalts may be ascribed to two major groups, extracted from contrasting mantle sources: 1. A depleted group, characterized by variably strong LREE-depletion, elevated Zr/Nb ratios (>30), distinct negative Ti anomalies, and highly radiogenic Nd isotopes (εNd600 from +7.8 to + 9.3). Variable Th abundances in this group of samples are reflected by the absence or presence of negative Nb anomalies on chondrite-normalized plots. However, multi-element patterns normalized to NMORB all show well-defined negative anomalies of Nb, and to a lesser degree, Zr and Ti. This excess of LREE
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Pin and Waldhausrová and Th relative to HFSE suggests magma generation from depleted mantle sources fluxed by hydrous fluids and probably also indicates silicate melts derived from an oceanic slab and subducted sediments. Relatively low Ti/ V ratios are interpreted to reflect relatively oxidizing conditions compared to those prevailing at mid-ocean ridges (Shervais, 1982), possibly caused by higher water contents. The elevated εNd600 values preclude any significant contribution from old sialic crust, either as a contaminant during magma ascent and emplacement or as a component recycled into the mantle source by subduction. 2. An enriched group, comprising both a subgroup of samples mildly enriched in incompatible elements (Zr/Nb 12–18), and a subgroup of samples even more strongly enriched in LREE and especially Nb, with very low Zr/ Nb (4–7) and elevated Ti/V ratios. These trace element features are typical of the within-plate style of mantle enrichment, related to the overall incompatibility of trace elements during mantle melting under relatively “dry” conditions. Although markedly radiogenic, Nd isotope signatures vary from elevated εNd600 (from +6.7 to +8.2) in the mildly enriched subgroup, to lower εNd600 values (from +5.1 to +3.8) in the strongly enriched subgroup, which includes a metatrachyte with εNd600 of +3.2. The decoupling observed between enriched trace element characteristics (e.g., LREE-enrichment, which requires an enriched source) and radiogenic Nd isotopes (which imply that the mantle source was depleted in Nd relative to Sm on a time-integrated basis) may suggest that the depleted source region was metasomatized shortly prior to the igneous event. Alternatively, recent magma mixing between depleted (MORB-like) and enriched melts could have been involved. In this case, mixtures containing between ~60 and ~80–90% of the depleted end-member would exhibit LREE enrichment while retaining positive εNd values (Anderson, 1982).
DISCUSSION Interpretation of Geochemical Results The results obtained in this study generally corroborate the conclusions of earlier geochemical investigations (e.g., Waldhausrová, 1997a, and references therein), which revealed the presence of several distinct igneous suites among the Neoproterozoic low-grade metabasalts from the Barrandian area, specifically, an earlier tholeiitic series, that was followed by an alkaline series through a stratigraphically intermediate transitional series. Our new results, based on alteration-resistant trace elements and Nd isotope data, allow us to gain further insight into the mantle sources involved and to try to assess in a more detailed manner the geotectonic significance of these metabasalts. First, Nd isotope signatures combined with Th/Nb systematics rule out any significant role for crustal assimilation during
magma ascent through the crust. Indeed, highly radiogenic signatures are measured in most samples, irrespective of their degree of enrichment in Th, Nb, and LREE. There is no obvious negative correlation between εNd values and Th/Nb, as would occur in case of assimilation of old sialic crust with low time-integrated Sm/Nd ratios (i.e., low εNd) and high Th/Nb ratios. Even the LREE-enriched samples from the second major group have distinctly positive εNd values, and their Th/Nb ratios are very low (0.09–0.10), unlike what would be observed in case of significant contamination by continental crust or sediments derived therefrom. This distinction implies that the two major groups of basaltic magmas parental to the Barrandian metavolcanics were emplaced into extremely attenuated crust, or even into a purely ensimatic (i.e., oceanic) setting, and that most, if not all, of their variability in terms of highly incompatible trace elements and Nd isotopes reflects mantle source processes. Second, trace element data emphasize a major, twofold distinction between contrasting types of mafic magmas that were extracted from different mantle sources, namely, a depleted group with supra-subduction zone affinities, and an enriched group resembling within-plate OIB. The generally primitive, incompatible element–depleted group resembles N-MORB in several characteristics, especially LREE depletion and highly radiogenic Nd isotopes. Because they are all notoriously prone to remobilization during low-grade alteration and metamorphism, it is not possible to make a reliable use of large-ion lithophile elements (LILE; alkalis and alkalineearths), whose enrichment provides a sensitive monitor of the addition of a subducted component to a depleted mantle source (e.g., Saunders and Tarney, 1984). However, multi-element diagrams normalized to N-MORB display negative anomalies of Nb (Fig. 3). Bearing in mind that Nd isotope data preclude crustal contamination as a possible cause, this style of HFSE/REE and Th fractionation is believed to be diagnostic of magmas produced by partial melting of depleted mantle that was metasomatized and fluxed by a hydrous component derived from a subducting oceanic plate. Whether the metasomatizing agent, enriched in LILE but deficient in HFSE, was a hydrous fluid (e.g., Eiler et al., 1998), a silicic melt (e.g., Prouteau et al., 2001), or both, is an open question. However, low-degree partial melting of the upper part of eclogitized, subducted crust, including any sedimentary component, is likely to occur at depths of 125–175 km beneath back-arc basins (Sinton and Fryer, 1987). The very high εNd values measured in the samples with supra-subduction zone affinity (depleted group) put some constraints on the nature of subducted sediments, that is, their ultimate continental or oceanic provenance. In the present case, it is clear that any subducted sedimentary component should have been derived from a juvenile source, characterized by radiogenic Nd isotope signatures. This derivation would favor a model involving either a trench starved of sediments or containing mostly volcaniclastic sediments derived from an ensimatic arc. Furthermore, the low concentrations, compared to those of N-MORB, of most “conservative” incompatible elements (i.e.,
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts those believed to be largely independent of the subduction component), such as Ti, Zr, Nb, HREE, and Y (Pearce and Peate, 1995), and their relative fractionation (e.g., high Zr/Nb or Yb/Nb ratios) suggest that the depleted group was generated from a NMORB source mantle that was already depleted in incompatible elements, as commonly observed in intraoceanic arcs with active back-arc basins (Pearce and Parkinson, 1993). The later-stage group of samples enriched in incompatible elements, particularly Nb and LREE, does not show any feature of subduction-related magmatism. Instead, this group has clear affinities with within-plate magmas, such as OIB, believed to be generated by a relatively low-degree of partial melting from enriched mantle sources. In very general terms, a popular model relates such magmas to very deep mantle sources connected to the surface by narrow ascending plumes (e.g., Hofmann, 1997), without direct relationship to plate tectonic processes. However, the geochemical data themselves can hardly provide unambiguous evidence on the ultimate depth of mantle reservoirs. Indeed, geological observations favor alternative models based on the occurrence of enriched domains fairly ubiquitous in the shallow mantle (including supra-subduction wedges; e.g., Morris and Hart, 1983; Gill, 1984) where they could form a so-called “perisphere” layer (Anderson, 1995). This enriched reservoir might consist of a relatively refractory mantle veined by enriched components, such as frozen low-degree partial melts extracted from the underlying mantle. The incompatible element composition of partial melts from such metasomatized sources would be dominated by vein materials (e.g., Wood et al., 1980). This enriched reservoir might be tapped wherever the overlying lithosphere fails under extensional stress, thereby allowing low-degree melts to egress (Anderson, 1995). Based on these concepts, it is inferred that the OIB-like magmas forming late-stage volcanic build-ups in the Neoproterozoic Barrandian basin correspond to relatively low-degree partial melts that are extracted, as a response to limited extension, from mantle sources that were not dominated by subduction-related components, but instead resembled enriched domains believed to be broadly ubiquitous beneath lithospheric plates. Possible Geological Implications The geochemical and Nd isotope data clearly favor an intraoceanic supra-subduction zone setting for the earlier, depleted group of the Barrandian metabasalts. More specifically, the relatively minor contribution of subduction-related component in several samples otherwise similar to N-MORB is reminiscent of extensional back-arc basins, which are commonly floored by MORB and basalts (commonly referred to as “back-arc basins basalts” (BABB; Fryer et al., 1981) that are geochemically transitional between MORB and island arc tholeiites (e.g., Saunders and Tarney, 1984; Volpe et al., 1987). Indeed, a back-arc or inter-arc rift environment would be in keeping with the thick pile of submarine clastic sedimentary rocks (mostly turbiditic graywackes) associated with these early-stage volcanics, and this geodynamic
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setting has already been advocated for a long time (see references in Krˇibek et al., 2000). In terms of modern analogue, the depleted metabasalts might correspond to magmas emplaced into intraoceanic back-arc basins, such as the Lau basin (e.g., Hawkins, 1995; Pearce et al., 1995) or the Sumisu rift in the Izu-Bonin arc (e.g., Hochstaedter et al., 1990), where very high sediment accumulation rates (up to 4 km/Ma) have been reported (Marsaglia et al., 1995). The incipient, rifting stage of back-arc basin opening is commonly accompanied by the emplacement of a bimodal association of basalts and Na-rich felsic lavas, as exemplified by the nascent Sumisu rift (Hochstaedter et al., 1990) or the Northern Mariana Trough (Gribble et al., 1998). In contrast, only basaltic magmas are erupted during the more mature spreading stage, which may begin when the back-arc basin is 100–150 km wide (Gribble et al., 1998). The lack of significant amounts of felsic lavas in the studied area might therefore suggest that the basaltic magmas of the depleted group were formed by decompression melting in a relatively mature extensional setting. Moreover, by analogy with the systematic variation of the composition of the subduction-related component with the distance to the arc observed in the Lau basin (Pearce et al., 1995), the clear Th signal documented in the Barrandian samples (Fig. 3) might be tentatively interpreted to reflect generation relatively close (~50 km or less) to the arc. Although an ensimatic back-arc basin setting is favored by geochemical features of volcanic rocks, the preliminary Nd isotope data obtained on sedimentary rocks (e.g., model ages far in excess of deposition ages) indicate a substantial contribution from continental crustal components, mixed with a juvenile, presumably arc-derived component. Indeed, a few 2.0-Ga-old grains were found among detrital zircons of a sample of graywacke from the upper part of the Neoproterozoic Barrandian, with a maximum deposition age of 564 ± 16 Ma (Drost et al., 2004). Although these 2-Ga grains might have suffered multiple sedimentary recycling and do not necessarily imply a direct provenance from Paleoproterozoic basement, their presence and the scarce Nd isotope data available suggest that the intraoceanic arc/back-arc system inferred from the chemical and isotopic signature of igneous rocks was not far removed from a continental land mass, at least during the latest Neoproterozoic. Based on major and trace element discrimination diagrams, Drost et al. (2004) favored a back-arc setting of deposition and a continental island arc provenance for Neoproterozoic graywackes. A composite provenance, involving juvenile volcaniclastic components from an oceanic island arc on the one hand, and epiclastic components from continental source(s) on the other hand, might provide an alternative interpretation. The association (e.g., at Mitov and Koterov) of enriched, OIBlike pillow-lavas with sediments deposited in an inter- or supratidal environment (Pouba et al., 2000) shows that the late-stage, enriched magmas were able to build significant volcanic edifices, forming seamounts and even islands during the late evolutionary stage of the “spilitic series” of the Barrandian Neoproterozoic. Basalts containing a nonsubduction-related component, enriched in Nb, Zr, and LREE and geochemically similar to E-MORB or
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OIB, are not rare in recent arc and back-arc settings (e.g., Morris and Hart, 1983; Gill and Whelan, 1989). Examples from backarcs include the North Fiji Basin (Price et al., 1990), the Sumisu rift (Hochstaedter et al., 1990), and particularly the Japan Sea, where mildly alkaline and alkaline basalts form seamounts and volcanic islands (e.g., Pouclet et al., 1995). The source heterogeneity documented by these rocks has been variably interpreted to reflect either veins or blobs of enriched mantle contained in a more depleted matrix (e.g., Wood et al., 1980), or injection of new mantle. The first class of models may account for the generation of enriched magmas during the early stages of rifting, through preferential melting of shallower enriched domains as a result of higher geothermal gradients. In contrast, the occurrence of enriched magmas at a late-stage, and the switch from early subduction-related to younger within-plate magmatism, would rather favor the alternative hypothesis invoking injection of new mantle. Vanishing or even missing subduction-related components are indicated in the late-stage, enriched Barrandian metabasalts. This dearth in turn suggests that the underlying upper mantle was no longer fluxed by fluids and/or melts from a descending slab and that mantle flow allowed injection of new, more enriched material in the melt source, replacing the former hydrous mantle wedge. Similar changes in recent arc and back-arc systems have been interpreted to reflect “plume mantle” flowing from beneath the subduction hinge into the back-arc region, either laterally around the edges or through gaps (tears) of retreating subducting slabs (Schellart, 2004, and references therein). However, the mantle flow compensating for the retreat (roll-back) of subducting slab may also involve enriched mantle dragged from the base of the overriding plate toward the back-arc basin (e.g., Martinez and Taylor, 2002). In both cases, mantle flow is associated with the retrograde motion of a subduction hinge (roll-back) that occurs as a natural consequence of the negative buoyancy of sufficiently old subducting oceanic slabs with regard to the surrounding mantle (Elsasser, 1971). For the same reason, an extensional regime commonly prevails in the overriding lithospheric plate (Hamilton, 1995). In the Barrandian case, such a retreating slab scenario could account for (1) the inferred, strong extension of the overriding plate, with opening of a basin, accompanied and followed by BABB magmatism and the deposition of thick clastic sediments; (2) the extinction of the supra-subduction zone geochemical component; and (3) the switch to enriched, within-plate style magmatism. An alternative interpretation would postulate not a mere ocean-ward migration of the subduction zone, but its death. For example, this extinction might have occurred as a result of impingement of a spreading ridge with the intraoceanic trench, leading to mutual annihilation and switch to a transform plate boundary. This process would be accompanied by the opening of a “slab window” beneath the upper plate of the extinct subduction zone, which would allow for the upflow of new mantle material in the former arc and back-arc system, thereby accounting for the generation of late-stage, enriched basalts. In the absence of precise chronological and tectonic constraints, it is not yet possible to favor any of these highly tentative interpretations.
Comparison with Other Late Proterozoic Segments The closer example of relatively well-preserved Precambrian formations occurs in the Brno Massif, which corresponds to the exposed part of the largely covered Bruno-Vistulicum basement block (Fig. 1, inset). On the basis of geochemical and Sr-Nd isotopes, three contrasting crustal blocks (“terranes”) were recognized (Finger and Pin, 1997; Finger et al., 2000a). The eastern Slavkov terrane consists mainly of ca. 590-Ma quartz-diorites, tonalities, and granodiorites with relatively primitive isotope signatures (87Sr/86Sri ~0.704–0.705; −3 < εNdi < –1). In marked contrast, the western Thaya terrane displays K-rich granitoids, also emplaced ca. 580 Ma, but characterized by crustal isotopes (87Sr/86Sri ~0.708–0.710; −7 < εNdi < –4). The intervening BrnoBreclav terrane consists of a fault-bounded belt of volcanic and plutonic mafic rocks that probably corresponds to a suture. A 725 ± 15-Ma Pb/Pb zircon evaporation age was measured on a subordinate metarhyolite (Finger et al., 2000b), interpreted to be cogenetic with the overwhelming mafic rock-types based on similar εNd 725 values of +6.8. Recent intraoceanic subduction systems may reach very great cumulative length (e.g., some 2500 km along the Izu-BoninMariana or the Tonga-Kermadec arc systems). This observation suggests that even larger-scale correlations throughout Cadomian Europe should be attempted. As long recognized, the Lower Brioverian series from northern Armorican Massif (northwest France) show similarities with the Central Barrandian domain. Sedimentary rocks consist of monotonous terrigeneous sediments with mixed volcaniclastic and reworked continental provenance (Dabard, 1990). A typical feature is the frequent occurrence of interbedded black cherts that are interpreted to represent silicified terrigeneous and evaporitic deposits (Dabard, 2000), as also inferred in the Barrandian. Concerning igneous rocks, an episode of intra-arc or back-arc extension is documented ca. 610 Ma, as exemplified by (1) the Paimpol spilites (containing minor rhyolites dated 610 ± 9 Ma) with arc tholeiite affinities (Egal et al., 1996) and εNd of approximately + 6 (Dabard et al., 1996); (2) the Erquy spilitic series, dated 608 ± 7 Ma (Cocherie et al., 2001) and the broadly equivalent Lanvollon bimodal suite, with still rather uncertain geodynamic settings (back-arc basin according to Cabanis et al., 1987; within-plate rift according to Lees et al., 1987, and Egal et al., 1996); and (3) the Yffiniac-Belle-Isle-en-Terre gabbros, emplaced 602 ± 8 Ma and 602 ± 4 Ma, respectively (Guerrot and Peucat, 1990). Certain amphibolites associated with these gabbros were reported to be chemically similar to oceanic tholeiites (Chantraine et al., 2001). Other scattered mafic volcanics occur within the Lower Brioverian series of the Lamballe and St. Lô formations, composed of terrigeneous sediments containing interbedded black cherts. The mafic volcanics show either a within-plate alkaline affinity in the Lamballe Formation (Cabanis et al., 1987) or a similarity with NMORB in the St. Lô Formation (Dupret et al., 1990). Close to the southeastern end of Armorican Massif, the Precambrian Mauges Group is composed of a several-kilometer-thick monotonous
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts sequence of clastic rocks ranging from quartz-feldspar metagraywackes to metapelites, together with black cherts and graphitic schists. Mafic meta-igneous rocks (gabbros, lava flows, breccias, dikes, and cinerites) display geochemical features similar to those of transitional tholeiitic basalts emplaced in a within-plate extensional setting (Cabanis and Wyns, 1986). In summary, a broadly extensional tectonic regime is clearly indicated by ca. 610–600-Ma mafic or bimodal volcanic suites, with both back-arc and within-plate geochemical affinities. However, Nd isotope data are too scarce to monitor the degree of contamination by crustal materials or to assess the ensialic versus ensimatic setting of the Lower Brioverian basin(s). This episode was bracketed by arc-related plutons dated from ca. 750 Ma to ca. 625 Ma (with strongly radiogenic Nd isotope signatures; Samson et al., 2003), and ca. 580-Ma syn- to late-kinematic diorite intrusions (e.g., Nagy et al., 2002, and references therein) with negative εNd values, interpreted to reflect generation in an active continental margin containing an ancient basement, such as the 2.1-Ga relics (Icartian) documented in that region. The Sierra Morena (southwest Spain) exhibits Precambrian rocks variously overprinted by Cadomian and Variscan metamorphisms and deformations (e.g., Quesada, 1990; Eguiluz et al., 2000). Among these formations, the Serie Negra consists of a >5-km-thick succession of graphite-rich, turbiditic metapelites and metagraywackes containing laminated black cherts and marbles (Quesada, 1990). This series has been subdivided into a lower group (Montemolin Group) made of thin-bedded quartzrich graywackes and graphite-rich pelites with black cherts and local marbles, containing abundant amphibolites in its upper part, and an upper one (Tentudia Group), consisting of progressively thicker beds of massive graywackes that contain an increasing proportion of calc-alkaline volcanic clasts. The maximum age of deposition of the Tentudia group is constrained by the 564 ± 30-Ma age of the youngest detrital zircons dated in a metagraywacke (Schäfer et al., 1993). The youngest detrital grains measured in a graphite-rich metapelite from the Montemolin Group gave an average SHRIMP age of 591 ± 11 Ma (Ordoñez Casado, 1998). Ages of 600 and 610 Ma have been reported for the igneous protoliths of amphibolites from the Montemolin Group and the Badajoz-Cordoba shear zone, respectively (Schäfer, 1990, as cited in Bandres et al., 2002, 2004). Sm-Nd analysis of a sample from these broadly tholeiitic rocks gave an εNd600 value of +7.4 (Ordoñez Casado, 1998), indicative of a mantle source with strong time-integrated depletion of LREE. In the absence of more detailed geochemical studies, it is not possible to infer whether these rocks reflect subduction- or nonsubduction-related extensional setting. However, the sedimentary successions bear similarities with Lower Brioverian formations of Armorican Massif and with the Central Barrandian graywackes. Moreover, dioritic plutons of probable continental arc affinity were emplaced ca. 580 Ma (Merida Massif; Bandres et al., 2004), a situation reminiscent of that documented in northern Armorican Massif and the Brno Massif. In the southeast of the Ossa-Morena zone, conglomerates, arkoses, and shales of lowermost Early Cambrian
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age unconformably overlie the San Jeronimo Formation, composed of andesites interbedded with conglomerates, sandstones, and lutites, with a minimum thickness of 1 km. Based on microfossil evidence and stratigraphic context, the San Jeronimo Formation was ascribed to the Varangerian glacial stage (Quesada et al., 1990), possibly indicating a ca. 580-Ma age by correlation with the Gaskiers Formation glacial deposits in Newfoundland (Bowring et al., 2003). The andesites show typical calc-alkaline chemistry and strongly radiogenic Nd isotope signatures (Pin et al., 2002), implying a time-integrated depleted mantle source. However, sedimentary rocks interbedded with the andesites have negative εNd values indicative of continental crust sources, which preclude a purely intraoceanic arc setting. These combined pieces of evidence suggest generation in a supra-subduction environment located on relatively juvenile crust, such as an island arc previously accreted to a continental margin. This short review shows that subduction-related arcs and back-arc basins were a general feature throughout the European domain during the 750–580-Ma period, as part of a much wider domain extending from present-day northwest Africa (Morocco) and Avalonia to the Arabo-Nubian shield (Fig. 7). The opening and spreading of back-arc basins ca. 610 Ma is suggested to have occurred in all three European examples, but whether only ensimatic or both ensialic and ensimatic basins were involved is still
Figure 7. Terrane distribution within the Avalonian-Cadomian realm (after Nance and Murphy, 1994) in the Neoproterozoic continental reconstruction of Torsvik et al. (1996). A—Armorican Massif; B—Bohemian Massif. From Krˇíbek et al. (2000).
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an open question, given the lack of adequate geochemical data. These basins acted as efficient traps for detrital sediments derived from both juvenile arc-related and recycled continental sources that played a significant role in local crustal growth. Overall, this picture is reminiscent of modern arc and back-arc systems from the western Pacific region, where large intraoceanic subduction systems fringe the major continental masses of the Asian and Australian plates, along with a complex mosaic of microplates and magmatic arcs that include intervening basins floored either by oceanic crust or attenuated continental crust. It is speculated that the Neoproterozoic paleogeography of the Cadomian realm broadly resembled such patterns. CONCLUSION Based on combined Nd isotope and trace element evidence, the spilitized basalts of the Central Barrandian Neoproterozoic reflect contrasting magmas extracted from fairly different mantle sources. During an earlier stage (tentatively dated at 605 ± 39 Ma based on a whole-rock Sm-Nd isochron), basalts broadly similar to N-MORB but showing negative anomalies of HFSE were extracted from a source that was strongly depleted in LREE on a time-integrated basis and emplaced as lava flows in a strongly subsiding sedimentary basin. Subsequently, Nb-rich basalts extracted from enriched mantle sources were emplaced as shallow to subaerial volcanic edifices. This kind of evolution of mafic magmatism is reminiscent of some recent intraoceanic back-arc basins, where a switch from supra-subduction zone to withinplate–like magmatism is documented. This change might have occurred either simply because of ocean-ward migration of the subduction zone or as a result of impingement of a spreading ridge with the intraoceanic trench, leading to mutual annihilation and evolution to a transform plate boundary. Even though the geochemical affinities of igneous rocks clearly favor an ensimatic setting, the presence of a thick sedimentary pile containing both juvenile and recycled old crustal components suggests that the inferred Late Proterozoic intraoceanic arc and back-arc system was located near a continental mass. ACKNOWLEDGMENTS This study was supported by travel grants to one of us in the scope of the Czech-French cooperation (Barrande Project). We are grateful to Mr. Dašek for drawing some of the figures. Constructive reviews of the manuscript by Dr. R. D’Lemos and Dr. S. Samson are gratefully acknowledged. This article is a contribution to the International Geological Correlation Program Projects 453 and 497. APPENDIX: SAMPLE LOCATIONS Kli-5: Metabasalts (tholeiitic), outcrops of a massive flow along the road N of the Klícˇava dam, north of Zbecˇno. Main central volcanic belt in the Rakovník area.
Zb-1: Metabasalts (tholeiitic), from the Zbecˇno quarry, opposite the railway station. Main central belt in the Rakovník area. Rou-1: Porphyritic basaltic meta-andesite, outcrops on the Roupov castle hill (southwest of Prˇeštice). Main central belt in the Prˇeštice sector. Kru˜s1: Metabasalts (tholeiitic), massive flow in the Krušec quarry near Chudenice village. Main central belt, in the Klatovy area. Teb-1: Metabasalt, massive flow in a small abandoned quarry east of Trˇebobuz village, southwest of Všeruby. Strˇíbro volcanic belt. Lhv-2: Metabasalt, outcrops of a massive flow on top of the hill northeast of Luhov village, southwest of Všeruby. Strˇíbro volcanic belt. Kli-1: Metabasalt from an abandoned quarry, at the northern end of the Klícˇava dam, north of Zbecˇno. Main central belt in the Rakovník area. La-2: Basaltic meta-andesite, outcrops of a massive flow on the hill Cˇ ihadlo, south of Lány village. Main central belt in the Rakovník area. UP-1: Basaltic pillow lava, “Kneˇžská skála,” rocky outcrops on the left bank of the Berounka River near Nezabudice village, north of Skryje. Main central belt in the Hrˇebecˇníky sector. UP-2: Basaltic pillow lava, “Cˇertova skála,” rocky outcrops on the left bank of the Berounka River near the Týrˇovice village, north Skryje. Main central belt in the Hrˇebecˇníky sector. Chrˇ-1: Metabasalt (transitional), outcrops of a massive flow in the vicinity of Chrˇícˇ village, northwest of Zvíkovec. Main central belt in the Hrˇebecˇníky sector. Li-1: Vitrocrystalloclastic, basaltic laminated tuff, outcrops from the locality “Liška” in the vicinity of Polenˇ village. Main central belt in the Klatovy area. Reb-1: Metabasalt, outcrops of a massive flow near Řebrˇí village, in the vicinity of Svojšín village, west of Strˇíbro. Svojšín volcanic belt. Lit-1: Metabasalt, massive flow from the Litice quarry, south of Plzenˇ. Main central belt in the Plzenˇ area. Lit-2: Actinolitized basalt in the Litice quarry. Zchl-1: Metabasalt, outcrops of a massive flow near Záchlumí, NNW of Strˇíbro village. Svojšín volcanic belt. Kli-3: Mugearite, outcrops at the northwest end of the Klícˇava dam north of Zbecˇno village. The main central belt in the Rakovník area. Kot-1: Metabasalt from the abandoned quarry near Koterov village, southeast of Plzenˇ town. Main central belt in the Plzenˇ area. Boro-1: Transitional basalt, small outcrops near Borovno village, northeast of Blovice town. The southern volcanic zone in the Blovice sector. Mit-1: Alkali basalt, massive flow from the quarry west of Mítov village, east of Blovice town. Southern volcanic zone in the Blovice sector. Mit-2: Alkali basalt, pillow lava, from the Mítov quarry. Si-2: Devitrified glassy trachyte, massive flow from the Slatina quarry, southwest of Rakovník town. Slatina-Pavlíkov strip. Viš-1: Graywacke, outcrops on the left bank of the Berounka River, at the locality Višnˇová near Roztoky u Krˇivoklátu. Main central belt in the Hrˇebecˇníky sector. Nml-1: Graywacke, outcrops on the left bank of the Berounka River, Nezabudice mill near Nezabudice village. Main central belt in the Hrˇebecˇníky sector. Krus-2: Black shale from an intercalation in basaltic flows, Krušec quarry near Chudenice village. Main central belt in the Klatovy area.
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Saunders, A.D., and Norry, N.J., eds., Magmatism in the ocean basins: London, Geological Society of London Special Publication 42, p. 313–345. Tanaka, T., Togashi, S., Kamioka, H., Amakawa, H., Kagami, H., Hamamoto, T., Yuhara, M., Orihashi, Y., Yoneda, S., Shimizu, H., Kunimaru, T., Takahashi, K., Yanagi, T., Nakano, T., Fujimaki, H., Shinjo, R., Asahara, Y., Tanimizu, M., and Dragusanu, C., 2000, JNdi-1: A neodymium isotopic reference in consistency with LaJolla neodymium: Chemical Geology, v. 168, p. 279–281. Taylor, S.R., and McLennan, S.M., 1985, The continental crust: Its composition and evolution: Oxford, Blackwell Scientific, 312 p. Tomek, Cˇ., and Dvorˇáková, V., 1994, Deep seismics in western Bohemia, in Geological model of western Bohemia in relation to the deep borehole KTB in the FRG. Abstracts: Prague, Czech Geological Survey, p. 44–47. Tomek, Cˇ., Dvorˇáková, V., and Vrána, S., 1997, Geological interpretation of the 9HR and 503M seismic profiles in western Bohemia, in Vrána, S., and V. Šteˇdrá, V., eds., Geological model of western Bohemia related to the KTB borehole in Germany: Sborník Geologických Veˇd, Geologie, v. 47, p. 43–50. Torsvik, T.H., Smethust, M.A., Meert, J.G., Van der Voo, R., Kerrow, W.S., Brasier, M.D., Sturt, B.A., and Walderhaugh, H.J., 1996, Continental break up and collision in the Neoproterozoic and Palaeozoic—A tale of Baltica and Laurentia: Earth Science Reviews, v. 40, p. 229–258, doi: 10.1016/0012-8252(96)00008-6. Vavrdová, M., 2000, Microfossils in carbonaceous cherts from Barradian Neoproterozoic (Blovice Formation, Czech Republic): Veˇstnik Cˇeského geologického ustavu, v. 75, p. 351–360. Volpe, A.M., Macdougall, J.D., and Hawkins, J.W., 1987, Mariana Trough basalts (MTB): Trace element and Sr-Nd isotopic evidence for mixing between MORB-like and arc-like melts: Earth and Planetary Science Letters, v. 82, p. 241–254, doi: 10.1016/0012-821X(87)90199-3. Waldhausrová, J., 1997a, Proterozoic volcanics geochemistry and mineral chemistry: A contribution to the Barrandian Upper Proterozoic stratigraphy (Bohemian Massif, Czech Republic): Krystalinikum, v. 23, p. 151–180. Waldhausrová, J., 1997b, Geochemistry of volcanics (metavolcanics) in the western part of the TBU Precambrian and their original geotectonic setting, in Vrána, S., and Šteˇdra, V., eds., Geological model of western Bohemia related to the KTB borehole in Germany: Sborník Geologických Veˇd, Geologie, v. 47, p. 85–90. Wharton, M.R., Hathway, B., and Colley, H., 1995, Volcanism associated with extension in an Oligocene-Miocene arc, southwestern Viti Levu, Fiji, in Smellie, J.L., ed., Volcanism associated with extension at consuming plate margins: London, Geological Society of London Special Publication 81, p. 95–114. Wood, D.A., Joron, J.L., and Treuil, M., 1979, A re-appraisal of the use of trace elements to classify and discriminate between magma series erupted in different tectonic settings: Earth and Planetary Science Letters, v. 45, p. 326–336, doi: 10.1016/0012-821X(79)90133-X. Wood, D.A., Joron, J.L., Marsh, N.G., Tarney, J., and Treuil, M., 1980, Major and trace element variations in basalts from the north Philippine Sea drilled during Deep Sea Drilling Project Leg 58: A comparative study of back-arc basin basalts with lava series from Japan and Mid-Ocean Ridges, in de Vries, K.G., and Kobayashi, K., et al., Initial Reports of the Deep Sea Drilling Project, v. LVIII: Washington, D.C., National Science Foundation, p. 873–894. Zindler, A., 1982, Nd and Sr isotopic studies of komatiites and related rocks, in Arndt, N.T., and Nisbet, E.G., eds., Komatiites: London, George Allen and Unwin, p. 399–420. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
Structural evolution of the Prague synform (Czech Republic) during Silurian times: An AMS, rock magnetism, and paleomagnetic study of the Svatý Jan pod Skalou dikes. Consequences for the nappes emplacement Tahar Aïfa* Géosciences-Rennes, CNRS UMR6118, Université de Rennes 1, Bat.15, Campus de Beaulieu, 35042 Rennes Cédex, France Petr Pruner Martin Chadima Petr Štorch Institute of Geology, Academy of Sciences of the Czech Republic, Rozvojova 135, 165 02, Prague 6, Czech Republic
ABSTRACT Silurian effusive basalts and volcaniclastics compose the Svatý Jan volcanic center, which is located in the northwestern limb of the Prague synform, where three major volcanic phases have been recognized: the first one of early to mid-Wenlock and the last of mid-Ludlow age. Two alkaline basalt dikes of late Wenlock to mid-Ludlow age, respectively tilted to the west and to the northeast, as observed in a 100-m-thick tuff sequence, which represents the second volcanic phase, have been extensively sampled. An anisotropy of the magnetic susceptibility (AMS) study of seventy-nine specimens taken from a 5-m-thick dike (dike1) and thirty-two specimens cored in a 3.5-m-thick dike (dike2) shows two different fabrics, carried mainly by Ti-magnetite and/or magnetite, which are considered to be related to the transtensional opening phase of the dikes. Four components of magnetization, attributed to Middle-Late Silurian (C1), Middle-Late Carboniferous (C2), Cretaceous (B), and Paleocene (D), in agreement with already-published directions for the Bohemian Massif, have been isolated. They are carried by Ti-magnetite for components C1 and C2, hematite and goethite for components B and D. The opening mode, which controlled both dikes, corresponds to a dextral transtensional regime, as deduced from the AMS K1 axis. They may have been opened during several magmatic stages related to different injections during late Wenlock to mid-Ludlow times. The first stage is dominant and controlled by the primary fabric, which is mainly oblate. With a NNW-SSE strike, perpendicular to the shortening direction, this fabric is in agreement with the direction of emplacement of the nappes during the Late Devonian. At that time the nappes emplacement that *E-mail:
[email protected]. Aïfa, T., Pruner, P., Chadima, M., and Štorch, P., 2007, Structural evolution of the Prague synform (Czech Republic) during Silurian times: An AMS, rock magnetism, and paleomagnetic study of the Svatý Jan pod Skalou dikes. Consequences for the nappes emplacement, in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, p. 249–265, doi: 10.1130/2007.2423(11). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Aïfa et al. pre-dates this direction was probably associated with the sinistral closure of the Rheic Ocean, in agreement with post-Givetian folding and faulting, which deformed the synform infill and closed the Barrandian marine sedimentary cycle. Keywords: Prague synform, dikes, stress, rheic, AMS, magnetization, Variscan orogen
INTRODUCTION The Paleozoic evolution of the Prague basin (Czech Republic) before its synformal deformation has been a new topic of interest for several years (e.g., Melichar, 2004). Many results previously obtained thanks to tectonics studies, sedimentology, and paleontology are actually in contradiction with the most recent data. It is now considered that the evolution of the Prague synform has been mainly controlled by allochtonous units. The previous results were mainly based on the sedimentary and volcanic history of the present synform, but few data were published regarding its structural evolution (Matte et al., 1990; Matte, 1991, 2001; Havlícˇek, 1998; Kachlík and Patocˇka, 1998; Žák et al., 2005). Recent work on zircon dating using U/Pb sensitive highresolution ion microprobe (SHRIMP) and Nd isotopic analysis has been published by Linnemann et al. (2004) in which they propose a geodynamical model suggesting a southwest transport direction of the ophiolitic complexes between the Moldanubian and Saxo-Thuringian zones from the Late Devonian until the Lower Carboniferous, that is, during the ultimate stage of closure of the Rheic Ocean (Kröner and Hahn, 2003). The purpose of this article is to shed some light on this problem. The Rheic Ocean, the existence of which has been demonstrated in eastern North America (Keppie et al., 2000; McKerrow et al., 2000; Murphy and Nance, 2003), is still widely debated in western Europe, mainly because its width cannot be estimated using paleomagnetic data because the orientation is northwest-southeast. The use of the anisotropy of the magnetic susceptibility (AMS) to constrain the mode of opening of the dikes combined with the paleomagnetic technique, which can be used for dating the fabrics, represents a useful tool to check the direction of the stress existing at the time of the opening. Consequently, the direction of the displacement of the nappe is possibly related to the closure of the Rheic Ocean, if the latter really existed. Because dikes are good stress indicators, we first check these two techniques on our two dikes, which are both supposed to be Silurian in age (Štorch, 1987). We then examine the magnetic mineralogy of the main carriers for better constraints on the magnetic components and their magnetic ages. GEOLOGICAL BACKGROUND The Prague synform, which is preserved in the central part of the Barrandian area (Bohemian Massif) comprises a pile of Ordovician, Silurian (Fig. 1), and Devonian rocks more than 2.5 km thick. Unmetamorphosed sediments, moderately deformed by the Variscan orogeny and famous for their fossils and their detailed
stratigraphy, outcrop in the Prague synform. The sedimentation was associated and temporarily disturbed by rather intensive and largely submarine basaltic volcanism. Basaltic volcanics first appeared during the late Early Ordovician and then formed the large Komárov complex in the southwestern part of the newly originated basin. This volcanism culminated in the late Llanvirn Series and again, but less intensively, in the late Caradocian Series. It revived again in the Early Silurian but remained localized to the northern limb of the central and northeartern parts of the present Prague synform. The last isolated submarine eruptions of basaltic magma known from the late Emsian succession occurred in the central part of the Prague synform (Havlícˇek, 1987). Extensive outcrops of Silurian effusive basalts and volcaniclastics (Patocˇka et al., 1993) belonging to the major Svatý Jan volcanic center are located between Beroun-Lištice, Svatý Jan pod Skalou, Záhrabská, and Lodeˇ nice. Three major volcanic phases have been recognized in the Svatý Jan center. The earliest phase started around the early mid-Wenlock (Chlupácˇ et al., 1998), the second phase is of late Wenlock age, and the latest ceased in about the mid-Ludlow. Two dikes made of alkaline basalt showing well-developed feldspar phenocrysts have been found, cropping out in a small gorge associated with the steep slope of the left bank of the Kacˇák Creek, which is located between Sedlec and Svatý Jan pod Skalou (Fig. 2), 600 m northeast of the Svatý Jan Monastery (49.975°N, 14.136°E). These two dikes have been sampled in detail (Fig. 3A). They are situated in the lower part of a 100-m-thick tuff sequence (corresponding to the second volcanic phase) and represent either volcanic channels feeding the upper part of the volcaniclastic succession or, more probably, fissures that supplied the basaltic magma to the lava shield, characterizing the third (and last) volcanic period. A late Wenlock to mid-Ludlow age can thus be assumed for these dikes. During the Middle Devonian (Givetian), the first Variscan orogenic movements of the early Bretonian phase (Havlícˇek, 1963) terminated the sedimentation in the Prague synform and uplifted the whole Barrandian area (Kukal and Jäger, 1988; Havlícˇek, 1998). Post-Givetian folding and faulting affected the synform infill and closed the Barrandian marine sedimentary cycle. As a result of these movements, the Silurian sediments actually dip toward the southeast (between 14° and 35°) in the studied area (Fig. 2). Tuffs, which rest slightly unconformably on the shales, dip at ~18°–45° above the contact (Fig. 3A). Dike1, which corresponds to the northwesternmost outcrop, is broadly north-south oriented and dips by 70° to the west. It is ~5 m in thickness and is characterized by wavy contacts on both sides. It was first suggested (Aïfa et al., 2002) that this dike was
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Structural evolution of the Prague synform (Czech Republic) during Silurian times emplaced in an original fissure, which may have been opened repeatedly and used by subsequent intrusions. This scenario was suggested by some magnetic data and clear internal planar surfaces at 1 m from the dike’s western contact and at 1.40 m from its eastern border. It is not necessarily always the case, however, because the difference in time between two successive intrusions can be very short and not discriminated by paleomagnetic data. The wall surface at the southwest contact of the dike has a strike of 78° and a dip of 70° at the site of the sampled profile. The same contact measured a few meters away gives values of 268°/78° and 262°/81°. The dip direction (strike) and magnitude (dip) of the northeast contact of the dike are 270° and 70°, respectively. Neighboring tuffs with volcanic bombs follow the bedding, although the bedding itself is difficult to measure. The dike steeply penetrates the tuffs. In the lower part of the outcrop the dike penetrates shales and thin-bedded limestones of mid-Wenlock age, dipping by 35° to the southeast (strike 120°). Two meters off the northeast contact there is a prominent strike-parallel tectonic plane within the dike. It may be considered as the boundary between two successive intrusions that used the same fracture. Dike2, exposed on a steep slope above the creek, dips 38° to the southeast (strike 145°; Fig. 2) and shows moderately sinuous contacts with the host rock. This dike displays feldspar phenocrysts, which are more or less parallel to its margins. This pattern can still be observed up to at least 10 cm from the dike margin. Neighboring tuffs close to the southwest contact exhibit a strike of 184° and a dip of 45°. Calcareous shales and laminated limestones of mid-Wenlock age cropping out below the tuffs are cut by the dike. Their topmost bedding plane, just below the overlying tuffs, dips by 30° to the south (strike 160°). There is a clear difference between the dips and the strikes of the sediments and those of the volcaniclastics. Tuffs may have been deposited on the slope of a volcanic cone (although no slumps are observed), or the volcanites may have been deformed by some magma flow before the deposition of the volcaniclastics. SAMPLING All together, seventy-nine specimens were taken from dike1 following one transverse (perpendicular to the edges) section and two dike-margin parallel sections with a spacing (cracks and rock weathering permitting) of ~10 cm (Fig. 4). In the first stage of our study, seven blocks were taken in order to cut pilot samples. They yielded fifteen cubic specimens named “SV1” to “SV7.” In the second phase, a portable drilling machine was used and gave us thirty-three core samples, which provided forty-three cylindrical specimens. Twenty-five core samples were also sampled in dike2 using the same device. They were completed later, with a total collection of thirty-two oriented specimens. Sampling was carefully made, taking care of the distances between specimens, but also of fractures, flow lines, and chilled margins. Sampling parallel to the borders was made to characterize the mode of opening of the dike (Aïfa and Lefort, 2000), because such samples illustrate the first stages of the emplacement mechanism (Smith et
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al., 1993; Lefort et al., 2006). If the opening of the dikes results from a transtensional opening, the contemporaneous stress directions can be deduced from it and dated through the components of magnetization. AMS is used to determine the type of magnetic fabrics recorded in the dikes and thus their regional geotectonic environment. Sampling along the edges is useful to characterize the magma flow in a 3-D space and may help to discover possible remagnetizations associated with fluid circulations or alterations. ANISOTROPY OF THE MAGNETIC SUSCEPTIBILITY In general, the opening of the doleritic dikes results from a vertical, oblique, or horizontal flow. In these conditions the regional stress is parallel to the trend of the dikes. This type of opening is responsible for a convergent lateral tiling of the feldspars (Blanchard et al., 1979; Moreira et al., 1999) and for a superimposed magnetic tiling (Aïfa and Lefort, 2000). In some rare cases the opening of the dikes is controlled by a transtensional mechanism (Lefort et al., 2006). In these conditions the petrographic and magnetic tilings show an oblique and identical trending of the tiling on both sides of the dike. This type of opening results from a regional stress oblique to the dike (Smith et al., 1993). The magnetic tiling (Aïfa and Lefort, 2000; Lefort et al., 2006) is usually associated with K1 or K2 (maximum and intermediate axes, respectively, of the AMS tensor). Thus it is important to check the K1 or K2 directions if we want to deduce the regional stress from the anisotropy tensors. Using KLY-3S Kappabridge (Agico Brno), AMS was measured for all samples. Anisotropy parameters, such as corrected anisotropy degree P′, shape parameter T, and direction of maximum and minimum magnetic susceptibilities (K1, K3), were counted using the tensor notation of AMS (Jelínek, 1978). Figure 3 shows the distribution of the various K1 values after the unfolding of each dike using the software of Pangaea Scientific (Stesky and Pearce, 1995) to draw the isocontours (Kamb 1959). For dike1 (Fig. 3B left), note that the eastern border shows a good cluster of K1 values with an oblique orientation with respect to the dike margin. The center of the dike shows a group of values that is clearly clustering along a great circle, oblique to the dike trend (but it also displays some mixture of the distribution along a great circle on the opposite side). The western border is nearly similar to the center, with two oblique distributions of K1, one being more clustered than the other. Therefore, we think that a postmagmatic emplacement fabric overprints the primary fabric. Because K1 shows the same obliquity with respect to the borders of the dike, we assume that we are dealing with the second type of opening (see above). In these conditions the grains contributing to the AMS results were necessarily aligned perpendicularly to the stress direction before the cooling of the magma. We can thus assume that the opening of this dike resulted from a dextral transtensional regime (Aïfa and Lefort, 2000; Lefort et al., 2006). The same data also show the existence of a shallow inclination of maximum axis K1 (between 16° and 21° from west to east; Fig. 3B).
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Dike2 (Fig. 3B right) has in its center a well-clustered K1 distribution, which suggests an east-west flow oblique with respect to the edges of the dike. K1 distribution is also oblique to the dike walls on both margins, which again suggests a dextral transtensional opening with a shallow dipping magma flow (between 28° and 3°; Fig. 3B). In the field, some imbricated (Blanchard et al., 1979) phenocrysts have also been observed more or less parallel to the margins of the dike (which may extend up to 10 cm away from the margins). Their distribution may correspond to the initial direction of the flow (Philpotts and Asher, 1994) before the transtentional opening.
values than the heating curve (Fig. 4A, B, D); the second (type II), less frequent, shows the heating curve presenting higher values of susceptibility than for the cooling curve (Fig. 4C, E, F, G). Three specimens from this second type were heated and cooled twice, and one of them (Fig. 4C) was submitted to an Argon flow (reducing environment). We note that the heating curve of the second run is close to the cooling curve of the first run. We also note that type I corresponds to the borders of the dike whereas type II corresponds to its center.
ROCK MAGNETIC ANALYSIS
Some pilot samples from both borders and the center of dike1 were submitted to isothermal remanent magnetization (IRM) to saturation to characterize their coercivity spectra. Sample SV2/1, located 5 cm from the eastern border of dike1, shows relatively low coercivity values (Fig. 5A), whereas sample SV6/2, located 101 cm from the same border, shows higher coercivity values (Fig. 5B). Our interpretation is that during the magma injection and opening of the dike, cooling of the magma in contact with the host rock is very rapid, which implies small magnetic grains (single to pseudo-single domain, SD). These grains grow and develop (multidomain size, MD) toward the center of the dike because cooling becomes slower toward the center. Hence, theoretically we may expect high susceptibility values toward the center of the dike if cooling is homogeneously axisymmetric (Moreira et al., 1999; Aïfa and Lefort, 2001; Lefort et al., 2006). The discrimination between the two values of coercivity could be related to the presence of different injection phases. One may expect that the high coercivity value is related to the SD size of magnetite or Ti-magnetite, whereas the lower coercivity value holds for the MD size of the same type of mineral, the normalized magnetic moment being within the same range.
Temperature Variation of Magnetic Susceptibility Pilot samples were used for a rock magnetic study. The temperature dependence of magnetic susceptibility measurements were carried out on several powder samples. Samples were heated to 700 °C and subsequently cooled down to room temperature in a CS-3 (Agico Brno) furnace (Parma and Zapletal, 1991). Magnetic susceptibility was simultaneously measured using a KLY3S Kappabridge (Jelínek and Pokorný, 1997). Thermomagnetic curves of pilot samples from the borders and the center of dike1 show that no alteration occurred in the center (Fig. 4) where the presence of magnetite or Ti-magnetite (Curie temperature ~580 °C) is dominant. This Ti-magnetite sometimes co-exists with either a small amount of possible goethite or hematite, because we distinguish an increase in the heating part of the curve after 120 °C (Fig. 4C, E, F, G) or after 540 °C (Fig. 4A, B, D). We are dealing here with two common “types” already mentioned by Hrouda et al. (2003): the most frequent of them (type I) shows a cooling curve with much higher susceptibility
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Variation of Magnetic Properties within the Dikes To study the mineralogical changes within the dikes, we used 3-D block diagrams to better illustrate the distribution of the bulk magnetic susceptibility K, the corrected anisotropy degree P′, the intensity of the natural remanent magnetization (NRM) M, and the shape parameter T as functions of distance, instead of plotting the scalar parameters along cross-sections. One purpose of this technique was to check the possible existence of abrupt changes of these parameters and thus to detect fractures or a possible magmatic “zoning.” To this end we measured within each dike the various magnetic characteristics with a (0,0) reference at the lower left corner of the diagram. Gridding using the krigging technique was applied and after interpolation, four diagrams were obtained for each dike; they are presented in Figure 6. For dike 1, they show a large area of high K (Fig. 6A) on the right side of the dikes, this high value also affects the other parameters of the diagrams—M (Fig. 6B), P′ (Fig. 6C), and T (Fig. 6D). Note that M and K exhibit a positive linear correlation. Most of the high values are concentrated along the eastern borders of both dikes. A perpendicular section crosscutting dike1 displays high values for NRM (NRMmax = 1.4 A/m) and the bulk susceptibility (Kmean = 0.06 SI). A profile parallel to the eastern border also shows higher values for both K and M (Fig. 6A, B). An east-west section across dike1 (Fig. 6C) shows evidence of two major P′ peaks, whereas a single value can be seen on dike2. The peak is located 1 m from the eastern edge of dike1 and reaches a value of 1.03. The other peak, which reaches 1.06, is located ~1.40 m from the dike’s western edge. Both peaks correspond on the field to the two fractures that may have resulted from the rejuvenation of previous north-south cooling cracks. We speculate that this possible rejuvenation mainly affected the easternmost linear crack zone, because this structure shows a P′ value higher than 1.05 (Puranen et al., 1992). If we accept this criterion, this rock disruption could be considered as a fracture. The linear structure located in the west, which attains a value of only 1.03, could also be considered as a possible fracture. COMPONENTS OF MAGNETIZATION The processing of these dolerites also included a progressive thermal demagnetization using the MAVACS (Magnetic Vacuum Control System [Geofyzika Brno]; Prˇíhoda et al., 1989) equipment at temperatures ranging between 80 and 680 °C with step intervals of 60–30 °C. Demagnetization using alternating field (AF) technique has been applied using an LDA-3 apparatus (Agico Brno) until 100 mT, with steps every 5–20 mT. Separation of the remanent magnetization components was carried out with the help of multicomponent analysis (Kirschvink, 1980; Man, 2003). At each step of the thermal demagnetizations, the bulk magnetic susceptibility was measured to track any mineralogical changes. In fact, in an oxidizing environment, magnetic susceptibility versus temperature shows dramatically increasing
values (Fig. 7A4, B4). This increase is in agreement with a high (Fig. 7A3) or low (Fig. 7B3) unblocking temperature probably of Ti-magnetite (goethite; Fig. 7A4, B4), which transforms above 450 °C (120 °C) to magnetite (hematite), as shown by the normalized M/Mmax intensity values. For thermal demagnetization we used twelve specimens; twenty-six specimens have been subjected to AF demagnetizations. At least two components were extracted from each specimen, leading to the following components (Figs. 7 and 8; Table 1): • Component B is of low field and low blocking temperature. The computed mean direction seems to be slightly older than the present-day field. • Component C1 lies in a temperature range between 200 and 540 °C, and its AF demagnetization field is between 10–20 and 40–65 mT, reflecting probably the presence of magnetite or Ti-magnetite, with a component of magnetization (D = 204.3°, I = –15.2°, α95 = 7.9°). Two preliminary conclusions can be drawn: (1) the data fit the Middle to Late Silurian directions if we compare with the results obtained for black shales from the Kosov Quarry near Karlštejn, Bohemian Massif (D = 205°, I = –28°), with a paleorotation of 175–185° (Patocˇka et al., 2003); and (2) the magnetization measured in Silurian dikes is likely to be early Permian to late Carboniferous overprint. • Component C2 lies in the temperature range between 200 and 540 °C and its AF demagnetization field is between 10(20) and 40(65) mT, reflecting probably magnetite or Ti-magnetite. Its component of magnetization fits the Carboniferous direction for Bohemian Massif (Krs et al., 2001; Edel et al., 2003; Patocˇka et al., 2003). Tilt-corrected mean direction of remanent magnetization (D = 179.72°, I = 11.8°, α95 = 13.3°) corresponds to the Middle or Late Carboniferous direction for the Bohemian Massif (with no significant rotation). • Component D shows thermal demagnetizations (only samples SV5–SV7 were processed) with a temperature ranging between 580 °C (in some rare cases, 620 °C) and 680 °C and carried mainly by hematite, the AF demagnetization fields are between 20(40) and 80(100) mT, similar to the B component. For dike2 only seventeen specimens out of thirty-two were subjected to AF demagnetizations. The C1 component of magnetization is not recorded in this dike. The other components of magnetization (B, C2, and D) are isolated on the borders as well as in the center of the dike (Figs. 7 and 8; Table 1). As an example specimen SV2–26 from dike2 recorded B and C2 components that show anti-parallel directions in the orthogonal diagram (Fig. 7D1). In this case great circle analysis has been carried out to isolate the best-fitting component. It can be observed on a declination versus inclination plot of the distribution of C1 and C2 components (Fig. 8E) that C1 and C2 directions are never superimposed. In addition, the mean value of C1 inclinations is –15.2° with good cluster (α95 = 7.9°), C2
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N N Structural evolution of the Prague synform (Czech Republic) during Silurian times
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Figure 7. Examples of in situ stereographic projections of isolated components of magnetization B, C1, C2, and D for (panels A, B) dike1 and (panels C, D) dike2. Full circles: lower hemisphere, open circles: upper hemisphere. Orthogonal projections of thermal (in °C) (panels A2, B2 for dike1) or alternating field (in mT) (panels C2, D2 for dike2) demagnetizations. Open (solid) circles indicate projection onto the vertical (horizontal) plane. Normalized intensity of magnetization vs. temperature (panels A3, B3) and vs. demagnetizing field (panels C3, D3). Magnetic susceptibility vs. temperature showing mineralogical changes of probably hydroxides above 400 °C (panel A4) and 200 °C (panel B4).
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TABLE 1. COMPONENTS OF MAGNETIZATION ISOLATED IN THE STUDIED SITES (49.975°N, 14.136°E) VGP VGP Dike Position Cp N D I D I Lat Long Paleolat dp dm Lat Long Paleolat dp dm α95 α95 no. (geo) (geo) (geo) (tilt) (tilt) (geo) (geo) (geo) (geo) (geo) (tilt) (tilt) (tilt) (tilt) (tilt) (tilt) 1 All B 35 357.8 83.2 5.7 109.1 58.2 5.7 63.37 13 76.59 11 11.2 18.4 65 38.8 6.2 8.5 1 All C1 13 196.6 –9.3 7.9 204.3 –15.2 7.9 42.59 171.38 –4.68 4 8 43.2 160.1 –7.8 4.2 8.1 1 All C2 11 191.2 26.8 13 179.7 11.8 13.3 –25.1 2.13 14.17 7.8 14.4 –34.1 14.5 6 6.9 13.5 1 All D 29 2.3 85.9 6.4 113.3 56.7 6.4 58.13 14.75 81.84 12.6 12.7 15.2 63.3 37.3 6.7 9.3 1 East B 19 355.4 81.5 7.4 96.5 67.6 10.9 66.53 10.83 73.36 14.9 15.4 33 63 50.5 15.2 18.2 1 East C1 9 197.2 –8.3 11 202 –12.7 9.9 41.95 170.77 –4.17 5.3 10.6 42.7 163.7 –6.5 5.1 10.1 1 East C2 6 196.2 28.9 17 184.8 24.5 20.2 –23.05 –2.86 15.43 12.8 23.3 –27 8.9 12.9 11.6 21.6 1 East D 11 211.4 88.4 8.5 120.6 60 9.9 47.22 11.68 86.8 17 17 14.7 56.4 40.9 11.3 14.9 1 Center B 12 332.5 85.5 9.8 115.4 58.7 9.8 57.69 6.42 81.05 19.3 19.5 15.8 60.6 39.4 10.8 14.6 1 Center C1 2 200.4 –14.1 15 210.4 –17 43.91 165.45 –7.16 41.6 152.1 –8.7 1 Center C2 4 191.4 24.1 25 181.3 9.7 24.9 –26.62 1.68 12.61 14.2 26.6 –35.1 12.6 4.9 12.7 25.2 1 Center D 14 16.6 82.6 9.9 107.1 56 9.9 63.65 23.45 75.44 18.8 19.3 17.8 67.8 36.6 10.2 14.2 1 West B 4 46.5 79.9 28 104.5 51.1 27.5 60.45 43.71 70.39 50.3 52.6 15.4 72.8 31.8 25.1 37.2 1 West C1 2 190.2 –8.7 72 198.6 –18.4 43.59 –180 –4.38 46.6 166.9 –9.4 1 West C2 1 163.8 21.9 160.1 –4.6 –27.04 32.02 11.36 39.42 –140 –2.3 1 West D 4 338.1 79.7 32 106.2 62.5 32 67.45 –5.27 70.03 58.4 61.1 23.6 63.3 43.8 39 49.9 2 All B 9 15.5 75.7 12 123 59.4 12.3 74.51 41.17 62.99 20.7 22.6 13.1 55.3 40.2 13.8 18.4 2 All C2 6 210.7 42 11 191.1 19.8 11.3 –10.95 –14.17 24.24 8.5 13.8 –29 1.7 10.2 6.2 11.8 2 All D 8 144.8 86.6 16 145 48.6 15.6 44.3 19.59 83.22 30.8 30.9 –4.6 44.2 29.5 13.4 20.5 2 Northeast B 6 15.1 74.7 20 121.3 59.8 19.6 75.89 44.99 61.32 32.4 35.6 14.2 56.1 40.7 22.3 29.5 2 Northeast C2 3 204.5 38.9 8.9 188.9 14.6 8.9 –14.84 –9.31 21.97 6.3 10.6 –32.1 3.7 7.4 4.7 9.1 2 Northeast D 5 92.7 85.2 24 139.2 48.9 23.5 48.61 28.63 80.47 46.1 46.6 –2.4 48.7 29.8 20.5 31 2 Southwest B 3 16.3 77.5 16 126.1 58.5 16.1 71.85 35.55 66.09 28.2 30.1 11 53.8 39.2 17.7 23.9 2 Southwest C2 3 217.7 44.9 28 193.4 25.2 28 –6.54 –19.3 26.49 22.4 35.4 –25.7 –0.4 13.3 16.2 30.2 2 Southwest D 3 200.9 82.5 34 154.2 47.4 33.8 35.98 7.69 75.25 64.3 65.9 –8.2 36.9 28.6 28.5 43.8 Note: Cp—name of the magnetic component measured. D, I—declination and inclination, respectively, before (geo) and after (tilt) bedding correction of the host rock for specimens from both borders (east, west) and for the center of the dike; α95—Fisher statistic parameter; dp, dm—semiaxes of the oval of 95% confidence about the mean pole; N—number of specimens used to compute the mean values; Paleolat—paleolatitude (°N); VGP—virtual geomagnetic pole (lat [°N], long [°E]).
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Structural evolution of the Prague synform (Czech Republic) during Silurian times being characterized by mean inclinations of 11.8° (α95 = 13.3°) for dike1 and 19.8° (α95 = 11.3°) for dike2. Thus, they are very different. Nevertheless, these components (C1 and C2) are associated to the “same” magnetic minerals (Ti-magnetite and/or magnetite) and correspond both to temperatures ranging between 200 and 540 °C (Fig. 7A) or to AC fields ranging between 20 and 65 mT (Fig. 7C). Taking into account of all these observations we can, however, suggest that C1 and C2 are different: C1 is probably primary and possibly overprinted by C2. It has been shown by numerous findings that many of the pre-Variscan rock formations of the Bohemian Massif were partly or totally remagnetized during the Variscan orogeny, most probably during the Carboniferous to the Early Permian (Krs et al., 2001; Edel et al., 2003). Because of the differences in their locations it is likely that the C1 direction corresponds to the direction of the original component, whereas C2 corresponds to either the rotation of the C1 component or to some remagnetization during the Hercynian orogen. The mean difference in orientation between C1 and C2 is significant and large enough (δD = 24.6°, δI = 32°) to explain
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a possible rotation of component C1 with respect to C2. According to conclusion 1 above, the distribution of virtual geomagnetic pole (VGP) fits remarkably well with the apparent polar wander path (APWP) of the Bohemian Massif, and the poles are located very close to the Silurian pole of that massif. According to conclusion 2, the distribution of VGP fits remarkably well with the APWP of the Bohemian Massif, and the poles are located very close to the Carboniferous poles of the massif. The distribution of VGPs after (tilt) bedding correction for dike1 and dike2 are documented in Figure 9. A detailed examination of the data suggests that a small amount of rotation may have occurred preferentially near the eastern edge of the dike. Because the magnetic component C1 is probably of the same age (Wenlock–Ludlow) as the basalt and picritic basalt lava flows, C1 can be considered as associated with the dike emplacement. It is interesting to note that in the center of the dike, where no major disruptions are known, there are only a few C1 directions still preserved. It is important to note that all the components—B, D, C1, and C2—do not show any preferential location in the dikes.
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Figure 9. Virtual pole positions after (tilt) bedding correction for dike 1 (d1) and dike 2 (d2). Names of the component are listed in Table 1. The virtual pole position 36V is of the Barrandian, Karlštejn, Middle Silurian, contact aureole of basalt sill. Apparent polar wandering path, inferred from the East European craton for Early Devonian (D1) to Middle Triassic (T2) time span, is presented by a dashed line.
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RELATIONSHIP BETWEEN MAGNETIZATION AND CARRIERS If we investigate the nature of the carriers we may discriminate between those that carry multiple components and those with single components. As an example of multicomponent carriers, samples SV2/1 and SV6/2 carry three components each: B and D components in common and C1 and C2, respectively, which is in agreement with their coercivity spectra. C1 is mainly carried by MD magnetite whereas C2 is carried by SD magnetite. In a first interpretation, we demagnetized thermally and by AF twelve and twenty-four specimens, respectively, from which four components of magnetization have been separated (Fig. 8, Table 1). Note that when P′ is high, C1 tends to disappear, but this is not a strict rule. As an example, in the center and the west side of dike1, no major shear zone has been observed. In samples with C1 components, lineations are still preserved. Regarding the component of magnetization C2, which is a Late Carboniferous remagnetization, obviously it is nearly missing in the western side of dike1, whereas it is well recorded in the dike’s eastern border, where P′ values are greater than 1.03. Significantly, there is a positive correlation between M and K, which is one reason to suspect that fractures may favor fluid circulation and overprint the primary component C1. The fabric is mainly oblate but can also be prolate in the eastern sides of the dikes (Fig. 6D). This combination is probably related to secondary minerals, as shown in Figure 4. In fact in this figure, which represents the distribution of the magnetic minerals along a cross-section of the dike, the distribution of maghemite or Ti-magnetite is associated with either goethite or hematite. We also notice in the thermomagnetic curves that type I (Fig. 4A, B, D) mentioned by Hrouda et al. (2003) is located mainly in the borders. If we take this criterion into account, we may define the width of each border: for dike1 the eastern border is more fractured and records higher values of P′, and the width may reach 140 cm, whereas the western border is limited to a maximum width of 20 cm. On the section along the x-axis it can be observed that the association of Ti-magnetite and hematite is usually located along the rims of the dike, whereas the samples characterized by Ti-magnetite only or Ti-magnetite and goethite are usually located in the center of the dike. This result suggests various interpretations: • It may indicate that the initial magma was mainly characterized by Ti-magnetite and that some of the primary Ti-magnetites were transformed into magnetites, because the alkaline basaltic tuffs, which constitute the host rock, contain a large amount of titanium. • It may be also associated with a late fluid circulation along the borders of the dike, as already observed by Aïfa and Lefort (2000). • It may at least represent a different magma injection, as suggested in our introduction. However, this interpretation
will not be favored here, because rare C1 components are still observed in the center of the dike. STRUCTURAL IMPLICATIONS AND CONCLUSIONS In a previous interpretation, the evolution of the present Prague synform during Silurian times was characterized by the movement of individual segments along deep synsedimentary faults (Krˇíž, 1998). The sedimentation and the widespread volcanism were considered to be controlled by three main faults (the Prague, Tachlovice, and Koda faults), which delineated three main stripes (the northern, central, and southern segments; Fig. 1). The latter two faults, however, have been interpreted by Melichar (2004) as planes of detachment (i.e., thrust faults separating different thrust units). The original orientation of these faults was not considered as typical of the Paleozoic but was thought to reflect the orientation of some deep Cadomian structures (Havlícˇek, 1963, 1998). This interpretation suggests that the predominant vertical movements recorded along the N65° faults (reaching 1000 m and even 2700 m between the Cambrian and the Lower Devonian; Fig. 1) did not result from a general extension of the lithosphere that controlled Ordovician-Devonian rock units of the Barrandian area but rather from a compressional regime. It is along these faults and along some N10°W faults that the calc-alkaline and sub-alkaline Silurian volcanism was supposed to link (Štorch, 1998). On the contrary, the late shearing episode previously thought to have taken place along these faults, even if limited, now appears to be unlikely, based on the most recent structural data. The general picture that can be given of the Prague synform during Silurian times strongly suggests the existence of a generalized piano-touch tectonics generated in a northeast-southwest compressional regime and followed by a general thrust and nappe tectonics. According to Melichar (2004), regarding the question of vergence in the Prague synform, field evidence agrees with asymmetrical indicators of tectonic movement on fault planes or in a proximal zone of simple shear. If we adopt this way of thinking, we can bring, with our AMS and paleomagnetic data, fresh information on the direction of displacement of the nappes in the Prague synform. This information fits our results of asymmetrical opening of the Svatý Jan pod Skalou dikes. The C1 component corresponds with inclinations known at the end of Silurian times. Component C2 is compatible with paleomagnetic results already known for the end of the Carboniferous. Component D is very similar to a previously published paleolatitude for the Paleocene. Component B, which is close to the D component but older than it according to the published paleolatitude, could be by comparison Cretaceous in age. We mainly concentrate here on C1 and C2 components for our interpretation. This restriction comes from the chronological limitation of the companion articles, which are all devoted to the Paleozoic evolution of the Bohemian Massif. The AMS results obtained on the two dikes of the Svatý Jan pod Skalou area show the existence of an asymmetrical
Structural evolution of the Prague synform (Czech Republic) during Silurian times
the slikensides are witness to a continuity of the stress direction (dextral) or not (sinistral). In any case, because the high P′ value (P′ = 1.06) is located on the eastern side of dike1, we can assume that the eastern disruption really corresponds with a fracture but we cannot say whether the western disruption corresponds with a fracture (P′ = 1.03) or with a cooling crack. According to the literature (Von Raumer et al., 2003; Linnemann et al., 2004; Schulz et al., 2004), the initial opening of the Rheic Ocean would have occurred in a northwest-southeast direction. Jelen´ska et al. (2001) suggested, based on olistoliths coming from the Bardo basin, that the nappes emplacement occurred in the Middle–Late Devonian. This observation suggests that, at that time, the Sudetes had already collided with Baltica and that the Saxo-Thuringian and the Teplá-Barrandian plates were already welded. So far as the subduction of the Rheic Ocean is concerned, our AMS data suggest two solutions (Fig. 10): (1) possible modification of the direction of this azimuth of subduction during the closure of the Rheic Ocean or (2) counterclockwise rotation of the Silurian shortening direction during the collision. Because the transtensional opening of the two dikes remains dextral during Silurian and Carboniferous times, which implies a counterclockwise rotation of the shortening direction and a late-stage sinistral transpressional collision (Fig. 10). If we follow this interpretation, the Prague synform shows, in the Silurian, some affinities with the convergence episode that affected Baltica, Avalonia, and Laurentia. These affinities are not
opening of both dikes. Study of the declination of the ASM K1 components shows that the two dikes were emplaced during a dextral transtensional opening. One of the most important results regarding these dikes is that they both opened by means of the same mechanism but show a difference in their remanent magnetization, as dike2 is devoid of C1 magnetic component. This difference implies that dike2 may be younger than dike1 if the C2 component is primary. Calculation of the difference between the directions of the two different stresses based on the AMS data show that the regional stress suffered a counterclockwise rotation of ~40° between the emplacement of dike1 and that of dike2. This result explains why these dikes display different inclinations. Our data do not provide any evidence on whether the Rheic Ocean existed, but we observe that the counterclockwise rotation of the stress as a function of time was also probably responsible for a modification of the direction of the displacement of the nappes. This counterclockwise rotation of the nappes emplacement strongly suggests that the Rheic Ocean (if its existence is really supported by other data) should have changed its azimuth of subduction between the emplacement of the two dikes or closed following a sinistral shearing. After field observations on dike1, two fracture zones developed in a direction nearly parallel with the border of the dike. Because of the existence of very discrete slikensides showing evidence for horizontal displacement, there is no criterion for the sense of shearing. Consequently, we do not know whether
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2a
Dike2
Figure 10. Reconstitution of the tectonic evolution of the Svatý Jan pod Skalou dikes. The orientations of the AMS lineations (K1) close to the border of the dike are oblique; they show that the emplacement of the magma resulted from a dextral transtensional opening mode. Large open arrow—regional shortening direction; small open arrows—direction of extension; solid line arrows—sense of shearing; α counterclockwise rotation angle (~40°) between the (A) Middle to Late Silurian and (B) the Middle to Late Carboniferous, represented by the shortening directions. (C) Diagram showing the hypothetical evolution of the Rheic Ocean between the Middle to Upper Silurian and the Middle to Upper Carboniferous times. Two solutions are possible: modification of the direction of the azimuth of the subduction plane during the closure of the Rheic Ocean, and counterclockwise rotation of the shortening direction during collision. Middle-Late Silurian shortening direction (1), Middle-Late Carboniferous shortening direction (2), and their respective corresponding nappes emplacement (1a, 2a). B—Baltica; G—Gondwana. If the second solution is valid, it implies a counterclockwise transpression for the closure of the Rheic Ocean.
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consistent with the rifting that has been supposed to affect the Armorican-Bohemian plates at that time (Lewandowski, 1997, 1998, 1999; Marheine et al., 2000; Schätz et al., 2002; Robardet, 2003; Torsvik and Cocks, 2004). However, if we accept the general idea that the Bohemian and Armorican massifs correspond to pieces detached from Gondwanaland and thus were located south of the Rheic Ocean (and not north of it), we must admit that some tightening may have existed between some of these pieces when they were rifting away from Gondwanaland. This suggestion would reconcile the apparent compression we have evidence of, the slow sedimentation that existed during the Silurian in the synform (Krˇ íž, 1998), and the Gondwana faunas that characterize this area. ACKNOWLEDGMENTS We thank the Ministry of Foreign Affairs (Direction des relations et de la coopération internationales) for a two-year grant through programme Barrande 2001–2002 (grant 03229QA), the Centre National de la Recherche Scientifique through GeosciencesRennes (UMR6118), and the Academy of Sciences of the Czech Republic (grant A 3013406). We are deeply indebted to Dr. J.-B. Edel and Professor D. Tarling for the careful comments and suggestions, which helped to improve the final text. We also are grateful to Dr. Linnemann for giving us the opportunity to contribute to this special volume. This article is a contribution to the International Geological Correlation Program Projects 485 and 497. REFERENCES CITED Aïfa, T., and Lefort, J.P., 2000, Fossilisation des contraintes régionales miocènes sous climat aride en bordure de filons doléritiques carbonifères en Bretagne. Apport de l’ASM et du paléomagnétisme: Comptes Rendus de l’Académie des Sciences, Paris, v. 330, no. IIa, p. 15–22. Aïfa, T., and Lefort, J.P., 2001, Relationship between dip and magma flow in the Saint-Malo dolerite dyke swarm (Brittany, France): Tectonophysics, v. 331, no. 1–2, p. 169–180, doi: 10.1016/S0040-1951(00)00241-9. Aïfa, T., Pruner, P., Chadima, M., Lefort, J.P., and Štorch, P., 2002. Mechanism of a Silurian dyke opening in the Prague basin: AMS and rock magnetic evidence: 8th Castle meeting, Paleo, Rock and Environmental Magnetism, 2–7 September, Castle of Zahradky, Czech Republic, p. 9. Blanchard, J.P., Boyer, P., and Gagny, C., 1979, Un nouveau critère de mise en place dans une caisse filonienne: le “pincement” des minéraux aux épontes: Tectonophysics, v. 53, p. 1–25, doi: 10.1016/0040-1951(79)90352-4. Chlupácˇ, I., Havlícˇek, V., Krˇíž, J., Kukal, Z., and Štorch, P., 1998, Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, 183 p. Edel, J.B., Schulmann, K., and Holub, F.V., 2003, Anticlockwise and clockwise rotations of the eastern Variscides accommodated by dextral lithospheric wrenching: Palaeomagnetic and structural evidence: Journal of the Geological Society of London, v. 160, p. 209–218. Havlícˇek, V., 1963, Tektogeneticke porušeni barrandienskeho paleozoika: Sbornik geologickych ved: Geologie, v. 1, p. 77–102. Havlícˇek, V., ed., 1987: The basic geological map of the CˇSSR, scale 1:25,000. Explanatory booklet to map sheet 12–411 Beroun: Prague, Ústrˇední ústav geologický, p. 1–100. Havlícˇek, V., 1998, Prague basin, in Chlupac, I., Havlícˇek, V., Kriz J., Kukal, Z., and Štorch, P., eds., Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, p. 39–78. Hrouda, F., Müller, P., and Hanák, J., 2003, Repeated progressive heating in susceptibility vs. temperature investigation: A new palaeotemperature indicator? Physics and Chemistry of the Earth, v. 28, p. 653–657.
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Printed in the USA
Geological Society of America Special Paper 423 2007
Cadomian and Variscan metamorphic events in the Léon domain (Armorican Massif, France): P-T data and EMP monazite dating Bernhard Schulz* Institut für Mineralogie, Brennhausgasse 14, D-09596 Freiberg, Germany Erwin Krenn Fritz Finger Abteilung für Mineralogie der Universität, Hellbrunner Strasse 34, A-5020 Salzburg, Austria Helene Brätz Reiner Klemd Institut für Mineralogie und Kristallstrukturlehre, Am Hubland, D-97074 Würzburg, Germany
ABSTRACT The Léon domain adjacent to the Cadomian realm in the North Armorican domain appears to be a displaced crustal block, as its metamorphism and rock types bear a resemblance to the South Armorican domain of the internal Variscan belt. The amphibolite-facies Conquet-Penze Micaschist unit overlies the high-grade Lesneven Gneiss unit in the central part of the Léon. Timing and conditions of the metamorphic evolution have been evaluated. At the base of the Lesneven Gneiss unit, a high-pressure eclogite-facies stage (700 °C at >13 kbar) was followed by a high-temperature event (800 °C at 8 kbar), which is characterized by the crystallization of garnet-cordierite assemblages in aluminous paragneisses. Maximal temperatures in the upper parts of the Lesneven Gneiss unit were 630 °C at 6 kbar. Zoned garnet in assemblages with staurolite recorded prograde P-T paths from 490–610 °C at 5–8 kbar in the upper and at 6–9 kbar in the lower parts of the Conquet-Penze Micaschist unit. Garnet Y, heavy rare earth elements, and Li are low in high-grade gneisses and display strong zonations in the micaschists. A younger population of monazite with a broad range of Y contents displays Th-U-Pb ages between 340 and 300 Ma. It crystallized subsequent to formation of foliations S1-S2 and Variscan peak metamorphic assemblages. In contrast, an older population of Cadomian monazite at 552–517 Ma is uniformly rich in Y, suggesting an earlier crystallization than garnet, however, at elevated temperatures. The findings do not support a South Armorican provenance of the Léon domain. The Léon units appear as part of a Cadomian crust at the northern margin of the former Armorican microplate. During a Variscan collision, this crust was strongly overprinted by underthrusting toward the southeast or east beneath
*E-mail:
[email protected]. Schulz, B., Krenn, E., Finger, F., Brätz, H., and Klemd, R., 2007, Cadomian and Variscan metamorphic events in the Léon domain (Armorican Massif, France): P-T data and EMP monazite dating, in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, p. 267–285, doi: 10.1130/2007.2423(12). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Schulz et al. the Central Armorican domain and by later uplift accompanied by Late Carboniferous dextral shear tectonics. The features are typical of the Variscan Saxo-Thuringian zone, which faced the Rheic Ocean to the north. Keywords: Armorican Massif, Léon domain, Armorica, Variscan, Cadomian, P-T paths, aluminous paragneisses, micaschists, metabasites, garnet trace element zonation, monazite Th-U-Pb dating, geothermobarometry
INTRODUCTION The Armorican Massif in western France is assembled out of several crustal domains. To the north, the Neoproterozoic Avalonian-Cadomian orogen was overprinted to a variable degree. To the south, structures and metamorphism of a Paleozoic continental collision are dominant. Within this frame of well-zoned Cadomian and Variscan orogenic belts, the Léon domain to the northwest appears as a strange (“exotic”) unit. Some arguments for displacement of the Léon domain arise from similar rock types, ages, and metamorphic events, as observed in the South Armorican domain, especially the occurrence of eclogites and orthogneisses (Cabanis and Godard, 1987; Le Corre et al., 1989). Tectonic studies revealed dextral shearing in ENE-trending zones along the southern border of the Lèon domain, which were interpreted as major displacement lines (Balé and Brun, 1986). The discussion of whether crustal domains represent displaced and allochthonous terranes is crucial for paleogeographic and plate tectonic models of the Ibero-Armorican segment of the Variscan belt (Franke, 1989; Martinez-Catalan, 1990; Matte, 1991; Dalziel, 1997; Shelley and Bossière, 2000, 2002; Robardet, 2002; Stampfli et al., 2002; Cartier and Faure, 2004). A detailed reconstruction of the magmatic, metamorphic, and structural evolution is essential for this discussion. The present article deals with the P-T evolution, the mineral trace element chemistry, and Th-U-Pb monazite ages from the two major lithotectonic units of the Léon domain. Geothermobarometry on garnet-bearing assemblages in paragneisses and micaschists and on Ca-amphibole in metabasites revealed single prograde-retrograde P-T paths at different temperatures and pressures in the upper and lower parts of the normal crustal pile. The majority of the Th-U-Pb monazite ages confirm that a Barrovian-type metamorphism can be assigned to the Variscan collision. However, an earlier Cadomian thermal event is documented in a distinct population of monazite and provides new details for the zoneography of the Variscan belt in the Armorican Massif. Regional Geological Setting Two major west–east-trending late Variscan shear zones separate the South, Central, and North Armorican domains (Fig. 1). Each of the domains is subdivided into distinct Upper Proterozoic to Lower Paleozoic lithotectonic units. The South Armorican domain is part of the internal Variscan belt (Cogné, 1988; Ballèvre et al., 1994). It involves greenschist-, amphibolite-, and blueschist-facies rocks as well as high-grade and eclogitic units,
with partly complex P-T evolution during eo-Variscan (463– 376 Ma) and Variscan (330–300 Ma) times (Jones and Brown, 1990; Audren and Triboulet, 1993; Ballèvre et al., 1994; Schulz et al., 2001; Lucks et al., 2002). In the Central Armorican domain a Brioverien (Upper Proterozoic) unit with a narrow southern amphibolite-facies zone (Schulz et al., 1998) can be distinguished from an unconformably overlying low-grade to epizonal “classical” Cambrian to Upper Devonian cover sequence (Le Corre et al., 1991; Paris and Robardet, 1994; Rolet, 1994). In the North Armorican domain a unique section across the Cadomian belt is developed (Strachan et al., 1989; Brun and Balé, 1990; Cogné, 1990; Ballèvre et al., 1994; Egal et al., 1996; Brun et al., 2001; Chantraine et al., 2001). The Cadomian Domnonean and Mancellian domains and subunits are separated by northeast-trending major thrusts and shear zones that turn to the northwest in the Trégor province (Fig. 1). Within this framework, the eastern part of the Léon domain is juxtaposed onto the Cadomian realm, whereas its southern part is linked to the Paleozoic of the Central Armorican domain. Geological Setting in the Léon Domain In the central part of the Léon domain two main metamorphic sequences, the amphibolite-facies Conquet-Penze Micaschist unit and the high-grade Lesneven Gneiss unit, can be identified (Rolet et al., 1994). At the southern border and in the hangingwall of these units (Fig. 1), the very low-grade phyllitic Proterozoic schists of L´Elorn have been intruded by a granodiorite, the later Gneiss de Brest, which was dated at 466 ± 25 Ma (Deutsch and Chauris, 1965; Michot and Deutsch, 1970; Cabanis et al., 1977; Le Corre et al., 1991). To the east, Silurian to Devonian schists overlie the micaschists of Conquet-Penze (Fig. 1B). The eastern margin of the Léon domain and the transition to the northwest-trending major Cadomian structures in the Trégor region are masked by the Carboniferous basin of Morlaix (Cabanis et al., 1979a). To the northwest, the fault of Porspoder-Guisseny with a sinistral sense of shear separates the Léon metamorphic pile from the migmatic complex of Landunvez-Plouguerneau (Outin et al., 2000). Two series of granites intruded the metamorphic pile: the older granite complex of Saint Renan-Kersaint with ages of 340–330 Ma and the younger granites of Aber-Ildut-Ploudalmézeau-Kernilis with ages of 300–280 Ma (Cogné and Shelley, 1966; Leutwein et al., 1969; Michot and Deutsch, 1970). Metamorphic structures and the southern margin of the Saint Renan-Kersaint granite were mylonitized by dextral shearing along the North Armorican shear
Léon domain granites ArmoricanCadomian Massifand Variscan metamorphic events in theCarboniferous
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Carboniferous basins (C, M)
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Figure 1. (A) Tectonic subdivision of the Armorican Massif, France, and occurrences of high-pressure rocks (stars). Cadomian units with ages of metamorphism according to Le Corre et al. (1991) and Ballèvre et al. (2001). A—Baie d´Audierne; B—Bois de Cené; C—Champtoceaux; CC—Carboniferous basin of Chateaulin; CAD—Central Armorican domain; E—Essart; G—Ile de Groix; L—Lesneven; M—Carboniferous basin of Morlaix; NAD—North Armorican domain; NASZ—North Armorican shear zone; SAD—South Armorican domain; SASZ-N, SASZS—South Armorican shear zone, northern and southern branches, respectively. (B) Geological map of the Léon region, modified from Cabanis and Godard (1987) and Le Corre et al. (1989, 1991). CAD—Central Armorican domain; ESZ—Elorn shear zone; NAD—North Armorican domain; NASZ—North Armorican shear zone; PGF—Porspoder-Guisseny fault. Variscan granites: AI—Aber-Ildut (290 Ma); B, P—BrigognanPlouescat (290 Ma); K—Kernilis (300 Ma); PL—Plounevez-Lochrist orthogneiss; RK—Saint Renan–Kersaint granite (340 Ma); T—Treglonou orthogneiss; Tr—Trégana granite. Sampling locations discussed in the text are shown in boxes. Sampling locations (Gauß-Krüger coordinate Rechtswert/Hochwert to the west of Greenwich): Penz, 1368900/5355700 (Penzer); PLi, 1368950/5356600 (Plage Porz Liogan); Bil, RenS, 1368850/5357075; Port, 1368450/5357650 (Plage de Portez); PortCo, 1368900/5358150; Pabu, 1368010/5358725 (Porz Pabu, Kermorvan); Sab, 1369600/5360375 (Plage des Blancs Sablons); Kerhorn, 1368900/5362870 (Anse de Porsmoguer); Lan, 1389150/5379300 (Lannilis amphibolites); Kao, 1410400/5385250; Lage, 1411250/534300; Kerz, 1411950/5384900; Trao: 1412450/5385150 (south of Plounevez-Lochrist). (C) Geological cross-section, modified from Le Corre et al. (1989) and Rolet et al. (1994). ESZ—Elorn shear zone; NASZ—North Armorican shear zone; PGF—Porspoder-Guisseny fault.
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zone (Goré and Le Corre, 1987). An offset of ~15 km along the shear zone has been estimated within the 329 ± 9-Ma PlouaretCommana granitic complex (Peucat et al., 1984). The lower part of the central Léon is a high-grade metamorphic sequence and was described as the Lesneven Gneiss unit (Cabanis et al., 1979b; Chantraine et al., 1986). The migmatized ortho-augen-gneisses of Plounevez-Lochrist and Treglonou appear in antiformal structures at its base (Fig. 1C). The orthogneisses provided ages of 400 ± 40 Ma (U-Pb zircon), 392 ± 14 Ma (Pb-Pb zircon; Chauris et al., 1998), and 385 ± 8 Ma (Rb-Sr whole rock), considered as protolith ages by the authors but also interpreted to date a major tectonometamorphic event (Le Corre et al., 1989, 1991). Lenses of eclogites occur within the orthogneisses and overlying paragneisses. The metabasites with a normal mid-oceanic ridge basalt (N-MORB)-type character (Cabanis and Godard, 1987) recorded 650–700 °C at minimum pressures of 13–14 kbar (Paquette et al., 1987; Godard and Mabit, 1998). A 439 ± 13-Ma U-Pb zircon lower intercept age was interpreted to date the highpressure metamorphism (Paquette et al., 1987). In the region of Plounevez-Lochrist, a distinct horizon of garnet-bearing aluminous paragneisses crops out in the vicinity of the eclogites and related amphibolitized eclogites (Fig. 1B). Mineral assemblages with fibrolitic sillimanite and K-feldspar, and stromatic migmatites prevail in paragneisses of the upper part of the Lesneven Gneiss unit. According to the huge anticlinal structure in the central Léon (Fig. 1C), the amphibolites of Lannilis should be part of the Lesneven Gneiss unit. The upper part of the Lesneven Gneiss unit, the Conquet-Penze Micaschist unit, and the Gneiss de Brest are exposed in an almost continuous coastal cross-section in the region of Le Conquet between the Anse de Porsmoguer and the Pte de St. Mathieu. The granodioritic gneiss of Pointe des Renards (565 ± 40 Ma Rb-Sr WR) occurs in lenses both in the micaschist and the gneiss units (Michot and Deutsch, 1970; Chauris and Hallégouët, 1989). Lenses and layers of amphibolites, partly with garnet, a metagabbro, and a meta-porphyroid are intercalated in the Conquet-Penze micaschists (Chauris and Hallégouët, 1989). In the coastal section, the main foliation uniformly strikes ENE. It is dipping steeply in the northern parts and 30–50° to SSE in the southern parts (Fig. 1B). A mineral lineation is dominant in the southern part and plunges 20–30° WSW. The ENE-trending Léon shear zone, where progressive synmetamorphic unroofing of the Léon domain (Jones, 1994) along dextral strike-slip transtensional movement should have been accommodated in Upper Devonian times (Balé and Brun, 1986), is located within the Gneiss de Brest to the south of the micaschist unit. A steeply increasing metamorphic grade toward the north, coinciding with a transition from micaschists to gneisses, was recognized in the coastal section (Jones, 1994). Garnet-bearing assemblages with staurolite prevail in the Conquet-Penze micaschists and display metamorphic conditions of 550–600 °C at 6.3–8.4 kbar (Jones, 1994). In the Kermorvan peninsula, slightly higher temperatures of ~630 °C were estimated from garnet-bearing gneisses (Jones, 1993, 1994). Although the metamorphic conditions in the Lesneven gneiss and Conquet-Penze micaschist units have already been evaluated by
geothermobarometric studies (Paquette et al., 1987; Jones, 1994; Godard and Mabit, 1998), data on metamorphic ages are still sparse. A possible maximum age of the high-pressure metamorphism is provided by the 439 ± 13-Ma U-Pb zircon lower intercept age from the eclogites (Paquette et al., 1987). However, the hightemperature metamorphism could be younger, if the Devonian ages of the orthogneisses are considered as protolith ages. Intrusion of the younger granites at ca. 300 Ma and the development of the North Armorican shear zone give a minimum age limit. It has been demonstrated in various parts of the Variscan belt and other metamorphic terrains (Finger and Helmy, 1998; Williams et al., 1999; Finger et al., 2002; Dahl et al., 2005) that in situ “chemical” Th-U-Pb dating of monazite by analysis with an electron microprobe (EMP monazite dating; Montel et al., 1994, 1996; Suzuki et al., 1994) provides valuable constraints on the timing of metamorphic events and allows the researcher to resolve and distinguish single thermal events in a polymetamorphic evolution. Metamorphic monazite crystallizes as an accessory phase in metapelites and metagraywackes with a limited range of Ca-poor bulk compositions. The presence of garnet with biotite, muscovite, plagioclase, quartz, and aluminosilicates in these rocks can be used to evaluate the metamorphic conditions and P-T paths by geothermobarometry. We combined the EMP monazite dating method with a detailed geothermobarometric study in garnet-bearing micaschists and gneisses. In addition, trace element analyses of garnet were used to evaluate the relative timing of monazite and garnet crystallization in the samples. Geothermobarometry on garnet-bearing metasediments gives insight into a limited part of the P-T evolution of a metamorphic unit. The study has been completed further by geothermobarometric data from metabasites. Monazite Th-U-Pb dating and related geothermobarometry is focused on the garnet-bearing aluminous paragneisses to the south of Plounevez-Lochrist in the lower part of the Lesneven Gneiss unit and the coastal section to the west, with the transition of the Conquet-Penze micaschists into the upper parts of the Lesneven gneisses (Fig. 1B). ANALYTICAL METHODS The main element whole-rock compositions of orthogneisses and of monazite-bearing micaschists and paragneisses were analyzed by X-ray fluorescence spectrometry (XRF). Trace element analyses were performed at Institute of Mineralogy University of Wuerzburg by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) with a single collector quadrupole AGILENT® 7500i ICP-MS equipped with a 266-nm Merchantek® LUV 266x laser. Argon was used as the carrier gas. The laser was adjusted to a scan speed of 5 μm/s at an energy of 0.25 mJ and a repetition rate of 10 Hz. It was traced along 1.6-mm-long and 50-μm-wide extraction lines on LiBO4 fusion glass beads (Brätz and Klemd, 2002). Bulk rock concentrations of SiO2 measured by XRF were used for normalization of LA-ICP-MS analyses. The glass reference material NIST SRM 612 with the values of Pearce et al. (1997) was used for external calibration and
Cadomian and Variscan metamorphic events in the Léon domain calculation of trace elements concentrations by the GLITTER® 3.0 Online Interactive Data Reduction for LA-ICP-MS Program version 2000 by Macquarie Research, Ltd. Mean results from five extraction lines coincide within 1σ error with the data from acid solution ICP-MS (Centre de Recherches Pétrographiques et Géochemiques–Centre National de la Recheche Scientifique, Nancy, France), except for the results for the element La. At measured La concentrations between 16 and 60 ppm, this divergence is not negligible. A constant amount of ~7 ppm La was introduced with LiBO4 into the glass beads and was empirically corrected by regression through comparison of samples with different La concentrations, as described by Sylvester (2003). From the means calculated from five extraction lines, the error is 2.0 wt% occur in monazite in high-grade samples without garnet (Fig. 6C). In the Léon monazites, Y2O3 apparently is not strictly related to metamorphic grade. Therefore, an interpretation of monazite Y2O3 in terms of metamorphic temperatures is difficult. On one hand, increases of Y2O3 in monazite should be correlated with increasing metamorphic grade when xenotime coexists (Heinrich et al., 1997). This trend is not matched by the Variscan monazite, which are lower in Y2O3 in aluminous paragneisses with garnet + sillimanite + cordierite + K-feldspar assemblages when compared to monazite in micaschist with garnet + staurolite assemblages, thereby indicating growth subsequent to the breakdown of xenotime. On the other hand, a low monazite Y content does not automatically indicate a low formation temperature. It may result from unavailability of Y, either because the
Sample Lage Lage Trao Trao Kerz Kerz
Cadomian and Variscan metamorphic Léon domain TABLE 2. CHEMICAL CHARACTERISTICS AND Th-U-Pb MODELevents AGESin OFthe MONAZITE FROM THE LÉON DOMAIN SiO2 Al2O3 CaO Y2O3 La2O3 Ce2O3 Pr2O3 Nd2O3 ThO2 UO2 PbO Total Th U Pb Th* 0.38 29.30 1.03 0.58 13.38 30.97 4.61 11.22 5.55 1.01 0.116 98.14 4.87 0.89 0.11 7.77 0.29 30.10 0.82 0.46 13.41 32.64 4.90 11.76 4.24 0.63 0.086 99.33 3.72 0.55 0.08 5.52 0.27 30.53 1.08 0.49 14.35 31.33 3.97 9.74 5.42 1.49 0.131 98.80 4.77 1.31 0.12 9.03 0.29 30.45 1.10 0.51 14.48 31.05 3.96 9.64 5.65 1.55 0.142 98.81 4.96 1.37 0.13 9.40 0.27 29.68 1.06 0.21 13.32 31.62 4.46 11.93 4.50 0.97 0.101 98.12 3.96 0.86 0.09 6.74 0.50 28.20 1.47 0.34 12.42 29.08 4.13 11.20 7.40 1.29 0.143 96.17 6.50 1.14 0.13 10.19
Age (Ma) 312 ± 23 325 ± 33 304 ± 27 316 ± 25 312 ± 36 292 ± 24
279
Pabu Pabu † Sab † Sab † Brent
0.50 0.14 0.13 0.24 0.25
30.44 29.90 29.72 29.56 29.05
1.26 1.15 1.03 1.00 1.08
1.46 1.33 2.12 2.54 2.53
12.62 12.91 12.67 12.38 12.08
29.00 30.13 29.33 29.16 29.58
3.86 4.04 4.25 4.15 4.16
12.22 11.17 11.37 11.63 11.47
3.39 3.29 3.81 3.33 4.14
2.26 2.41 1.00 1.55 1.06
0.139 0.151 0.090 0.117 0.098
97.14 96.62 95.52 95.64 95.50
2.98 2.89 3.34 2.92 3.64
2.00 2.12 0.88 1.36 0.94
0.13 0.14 0.08 0.11 0.09
9.45 9.79 6.20 7.36 6.68
307 ± 29 322 ± 28 302 ± 39 331 ± 33 307 ± 36
Penz Penz Penz PLiN1 PLiN1 Portez Portez
0.68 2.25 1.76 0.29 0.31 0.17 0.48
29.54 26.39 26.59 27.72 28.80 29.90 30.17
0.60 1.60 1.56 0.89 1.12 0.66 0.96
1.67 1.82 2.23 1.49 1.86 1.50 1.69
13.10 10.53 10.53 11.94 11.58 13.43 11.76
31.08 24.47 24.90 29.32 28.25 32.62 28.71
4.31 3.30 3.42 4.72 4.50 4.19 3.74
12.38 2.71 10.08 16.73 10.32 14.07 12.72 4.58 11.96 5.68 12.32 2.70 11.35 4.13
0.35 0.43 0.52 0.89 1.10 0.56 0.86
0.048 0.228 0.203 0.115 0.133 0.063 0.087
96.46 2.38 97.84 14.70 96.09 12.37 94.68 4.02 95.29 4.99 98.11 2.37 95.36 3.63
0.31 0.38 0.45 0.78 0.97 0.50 0.76
0.04 0.21 0.19 0.11 0.12 0.06 0.08
3.38 15.93 13.84 6.58 8.15 3.98 6.09
296 ± 80 297 ± 17 305 ± 19 363 ± 27 340 ± 22 328 ± 68 298 ± 44
Port1 0.22 29.99 1.14 1.22 10.82 30.05 4.84 12.63 5.91 0.31 0.167 97.30 5.20 0.27 0.16 6.09 566 ± 29 Port1 0.08 30.25 1.35 2.29 11.53 28.75 4.42 11.30 4.01 2.66 0.274 96.92 3.52 2.35 0.25 11.26 507 ± 16 Pabu 0.09 30.02 1.14 2.49 11.44 28.82 4.48 11.60 4.46 1.22 0.200 95.96 3.92 1.08 0.19 7.49 555 ± 24 Pabu 0.12 29.25 1.21 2.64 11.70 28.80 4.13 12.27 4.92 1.07 0.186 96.30 4.33 0.94 0.17 7.43 519 ± 36 † 0.17 29.13 1.03 3.18 11.49 28.78 4.01 12.00 4.38 0.56 0.141 94.87 3.85 0.50 0.13 5.49 533 ± 43 Sab † 0.16 29.22 1.09 1.55 11.61 29.39 4.10 12.39 5.01 0.42 0.152 95.09 4.40 0.37 0.14 5.63 558 ± 42 Sab Note: Model ages shown with 95% confidence level. Model ages and errors have been calculated from multiple analyses of one point. Locations of samples are shown in Figure 1. † Sample contains no garnet.
0.06
A
Ca
70
REE (wt.%)
B
0.04
Variscan (Al-rich) Variscan Variscan, no garnet Cadomian
0.02
60
Y2O3
Th+U
0
4 3
0.02
Y2O3
0
0.04
0.06
0.08
no Grt
0.10
0
1
2
3
4
C Figure 6. Mineral chemistry of monazite. (A) Linear positive correlation of U + Th and Ca of brabantite substitution in monazite. (B) Variation of Y2O3 and REE in monazite from various parent rocks and age populations. (C) Variation of monazite Y2O3 with Th-U-Pb model age. Variscan monazite from samples without garnet (Grt) have been treated as a separate group (see Table 2).
2 1
0 200
50
age (Ma) 300
400
500
600
700
280
Schulz et al.
500
0.3
400
Pb
0.2
0 4
2
6
8
Th* 400
Pb
0.2
0 2
4
6
8
552 310
22 Ma 14 Ma
Th*
500
0.3
400
0.2
Portez 304 21 Ma Port1 517 13 Ma
0 2
0
I
8
8
10 12 14 16 18
4
6
8
Th*
G Th* 0
2
4
6
8
10 12 14 16 18
Conquet-Penze Unit Lesneven Gneiss Unit
Cadomian
Variscan 30 0
H
0
10 12 14 16 18
1000
20 km
N Brent 308
16 Ma 22 Ma 14 Ma
520 Pabu 324
18 Ma 15 Ma
Port 517 304
13 Ma 21 Ma
Bil 306
18 Ma
PLi 340
16 Ma
Penz 305
10
K
T Treglonou AI
Lesneven
St. Renan
RK
U/Pb 15
20
Granites 290 Ma and 330 Ma
PL
Trao 306 12 Ma Kerz 300 10 Ma Lage 309 11 Ma
Penze
Migmatites Palaeozoic of NAD and CAD Gneiss de Brest Schistes d´Elorn
Landiviseau Plabennec
Ma
Ma
Léon Domain
P
Pl.-Loc.
Sab 552 310
Ma
5
0
B
F PG
500
Roscoff
EMP monazite ages
340 300 270 Ma
0
10 12 14 16 18
400
Lage 309 11 Ma Trao 306 12 Ma Kerz 300 10 Ma
Pb
Léon monazites
40
6
20
C
0.1
60 340 300 270 Ma
Pb
Th/Pb
80
6
4
0.2
Th* 4
2
500
0.1
2
0
Th*
0.3
340 300 270 Ma
0
10 12 14 16 18
F 0
400
Bil 306 18 Ma Brent 308 16 Ma
340 300 270 Ma
Penz 305 21 Ma
0
10 12 14 16 18
E
0.1
Sab Sab 0
Pb
0.2
B
0.1
8
500
0.3
340 300 270 Ma
6
400
0.2
Th* 4
2
0
Pb
0.1
0
10 12 14 16 18
500
0.3
PLiN1 340 16 Ma
500
0.3
340 300 270 Ma
D
0.1 Pabu 520 18 Ma Pabu 324 15 Ma
0
Pb
400
0.2
A
0.1
500
0.3
340 300 270 Ma
NASZ
Conquet-Penze Micaschist Unit Micaschists Amphibolites
Le Conquet
Tr
Lesneven Gneiss Unit Paragneisses and migmatites Brest o El
rn
Z
ES
Aluminous paragneisses Eclogites and amphibolites Orthogneisses of Treglonou
21 Ma
Figure 7. Th-U-Pb model ages in monazite. (A–G) Total Pb vs. Th* (wt%) isochron diagrams for various samples. Isochron ages derived from these diagrams broadly match weighted average ages calculated for the rocks according to Montel et al. (1996). (H) Distribution of Variscan and Cadomian Th-U-Pb monazite ages in the U/Pb-Th/Pb diagram of Cocherie and Albarede (2001). (I) Presentation of monazite ages in simplified map. See Figure 1 caption for abbreviations.
Cadomian and Variscan metamorphic events in the Léon domain host rocks are Y-deficient, or because Y is retained and fractionated into other minerals, such as xenotime and prevailing garnet (Pyle et al., 2001). Monazite is low in Y in samples with abundant low-Y garnet, as in aluminous paragneisses, and it is intermediate in samples with few garnets. High Y in monazite occurs in samples without garnet. These observations allow us to conclude that Variscan monazites should have crystallized subsequent to the breakdown of xenotime and during or subsequent to Mg-rich garnet, which represents the thermal peak of the metamorphic evolution. Furthermore, the potential for preservation of older populations of monazite increases with bulk rock Y and when the mode of garnet remains low and xenotime is absent. All ten studied samples contain monazite with Variscan and older chemical model ages (Fig. 7A–G, Table 2). The weighted averages (Ludwig, 2001) of the Variscan ages range from 304 ± 21 Ma (Port) to 340 ± 16 Ma (PLiN1) in the Conquet-Penze micaschists. In the Lesneven gneisses, the ages are younger and range from 300 ± 10 Ma in Kerz to 324 ± 15 Ma in Pabu (Fig. 7A– G). Three samples with Variscan monazites contain another monazite generation, which yielded ages at 520 ± 28 Ma (Pabu), 517 ± 13 Ma (Port) and 552 ± 22 Ma (Sab1). The “Cadomian” monazites occur both in the Lesneven Gneiss and in the Conquet-Penze Micaschist units. The single ages from individual grains display two distinct groups between 280–350 Ma and 560–510 Ma in the U/Pb-Th/Pb diagram (Fig. 7H) of Cocherie and Albarede (2001). No monazite was found with ages in between or linking these two groups. Thus, the monazites belong to entirely separate events. The existence of clearly separated groups of ages, observed within single samples, indicates that monazite has not been affected by possible lead loss or postcrystalline metasomatism. CADOMIAN AND VARISCAN METAMORPHIC EVENTS Thermobarometric data from metasediments and metabasites in the Lesneven Gneiss and the Conquet-Penze Micaschist units signalize a normal crustal pile and increasing maximal metamorphic temperatures and pressures with increasing structural depth. A change from prograde-zoned garnet with increasing XMg toward the rims to Mg-rich garnet with homogeneous cores and zoned outermost rims coincides with the mapped lithological border of the Conquet-Penze Micaschists and the Lesneven Gneisses. The common retrograde P-T evolution in the sillimanite and then andalusite stability fields was mainly recorded by assemblages with Ca-amphibole in the various metabasites. Mineral-chemical, thermobarometric, and Th-U-Pb monazite age data indicate a transitional character and no major metamorphic discontinuity between the Conquet-Penze and Lesneven units. Monazite Th-U-Pb model ages determined by the EMP range from 340 to 300 Ma, with younger monazite of 310– 300 Ma prevailing in lower parts of the Lesneven Gneiss unit (Fig. 7I). In aluminous paragneisses, monazite of 310–300 Ma coexisted with cordierite and serves as a temporal marker for the high-temperature metamorphic stage. The eclogitic high-pressure stage is linked to the high-temperature stage along a single
281
P-T path recorded by the prograde garnet zonation in the eclogites (Fig. 5A). This link provides an argument that the age of the high-pressure stage may be younger than the 439 ± 13-Ma U-Pb zircon lower intercept age given by Paquette et al. (1987). However, the Carboniferous monazite ages coincide with the KAr and Rb-Sr mica ages in the granites and orthogneisses and the intrusion of the late granites (l´Aber-Ildut, Ploudalmezeau) at 300–290 Ma. They postdate the early 340–330-Ma granite de Saint Renan-Kersaint, which was already post-tectonic and posterior to an Upper Devonian “phase bretonne” observed in low-grade Paleozoic sequences (Le Corre et al., 1989, 1991). In consequence, the metamorphic monazites support a distinct “late Variscan thermotectonic event” (Le Corre et al., 1989) in the Léon domain that lasted until the Lower Permian during the late generation of granite intrusion. Similar Late Variscan mica ages have been reported from the South Armorican domain (Brown and Dallmeyer, 1996). The question of whether the amphibolite-facies metamorphism in the Conquet-Penze unit is a Variscan or an older event (Jones, 1993, 1994) can now be discussed in view of new Cadomian monazite ages, the garnet trace element zonations, and the P-T conditions of metamorphism. Cadomian monazites were analyzed in three samples from different locations in both the Conquet-Penze and Lesneven units. Variscan monazite appears in the same samples. Cadomian monazite occurs in a garnet-free paragneiss and in samples with low modes of garnet. All monazites in garnet-rich micaschists and aluminous paragneisses are of Variscan age. Cadomian monazite is rich in Y, regardless of whether it is from samples with Y-rich or Y-poor garnet or without garnet, and contrasting low-Y Variscan monazite observed in the same samples. From their uniform mineral-chemical compositions and their narrow range of ages, a detrital origin of the Cadomian monazites can be excluded. They should record an early thermal event in the units and provide a minimum age of sedimentation. As Y-rich Cadomian and Y-poor Variscan monazite appear in the same sample with garnet, it can be concluded that Cadomian monazite crystallized previous to, and Variscan monazite subsequent to, the garnet; the garnet-bearing assemblages are bracketed by the two populations of monazite. New monazite can crystallize when Y is available through breakdown and retrogressive replacement of nearby garnet and xenotime (Foster et al., 2000, 2002; Pyle and Spear, 2000). This interpretation is supported by the observation of Y-rich Variscan monazites in samples without garnet. The existence of monazite previous to garnet crystallization is further supported by Y-poor garnet cores, which imply the former presence of another phase (monazite and/or xenotime), which fractionated Y previous to garnet growth. The Y content of a preserved early monazite should be dependent on bulk rock composition, metamorphic temperature, and the presence of xenotime, regardless of whether later garnet growth occurred. This dependence allows some conclusions to be drawn on the nature of the Cadomian thermal event. At increasing temperature, Y in monazite increases at the expense of xenotime, which is consumed
282
Schulz et al.
and disappears at higher grades (Heinrich et al., 1997; Pyle et al., 2001; Pyle and Spear, 2003). Xenotime was not evident in the studied samples and presumably was consumed; thus, highY Cadomian monazite could indicate an elevated temperature of metamorphism. The bulk rock composition was suitable for garnet crystallization in some samples, however, the observed garnet has low Ca and crystallized at medium pressures. Garnet would not be stable at the given Ca-poor bulk rock compositions when pressure is low (Spear, 1993; Pyle and Spear, 2003). Monazite and xenotime are consumed during garnet growth (Pyle et al., 2001; Pyle and Spear, 2003). Thus, the preservation of Cadomian monazite with high-Y content could be explained by the lack of coeval garnet, caused by low pressures during the Cadomian metamorphic event. One could speculate on a Cadomian contact metamorphism in the vicinity of intrusions like the Pointe des Renards metagranitoid. From the available sparse radiochronological data (565 ± 40 Ma, Rb-Sr WR; Michot and Deutsch, 1970) this possibility cannot be excluded. A Cadomian regional low-pressure metamorphism was described from the eastern parts of the North Armorican Cadomian domain (Ballèvre et al., 2001) and appears to be an alternative explanation for the Cadomian monazite ages. There exists a zoneography of metamorphic cooling ages within the Armorican Cadomian belt, ranging from Neoproterozoic (600 Ma) cooling in the Trégor unit to the northwest to Cambrian (520 Ma) cooling in the Fougères unit to the southeast (Fig. 1A). The northeastern part of the Léon is juxtaposed on the Trégor unit of the internal Cadomian belt along the main Cadomian thrust, which curves to the northwest along the Baie de Morlaix. To the east, the Léon continues into the Saint Brieuc and Saint Malo units of the external Cadomides (Fig. 1A and B). In the St. Brieuc unit and its subordinate Yffinac formation, maximal high-grade conditions at 9 ± 1 kbar and 700 ± 50 °C (Hébert, 1995) were achieved previous to cooling at ca. 570–560 Ma (Dallmeyer et al., 1993). This scenario is in contrast with the later Cadomian deformation and metamorphism in the St. Malo migmatitic unit at ca. 550–540 Ma, which is characterized by low pressures and high temperatures (45-cm-thick K-bentonite (altered ash-fall tuff) bed within the upper Barrios Formation (Ordovician Armorican Quartzite facies), in the Cantabrian zone of the Iberian Variscan belt were dated by isotope dilution–thermal ionization mass spectrometry. U-Pb analyses of six highly abraded single grains yielded concordant and overlapping error ellipses with a pooled concordia age of 477.47 ± 0.93 Ma. This age provides the depositional age of the Armorican Quartzite facies in the studied sector and establishes an absolute minimum age for the rifting that led to the opening of the Rheic Ocean in this section of northern Gondwana. This age is within error of the currently accepted interpolated age for the Tremadocian–Lower Ordovician Stage 2 (Floian) limit at 478.6 ± 1.7 Ma (Gradstein et al., 2004). Keywords: Iberian Peninsula, Ordovician, Armorican Quartzite facies, K-bentonite, zircon, U-Pb dating *E-mail:
[email protected]. Gutiérrez-Alonso, G., Fernández-Suárez, J., Gutiérrez-Marco, J.C., Corfu, F., Murphy, J.B., and Suárez, M., 2007, U-Pb depositional age for the upper Barrios Formation (Armorican Quartzite facies) in the Cantabrian zone of Iberia: Implications for stratigraphic correlation and paleogeography, in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, p. 287–296, doi: 10.1130/2007.2423(13). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION High-precision U-Pb dating by isotope dilution–thermal ionization mass spectrometry (ID-TIMS) of single zircon crystals is currently considered to be the most robust method to obtain precise ages of air-fall volcanic tuffs (often transformed into Kbentonites). The fundamental caveat is the presence or absence of a zircon population that was generated as a result of the volcanic event (indigenous zircon), as this population can be missing or very dilute in older detrital (remanié) zircons (see Bowring and Schmitz, 2003, for a review of details and analytical and natural complexities). Under any circumstances, precise U-Pb dating of air-fall tuffs and/or K-bentonites provides the most accurate absolute age constraints of depositional ages in those sedimentary formations in which they are interbedded, because they are considered to be instantaneous basin-scale deposits. Moreover, in fortunate cases where these rocks occur at boundaries of stratigraphic units, they are the main tool to constrain the absolute time frame of different subdivisions of the stratigraphic record (e.g., the Ediacaran-Cambrian boundary: Brasier et al., 1994; Amthor et al., 2003; Bowring and Schmitz, 2003; Condon et al., 2005; the base of the Tremadocian: Landing et al., 2000; or the compilations made by Sadler and Cooper, 2004) and greatly aid to fine-tune correlations based only on fossil occurrences (Cooper and Sadler, 2004). Precise age constraints on the depositional age of key formations representing main changes in basin development are fundamental in paleogeographic and paleotectonic studies, as they provide the time frame for development of back-arc basins, overstep sequences, rifted margins, platform development, rift-drift transitions, and the like. Given the diachronism of depositional events over large areas, the “one datum per rock unit” approach is not advisable, precisely because it is only a reliable and precise time frame that can provide powerful constraints on the evolution of sedimentary basins in relation to tectonics. Within this frame, Lower Ordovician sediments in Perigondwanan realms record a complex environmental and tectonic story (Avigad et al., 2003, 2005). From this point of view, the Ordovician period is characterized by widespread epicratonic seas (Ross and Ross, 1995) and abundant passive margin sedimentary successions. The Armorican Quartizite is an important stratigraphic facies of clastic deposits that extended along the Perigondwanan margin from west Africa through Iberia, Armorica, and continental Europe, probably as far east as Serbia, Saudi Arabia, or even Afghanistan. These clastic deposits contain similar trace fossils (abundant Cruziana of the rugosa group and vertical assemblages of Skolithos and Daedalus), linguliform brachiopods (giant obolids), and rare bivalves and trilobites that are rather diagnostic (but not exclusive) of West Gondwana (Romano, 1991; Seilacher, 1992; Cocks, 2000; Fortey and Cocks, 2003). Other localities with the same fossil assemblages as the Armorican Quartzite facies are also present as pebbles in a Triassic conglomerate located in the southern British Isles (Cocks and Lockley, 1981; Cocks, 1993) eroded during the early stages of the Variscan orogeny.
From a different point of view, the Gondwanan environment was controlled by the south pole position, which lay within or adjacent to the west African portion of West Gondwana, which spanned more than 100 degrees of latitude (Cocks, 2001, and references therein). Therefore, a strong temperature gradient from pole to equator throughout the Ordovician (Spjeldnaes, 1961) resulted in faunal endemism that, when combined with paleomagnetic evidence, constrains rift and drift of Avalonia and Carolina away from the Gondwanan margin (Cocks and Torsvik, 2002; Fortey and Cocks, 2003, and references therein). Furthermore, basin modeling of Avalonian strata (Prigmore et al., 1997), the occurrence of thick Late Cambrian–Early Ordovician platformal sedimentary (mostly siliciclastic) successions in Iberia and Armorica that contain Armorican Quartzite facies, and widespread magmatism linked to extension all suggest rifting of Avalonia had commenced by that time (Quesada, 1991; Simancas et al., 2003; Salman, 2004; Murphy et al., 2006). As Tremadocian to early Darriwilian fauna of Avalonia contains trilobites that are similar (at the species level) to those of West Gondwana, the Avalonian microcontinent probably remained close to West Gondwana at that time (Fortey and Cocks, 2003). Thus, deposition of the Armorican Quartzite facies occurred in the earlier stages of the rift to drift process of Avalonia and development of the Rheic Ocean. It is marked by several unconformities of lower and upper Arenigian age (Gutiérrez-Marco et al., 2002) and in-between unconformities, where sedimentation is more extensive and uniform in western Europe. By 460 Ma, however, paleomagnetic evidence suggests that Avalonia had migrated ~20° north of the Gondwanan margin (Van der Voo, 1988; Hamilton and Murphy, 2004); by the late Ordovician, Avalonian fauna reflects proximity with Baltica; and by the mid-Silurian, proximity with Laurentia (see Bassett and Cocks, 1974; Harper and Owen, 1984). Taking into account all aforementioned factors involving the origin of the Armorican Quartzite facies, which is composed of up to 99% detrital quartz, its deposition and provenance are not yet fully understood (Avigad et al., 2005). According to Linnemann and Romer (2002), the abundance of quartzite reflects increased reworking of older sediments and weathering of cratons during a hiatus in sedimentation. This interpretation is supported by lower abundances in elements sensitive to weathering in the quartzite (which reflects a dilution effect by quartz), and by relative enrichment of some trace-elements (Zr, middle rare earth elements [MREE], and heavy rare earth elements [HREE]), reflecting concentration of stable accessory minerals, such as detrital zircon and garnet. Sm-Nd TDM model ages of 1.9–1.5 Ga support detrital zircon-age data which indicate affinities with the West African craton (Fernández-Suárez et al., 2002). The importance of the Armorican Quartzite facies from a paleogeographic perspective resides in the widely accepted premise that deposition of this regionally extensive formation postdates the break-up unconformity (widely, but erroneously, known as the “Sardic unconformity” of Upper Ordovician age; Aramburu and García-Ramos, 1988; Brun et al., 1991) that
U-Pb depositional age for the upper Barrios Formation (Armorican Quartzite facies) equates to the “Toledanian” unconformity (Gutiérrez-Marco et al., 2002) and was deposited after the Cambrian–Early Ordovician magmatic activity in northwest Iberia (“Ollo de Sapo” belt and related rocks; Valverde-Vaquero and Dunning, 2000) and in Armorica (Le Corre, 1994). The Toledanian unconformity marks the rift-drift transition related to the origin of the Rheic Ocean (e.g., Gutiérrez-Alonso et al., 2005). In spite of its importance, no precise absolute-age constraints on the deposition of the Armorican Quartzite facies exist except for some low-precision U-Pb zircon data from western Armorican Massif (Bonjour et al., 1988; Bonjour and Odin, 1989). This article presents a precise date for a K-bentonite layer within an equivalent to the Armorican Quartzite facies in the Cantabrian zone of the Iberian Variscan belt (the Tanes Member of the Barrios Formation), which lies at the Tremadocian-Floian boundary (global Ordovician chronostratigraphy updated by the International Commission of Stratigraphy; Bergström et al., 2004, 2006). This age constrains the detailed timing for the break-up of Avalonia from northern Gondwana and pinpoints the absolute age of part of the well-known lower Palaeozoic detrital sequence of the western European Variscan belt. GEOLOGICAL BACKGROUND The Armorican Quartzite facies is widely exposed in western Iberia, providing an excellent marker to unravel the structure of the Variscan belt in this region, and it has been intensively studied from the stratigraphic and sedimentological points of view (Aramburu, 1989; Aramburu and García-Ramos, 1993). These studies focused on the Armorican Quartzite facies in the core of the curved west European Variscan belt, in the region known as the Cantabrian zone, where it is locally equivalent to the Tanes Member of the Barrios Formation (see below). This unit, first studied by Barrois (1882) and Comte (1937), was reviewed in detail by Aramburu (1989) and Aramburu and García-Ramos (1993), who divided the Barrios Formation into three members, of which the lowermost one (La Matosa Member, quartzites and shales) is of Cambrian age, the middle member (Ligüeria Member, conglomerates and siltstones) is of undetermined Tremadocian age, and the uppermost member (Tanes Member, mainly quartzites) is the alleged equivalent to the Armorican Quartzite facies in the rest of Iberia and from which the studied sample was collected. The Barrios Formation in the Cantabrian zone has not metamorphosed or recrystallized since its deposition. A good section of the La Matosa and Tanes members is exposed in the core of the Viyazón-Reigada syncline within the Somiedo unit, where a reference stratigraphic section (Embalse de La Barca in Aramburu, 1989; Aramburu and García-Ramos, 1993) was established (Figs. 1–3). The Barrios Formation is locally represented by the Tanes Member, which is composed of extremely pure, white quartzites interpreted to have been deposited in a braid-plain delta environment with marine influence. Paleocurrent data and facies distribution indicate that transport occurred from east to west (i.e., from the core of the Ibero-Armorican arc;
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Fig. 1) toward a deeper marine environment. Detailed sedimentological data and stratigraphic correlations are given in Aramburu and García-Ramos (1993). The lower boundary of the Tanes Member is an erosional paraconformity above the Upper Cambrian Oville Formation (Fig. 2), which is interpreted to represent the aforementioned Toledanian unconformity (Aramburu et al., 2004; previously claimed as “Sardic” by Aramburu and García Ramos, 1988), classically believed to be located at the Tremadocian-Arenigian boundary. A well-known feature of the Lower Paleozoic succession in northwest Iberia is the abundance of long-lived magmatism, which is represented in the Cantabrian zone by alkaline basalts and volcaniclastic rocks located mostly in Upper Cambrian and Lower Ordovician strata (Loeschke and Zeidler, 1982; Heinz et al., 1985; Gallastegui et al., 1992, 2004; Suárez et al., 1993; Barrero and Corretgé, 2002) together with two K-bentonite (according to the classification of Fischer and Schmincke, 1984) beds interstratified within the Barrios Formation (García-Ramos et al., 1984). These K-bentonite layers are locally known as the Valverdín bed and the more extensive and younger Pedroso bed, which does not spatially overlap the former. The Pedroso bed, the object of this study, extends over more than 1800 km2 with a thickness between 45 and 80 cm (Aramburu, 1989) and is located ~220 m above the base of the Tanes Member of the Barrios Formation, which in this section is anomalously thick for the Cantabrian zone, reaching a thickness of ~750 m. Recently, a new Kbentonite bed within the Barrios Formation was reported in the northernmost part of the Central Coal basin of the Cantabrian zone. This bed is interpreted to be an eastward extension of the Pedroso bed (Gutiérrez-Marco and Bernárdez, 2003; GutiérrezMarco et al., 2003). The K-bentonite Pedroso bed is always interstratified within Skolithos pipe beds (Fig. 2), indicating very low sedimentation rates, and is interpreted as altered ash-fall tuffs (García-Ramos et al., 1984; Aramburu, 1989). The upper and lower contacts with the bioturbated quartzites are very sharp, and the massive ash-fall apparently did not affect the population structure and the development of the benthic communities, which enhanced a rapid recovery and recolonization of the shallow marine environment in a way similar to that observed in other Ordovician and modern ash-falls (Huff et al., 1992, Kuhnt et al., 2005). According to previous descriptions, two different lithologies form the Kbentonite beds: (1) coarse-grained (type G) in thin layers located near the base of the Pedroso bed, showing graded and cross-bedding structures and (2) fine-grained (type F), consisting basically of the same lithology with horizontal lamination marked by pyrite (Fig. 2). Another widespread correlation event within the Armorican Quartzite facies of southwest Europe is a tsunami deposit recorded as a ubiquitous linguloid bed (Emig and Gutiérrez-Marco, 1997). In addition to the date presented here, a former study of detritral zircon U-Pb ages in quartzite beds a few meters above the studied K-bentonite Pedroso bed within this formation (Fernández-Suárez et al., 2002) revealed the presence of age clusters of
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Somiedo Nappe
B
C 0
500 km
CANTABRIAN ZONE
0
30 km Precambrian Pedroso and Fabar Beds
Unconformable Meso-Tertiary
IBERO-ARMORICAN ARC (Shaded area contains main exposures of the Armorican Quartzite facies)
Valverdín Bed
Figure 1. (A) Simplified geological map of part of the Somiedo unit where the Villazón-Reigada syncline is exposed, with the location of the studied sample. Also shown is the location of the sample used by Fernández-Suárez et al. (2002) for detrital zircon U-Pb dating (see text for details). Fm.—formation; Lst.—limestone. (B) Sketch of the Cantabrian zone, with inset for location of panel A. Patterns depict the extension of the Ordovician K-bentonites (Aramburu and García-Ramos, 1993; Gutiérrez-Marco and Bernárdez, 2003; Gutiérrez-Marco et al., 2003); arrows represent a summary of the paleocurrent directions in the Barrios Formation according to data from Aramburu and García-Ramos (1993). (C) Ibero-Armorican arc in western Europe.
ORDOVICIAN
5th & 6th Stages
Darriwilian
2nd & 3rd Stages
Luarca Fm.
Sample for detrital zircons from Fernández-Suárez et al. (2002)
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C
B Barrios Fm. (Tanes Mb.)
Tremadocian Furongian
A
U-Pb depositional age for the upper Barrios Formation (Armorican Quartzite facies) Castro Fm.
PEDROSO K-BENTONITE
Oville Fm.
*
Middle
CAMBRIAN
Láncara Fm.
0 cm
2m Lower
20 cm
0m
Herrería Fm. 1000 m
500 m
Skolithos EDIACARAN
Narcea Shales
Quartzites 0m
G Type F Type Pyrite Framboids
*
Studied sample
Figure 2. (A) Synthetic stratigraphic column of the Lower Paleozoic in the Somiedo-Cabo Peñas region of the Cantabrian zone (Aramburu et al., 2004) showing the stratigraphic location of the K-bentonite sample used for U-Pb dating and the stratigraphic location of the sample used by Fernández-Suárez et al. (2002) for detrital zircon U-Pb dating (see text for details). Fm.—formation; Mb.—member. (B) Detailed stratigraphic section of the Pedroso K-bentonite and the adjacent Skolithos-rich beds (after Aramburu, 1989, Fig. 41.18). (C) Detailed section of the Pedroso K-bentonite showing the distribution of G- and F-types in the surroundings of the sampled section (adapted from Aramburu, 1989, Fig. 43, and García-Ramos et al., 1984).
Figure 3. Picture of the Villazón-Reigada syncline, with location of the collected sample and the sample for detrital zircons in Fernández-Suárez et al. (2002).
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ca. 800–550, 1300–900, 2200–1800, and 2800–2500 Ma, and the youngest detrital zircon found in that study was 550 Ma. The studied sample in this work was collected at Mina Conchita, an underground mine owned by Caolines de Merillés that is in the core of the Villazón-Reigada syncline (N43°19′25.7″, W6°18′0.07″; Figs. 1 and 3). The sample consisted of 25 kg of white G- and F- type K-bentonite. PALEONTOLOGICAL DATING OF THE ARMORICAN QUARTZITE FACIES IN SOUTHWEST EUROPE Traditional time correlation of the Armorican Quartzite facies within southwest Europe placed this formation approximately in the Arenig, with its upper boundary being roughly equivalent to the Arenig-Llanvirn boundary. Moreover, over large areas of the Central Iberian zone and the Armorican Massif, volcanosedimentary units and coarse red beds, which lie above the Toledanian unconformity and directly below the typically massive whitish orthoquartzites, were tentatively dated as Tremadocian, owing to their stratigraphic position below the “genuine” Armorican Quartzite (Arenigian). The paleontological record of the Armorican Quartzite facies is mainly characterized by abundant ichnofossils of the Cruziana and Skolithos ichnofacies, both being representative of diverse settings in a range of wave-dominated to tide-dominated shallow-marine environments. The most typical ichnoassemblages are represented by post-Tremadocian forms of Cruziana (rugosa and imbricata groups) and by huge concentrations of vertical burrows, such as those of Skolithos and Daedalus (e.g., Seilacher, 1970; Crimes and Marcos, 1976; Baldwin, 1977; Kolb and Wolf, 1979; Pickerill et al., 1984; Durand, 1985; Romano, 1991, and references therein). The Armorican Quartzite facies also yielded a widespread suite of giant linguliform brachiopods (Lingulobolus, Ectenoglossa, Lingulepis, Tomasina, etc.), as well as some molluscs (essentially bivalves, rare cephalopods, gastropods, and rostroconchs), conularids, and a few trilobites and other arthropods, such phyllocarid crustaceans and xiphosurans (e.g., Rouault, 1850; Davidson, 1880; Guillier, 1881; Barrois, 1891; Babin, 1966; Henry, 1980; Gutiérrez-Marco et al., 1997; Coke and Gutiérrez-Marco, 2001; Babin and Hammann, 2001). Despite the paleoecological importance of the giant linguliform brachiopods, they constitute a biofacies that unfortunately lacks biostratigraphical potential. The only biostratigraphical resource coming from the Armorican Quartzite facies in Iberia is provided by a single graptolite occurence of middle Arenigian age (Gutiérrez-Marco and Rodríguez, 1987; Gutiérrez-Marco and Bernárdez, 2003). The aforementioned data are in agreement with the oldest graptolite assemblages found in shaley units immediately overlying the Armorican Quartzite, of middle to upper Arenigian age (Paris, 1990, and references therein). In addition, the micropaleontological record from the Armorican quartzites contains diverse chitinozoans, acritarchs, and leiospherids. Detailed study of chitinozoans showed that the whole formation belongs to the
Eremochitina brevis biozone (Paris, 1981, 1990; Paris et al., 1982), which is equivalent to a late early to middle Arenigian age and roughly equivalent to the Floian of the Ordovician System (= Time Slice 2c-basal 3a of Webby et al., 2004) in several places around Gondwana, including South America and China (Paris et al. 2004). The claimed diachronism for the top of the Armorican Quartzite facies is virtually nonexistent in all well-known continuous sections in northwest Europe, where, after re-examination of places where the formation was claimed to reach progressively higher horizons—even into the Darriwilian, this diachronism cannot be proven (Sá et al., 2003). The paleontological attribution of the basal parts of the Armorican Quartzite facies to the Tremadocian (e.g., Bouyx, 1970; Walter, 1982) was based on brachiopod assemblages that also lack chronostratigraphical value. Finally, using Ordovician event-stratigraphy correlation criteria, San José et al. (1992) considered a probable lower Arenigian age appropriate for the red beds located immediately below the “Armorican Quartzite” in the Central Iberian zone. ANALYTICAL TECHNIQUES Mineralogical characterization of the studied K-bentonite was performed by X-ray diffraction (XRD) using a Siemens™ D 500 XRD diffractometer with Cu Kα radiation and a graphite monochromator. The samples used were a random-powder specimen and an oriented aggregate obtained by sedimentation of the